Research Collection

Doctoral Thesis

Response of terrestrial paleoenvironments to past changes inclimate and carbon-cycling: Insights from palynology and stableisotope geochemistry

Author(s): Heimhofer, Ulrich

Publication Date: 2004

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

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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,

charcoal and bulk Corg was measured. The distinct δ13C 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

Holzresten, Blatt-Kutikulen und Gesamt-Corg ermöglichen die Erstellung einer δ13C 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).

-60ºE -30ºE-45ºE -15ºE-75ºE 0ºE

0ºN

30ºN

15ºN

-15ºN

Northern Gondwana province(arid to semi-arid)

3

2

1

transitional zone

Southern Laurasian province(subtropical to warm-temperate)

proto North Atlantic

western Tethys

EAG

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).

Fig. 2: Absolute species diversity of

Cretaceous macrofossil plant assemblages

(redrawn from Lidgard and Crane 1988).

Note the dramatic increase in the number of

angiosperm taxa from the Albian onwards.

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

600

500

400

300

200

100

0

160 140 120 100 80 60

U Jur Neocom Ba-Ap Alb Ce T-S Cmp Ma Pal

Angiosperms

Pteridophytes

Cycadales

Conifers

Ginkgoales

nu

mb

er

of sp

ecie

s

time (Ma)

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

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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

estimate changes in palaeoatmospheric CO2 levels across the OAE 1a, the δ13C 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

of organic (Corg) and carbonate carbon (Ccarb). The resulting δ13C 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).

Castellane

Mediterra

nean Sea

Grenoble

Die

Nice

Digne

Internal zones

of the Alps

Massif

Central

Provence Platform

Valence

0 50 km

Drowned platform faciesShallow open marine environmentsDeep open marine environmentspre-Triassic basement

Serre Chaitieusection

N

B

Spain

France

B

AIntermediate floral belt

Southern Laurasian province

(subtropical to

warm-te

mperate)

Northern Gondwana province(arid to semi-arid)

Serre Chaitieu

Roter Sattel

Cismon

40°

30°

20°

10°N

0 500 km

C

Luz

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

DeltaPlus 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

Mutterlose, 2003). (3) The lack of covariance between δ13Ccarb and δ18Ocarb (r2 = 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

6.1. Chemostratigraphic correlation of the δ13Ccarb 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

stable δ13Ccarb values of ~ 3.0 ‰ with the lowest value of 2.4 ‰ occurring at the base of the

studied interval. Within the first paper-shale horizon (PS-1), the δ13Ccarb signature shifts

Chapter 2

25

towards higher values of ~ 3.5 ‰. At 7.8 m the δ13Ccarb curve displays a second positive shift

and peaks at values of ~ 4.5 ‰ at 9.4 m within PS-5 (Fig. 2).

The distinct shifts in the δ13Ccarb 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

Serre Chaitieu record the stable δ13Ccarb 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

same absolute δ13Ccarb 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

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).

0

2

4

6

8

10

12

mete

rs

C7

C6

C5

C3

C4

C2

C1

δ13Ccarb

Cismon section(Menegatti et al. 1998)

Serre Chaitieu section(this study)

C5

C6

C7

C4

δ13Ccarb

Lo

we

r A

pti

an

Sta

ge

Fo

ram

.-zo

ne

Na

nn

o.-

zo

ne

Le

up

old

ina

ca

bri

R.

an

gu

stu

s (

NC

7A

)C

. li

tte

rari

us

(N

C6

)Lo

we

r A

pti

an

Le

up

old

ina

ca

bri

Glo

big

eri

nello

ides b

low

i

mete

rs

-258

-263

-268

-273

-278

Sta

ge

Fo

ram

.-zo

ne

Na

nn

o.-

zo

ne

limestone

marly limestone

marl

"black shale"

lamination

1 2 3 4 5

1 2 3 4 5

Liv

ello

Se

lli I

nte

rva

l

Niv

ea

u G

og

ue

l In

terv

al

PS-1

PS-2

PS-3

PS-4

PS-5

PS-6

Ma

ioli

ca

Sc

ag

lia

Va

rie

ga

taF

orm

atio

n

Ma

rne

s B

leu

es

Fo

rma

tio

n

Lith

olo

gy

Lith

olo

gy

C.

litt

era

riu

s (

NC

6)

R.

an

gu

stu

s (

NC

7)

section notaccessible

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

The carbon isotopic composition of the bulk OM displays a significant shift in δ13Corg from

mean values of ~ -25.5 ‰ in the lower part to values of ~ -23.8 ‰ prevailing in the upper part

of the section (Fig. 3). The increase in δ13Corg towards higher values has several superimposed

smaller-scale fluctuations (up to ~ 0.5 ‰). In comparison to the δ13Ccarb record, the bulk

δ13Corg 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

(Gelpi et al., 1970) and display a similar carbon isotopic shift as the bulk δ13Corg 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

for the latter. Due to the low abundance of C28 steranes in the upper part, δ13C values could

not be determined. The intermediate n-alkanes (n-C23, n-C24) cannot be assigned to a specific

marine source but again parallel the bulk δ13Corg pattern with a depletion in 13C of ~ 4.0 ‰.

In summary, the δ13C 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

comparison to δ13Ccarb, the biomarker record displays an increased overall shift of ~ 2.5 to 3.0

‰ (δ13Ccarb = ~ 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

composition of the carbon source (δ13C 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

a reduction in εp values. This εp decrease should result in a shift towards higher δ13C 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.

The ~1.5 ‰ positive shift in δ13Ccarb 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.,

2001). Hence, the δ13Ccarb 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

30°C, the isotopic composition of oceanic dissolved CO2 can be calculated from δ13Ccarb

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

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

δ13C 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.

02468

10

12

Meters

OAE 1a Interval

Lithology

δ13C

ste

roid

+ n

-alk

.

n-C

24

n-C

23

C28 s

tero

id

-32

.0-3

0.0

-28

.0-2

6.0

TO

C (

wt %

)

0.0

1.0

2.0

3.0

n-C

18

n-C

17

δ13C

n-a

lk (

sh

ort

-ch

ain

)

-32

.0-3

0.0

-28

.0-2

6.0

δ13C

bu

lkO

M

-27

.0-2

5.0

-23

.0

C5

C6C7

C4

12

34

5

δ13C

ca

rb

15

.01

9.0

23

.0

ε p

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

0 50 100 0 50 1000 50 100

Liv

ello

Se

lli I

nte

rva

l

bisaccate

pollen

Classopollis

% total sporomorphs

pteridophyte

spores

0 50 100 0 50 1000 50 100

Lo

we

r A

pti

an

Leu

po

ldin

a c

ab

ri

-258

-263

-268

0

2

4

6

8

10

12

mete

rs

Lo

wer

Ap

tian

Sta

ge

Fora

m.-

zone

Nanno.-

zone

Le

up

old

ina

ca

bri

R.

an

gu

stu

s (

NC

7A

)C

. li

tte

rari

us

(N

C6

)

section not

accessible

Lith

olo

gy

bisaccate

pollen

Classopollis

G. b

low

i

% total sporomorphs

Cismon section(Hochuli et al. 1999)

mete

rs

Sta

ge

Fo

ram

.-zo

ne

Na

nn

o.-

zo

ne

Lith

olo

gy

Interval of selective sporomorph preservation

pteridophyte

spores

R.

an

gu

stu

s (

NC

7)

C.

litt

era

riu

s (

NC

6)

OA

E 1

a

Serre Chaitieu section(this study)

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.

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38

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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

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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

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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

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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

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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

Southern Laurasian province(subtropical to warm-temperate)

Northern Gondwana province(arid to semi-arid)

Vocontian Basin

40°

30°

20°

10°N

0 500 km

C

Castellane

Mediterra

nean Sea

Grenoble

Die

Nice

Digne

Internal zones

of the Alps

Massif

Central

Provence Platform

Valence

0 50 km

Serre Chaitieusection

N

BDrowned platform faciesShallow open marine environmentsDeep open marine environmentspre-Triassic basementGravity reworked siliciclastics

Spain

France

B

A

Intermediate floral belt

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

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

NiveauGoguel

Lith

olo

gy

Mete

rs

Na

nn

ofo

ssil

Zo

ne

Fo

ram

inife

ral Z

on

e

Am

mo

nite

Zo

ne

Substa

ge

T. b.

P. n

.

Faisceau Nolan

Niveau Fallot

Niveau Noir

Niveau Clairs

Niveau Noir Calcaire

Niveau Blanc

black shale

marly limestone

dark-grey marlstone

studiedinterval

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

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

R2 = 0.66

n = 16

Phytoclasts (grains/mg sed.)

Po

llen

(g

rain

s/m

g s

ed

.)

(A)

20 40 60 80 100

10.0

8.0

6.0

4.0

2.0

0.0

0S

po

rom

orp

hs (

gra

ins/m

g s

ed

.)Dinoflagellate cysts (grains/mg sed.)

40

20

0

10

30

0 10 20 30 40 50 60 70 80

R2 = 0.86

n = 16

(B)

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

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.

02468

10

12

Meters

Lithology

OAE 1aP

S-3

PS

-2

PS

-1

PS

-5

PS

-6

PS

-4

Ca

CO

3 (

wt %

)

TO

C (

wt %

)

01

23

01

02

03

04

0

(A)

ca

lca

reo

us n

an

no

fossil

nutr

ient in

dex (

NI)

40

20

60

0

hig

hlo

w

(D)

(gra

ins/g

se

d.)

pa

lyn

om

orp

hs a

bso

lute

ab

un

da

nce

2 x

10

4

ma

rin

e d

ino

cysts

po

llen

an

d s

po

res

0

(C)

40

20

60

80

01

00

(% o

f to

tal p

aly

no

mo

rph

s)

(B)

ma

rin

e d

ino

cysts

po

llen

an

d s

po

res

fora

min

ife

ra lin

ing

s

pa

lyn

om

orp

hs r

ela

tive

ab

un

da

nce

4 x

10

46

x 1

04

8 x

10

4

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

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

Bis

accate

Polle

n

Vitre

isporite

s p

alli

dus

Ara

uca

ria

cite

s s

pp.

Cla

ssopolli

s s

pp.

Ephedripites s

pp.

Exe

sip

olle

nite

s s

pp.

Ina

pe

rtu

rop

olle

nite

s s

pp.

Scia

do

pitysp

olle

nite

s s

pp.

Oth

er

Gym

nosperm

s

Afr

opolli

s g

rou

p

An

gio

sp

erm

s

Cic

atr

icosis

porite

s s

pp.

Deltoid

ospora

spp.

Gle

ich

en

iidite

s s

pp.

Retitr

ilete

s s

pp.

Oth

er

Spore

s

0

2

4

6

8

10

12

Me

ters

Lo

wer

Ap

tian

Sta

ge

Fora

min

ifera

l Z

one

Nannofo

ssil Z

one

Leu

po

ldin

a c

ab

ri R. an

gu

stu

s (

NC

7A

)C

. litt

era

riu

s (

NC

6)

Lith

olo

gy

0 10 20 30 40 50%

PS-3

PS-2

PS-1

PS-5

PS-6

PS-4

OA

E 1

a

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

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.

02468

10

12

Meters

Lower AptianStage

Foraminiferal Zone

Nannofossil Zone

Leupoldina cabri

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

Lithology

PS

-3

PS

-2

PS

-1

PS

-5

PS

-6

PS

-4

Achomosphaera spp.Callaiosphaeridium asymmetricumCassiculosphaeridia reticulataCerbia tabulataCometodinium spp.Cribroperidinium spp.Dingodinium spp.Gonyaulacysta helicoideaOdontochitina operculataOligosphaeridium complexPseudoceratium securigerumPterodinium cingulatumSpiniferites spp.Subtilisphaera spp.Systematophora spp.Tanyosphaeridium variecalamumCirculodinium spp.Florentinia spp.Heslertonia heslertonensisOligosphaeridium asterigerumPinocchiodinium erbaeTrichodinium spp.Aptea polymorphaCoronifera oceanicaCribroperidinium orthocerasDruggidium apicopaucicumDruggidium spp.Gardodinium trabeculosumHystrichosphaerina schindewolfiiKiokansium polypesOligosphaeridium spp.Pterodinium spp.Batiacasphaera spp.Florentina deaneiGardodinium spp.Gonyaulacysta cretaceaKleithriasphaeridium simplicispinumPalaeoperidinium cretaceumChlamydophorella spp.Cleistosphaeridium spp.Exochosphaeridium spp.Hystrichodinium pulchrumKleithriasphaeridium spp.Exochosphaeridium phragmitesHystrichosphaeropsis spp.Rhynchodiniopsis aptianaCallaiosphaeridium spp.Odontochitina

cf. imparilis

Wallodinium lunumDapsilidinium spp.Coronifera spp.Dingodinium cerviculumKalyptea spp.Pseudoceratium spp.Sepispinula huguoniotiiKleithriasphaeridium fasciatumProlixosphaeridium parvispinumChytroeisphaeridia spp.Kleithriasphaeridium loffrenseMicrodinium opacumProlixosphaeridium spp.

Din

ofla

ge

llate

div

ers

ity

(# ta

xa

)

10

20

30

0

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).

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.

Sedimentation rate (cm ka-1)

TO

C (

wt.

%)

R2 = 0.61

n = 16

0.0 0.5 1.0 1.5 2.0 2.5 3.0 3.5

0.0

0.5

1.0

1.5

2.0

2.5

3.0

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.

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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.

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Chapter 3

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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-

plant particles. Correlation of two terrestrial δ13Cbulk 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.

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;

Algarvebasin

LagosBurgau

Sagres

PortimãoN

0 10 km

Al

ga r v e

Luz sectionBurgau section

Iberia

8°30'W9°W

37°N

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-

isotope composition of bulk OM. To avoid possible diagenetic alteration effects of the δ13Cbulk

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

low Tmax values (mean of 424.5°C). The absence of any significant correlation between δ13C

values and CaCO3- or TOC-content of the samples indicates the independence of the carbon-

isotope signature from lithological variations.

Chapter 4

78

Fig. 2: δ13Corg measurements of different types of land-plant particles (closed symbols) and the

corresponding bulk OM signature (open symbols) from the same horizon. Error bars of δ13Corg values

represent standard errors of means of repeated measurements.

5. Intrabasinal chemostratigraphic correlation

In order to test the terrestrial δ13Cbulk data for its consistency as well as for its potential as a

chemostratigraphic correlation tool, the carbon-isotope records of the Luz and Burgau

sections have been compared in detail (Fig. 3). Even though the δ13Corg record of the Luz

section is rather noisy and records an overall variation of ~8.5‰, pronounced shifts in the

magnitude of 5.0‰ to 7.0‰ can be observed. The most significant features of the δ13Cbulk

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

30

50

70

90

110

130

150

170

190

210

charcoal lignite cuticules leafs

-19-21-23-25-27 -19-21-23-25-27 -19-21-23-25-27 -19-21-23-25-27

Lo

we

r A

lbia

nU

pp

er A

ptia

nL

ow

er A

ptia

n

13Corg (%0 VPDB) δ 13Corg (%0 VPDB) δ 13Corg (%0 VPDB) δ 13Corg (%0 VPDB) δ

Lu

z M

arls

Po

rto

de

Mo

s F

m.

phytoclasts

bulk

height (m)

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

of the two records is furthermore supported by the similarity of the averaged δ13Corg 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

compared to an existing terrestrial δ13Cwood curve (Gröcke et al., 1999) and to a marine

δ13Ccarb 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

Fig. 3: Palynostratigraphy, lithological logs and terrestrial δ13Cbulk 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

Burgau

Luz

limestonemarl

silt

claystone (<25% CaCO3)

-28 -26 -24 -22 -20 -18

13C bulk OM (‰ VPDB) δ

-28 -26 -24 -22 -20 -18

-28 -26 -24 -22 -20

-28 -26 -24 -22 -20

0 m0

20

4040

60

80

90

70

505

30

10

100

0

20

40

60

80

90

70

50

30

10

100

120

140

150

130

110

Alb

ian

Low

er A

pti

anU

pp

er A

pti

an

13C bulk OM (‰ VPDB) δ

Low

er L

uz

Mar

lsU

pp

er L

uz

Mar

ls

h

eig

ht (m

)

sta

ge

form

atio

n ~ 6.5 km

E WAlgarve basin

Hys

tric

hosp

haer

idiu

m a

rbori

spin

um

Pro

toel

lipso

din

ium

coro

llum

Din

opte

rygiu

m c

ladoid

esM

uder

ongia

pari

ata

Subti

lisp

haer

a c

hei

tC

yclo

nep

hel

ium

pauci

marg

inatu

mM

uder

ongia

sta

uro

taO

donto

chit

ina a

nca

laTeh

am

adin

ium

ten

uic

eras

Bre

nner

ipoll

is r

etic

ula

tus

Afr

opoll

is aff

. ja

rdin

us

Afr

opoll

is j

ard

inus

Bre

nner

ipoll

is p

erore

ticu

latu

sB

renner

ipoll

is cf.

div

idus

Tri

colp

its

vulg

ari

sD

ichast

opoll

enit

es sp

.

Din

ocysts

Sp

ore

s

& P

olle

n

correlation of marker beds

I

II

III

IV

V

VI

VII

VIII

IX

X

ch

em

ostr

at.

se

gm

en

t

Chapter 4

81

Fig. 4: Tentative correlation of mid-Cretaceous terrestrial δ13Corg 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

δ13Ccarb 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

Isle of Wight

(UK)

-18-22-26-30

150

100

50

0

200

250

300

80

60

40

0

100

120

20

140

Algarve

(Portugal)

160

he

igh

t (m

)

sta

ge

he

igh

t (m

)

13Cwood (‰ VPDB) δ 13Cbulk OM (‰ VPDB) δ

sta

ge

Lo

we

r A

ptia

nU

pp

er A

ptia

nA

lb.

2.0 3.0 4.0 5.0

Barr.

Lower

Aptian

Upper

Aptian

Albian

13Ccarb (‰ VPDB) δ

-20

Composite

marine recordN

C7

NC

8N

C6

NC

5

L. cabri

G. blo

wi

G. fe

rr.

G. alg

.

?

Lo

we

r A

ptia

nU

pp

er A

ptia

nA

lb.

Barr

.

H. pla

n. -

T

. bej.

sta

ge

Pk F

ora

m. Z

one

Na

nn

ofo

s. Z

on

e

-28 -26 -22-24

jacobi

nutfield-iensis

martinioides

forbesi

deshayesi

bowerbanki

fissicostatus

Am

monite

Zonation

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

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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).

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.

-9°45' -9°15' -8°15'

40°30'

40°00'

39°30'

39°00'

38°30'

25 km

Figueira da FozCoimbra

Nazaré

Peniche

Cresminasection

Ericeira

São Juliãosection

Low

er

Tagus

Basi

n

Santa Cruz

Buarcos

Famalicão

Vale de Agua

Torres Vedras

Catefica

N

Palaeozoic Basement

Angiosperm mesofossil site

Naz

aré

Fault Zon

e

Faro

Silves

Lagos

25 km

-9°00' -8°30' -8°00' -7°30'

-37°00'Luz

section

Tavira

8°

IB

ER I A

Port

ugal

A

B

C

B

C

-8°45'

Cascais

Lisboa

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

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.

Lusitanian Basin

Tehamadinium tenuiceras

Muderongia staurota

Hystrichosphaerina schindewolfii

Dinopterygium cladoides

Ctenidodinium elegantulum

Pseudoceratium securigerum

Callaiosphaeridium trycherium

Rhynchodiniopsis aptiana

Pseudoceratium pelliferum

Pseudoceratium securigerum

Callaiosphaeridium trycherium

Odontochitina operculata

Palaeoperidinium cretaceum

Heslertonia heslertonensis

Ctenidodinium elegantulum

Dinopterygium cladoides(consistent occurrence)

Xiphophoridium alatum

Chichaouadinium vestitum

60

40

20

0

80

100

120

140

160

180

200

220

240

260

Lo

we

r B

ed

ou

lian

Up

pe

r B

ed

ou

lian

Up

pe

r A

ptia

n L

ow

er A

lbia

n

Po

rto

de

Mo

s F

m.

Up

pe

r L

uz M

arls

Lo

we

r L

uz M

arls

Pa

lorb

ito

lina

Be

ds

Algarve Basin

60

40

20

0

80

100

120

140

160

180

200

Lo

we

r A

lbia

nU

pp

er

Ba

rre

mia

nM

idd

le A

lbia

nL

ow

er

Be

do

ulie

n

Ro

dis

io

Fm

.

Ga

lé F

m.

Cre

sm

ina

Fm

.

Sta

ge

Me

ter

Fo

rma

tio

n

Lith

olo

gy

Sta

ge

Me

ter

Fo

rma

tio

n

Lith

olo

gy

Subtilisphaera perlucida

Cerbia tabulata

Ag

ua

Do

ce

Mb

.P

on

ta A

lta

Co

bre

Mb

.P

dL

Me

mb

er

C. decip

iens

Ma

rls

Dinopterygium cladoides(consistent occurrence)

Pachytraga paradoxa

Caprina douvilleiPraecaprina varians rudists

limestone

siltstone

sandstone/conglomerate

marl

claystone (< 25% CaCO3)

first occurrence (FO)

last occurrence (FO)

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

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

-60ºE -30ºE-45ºE -15ºE-75ºE 0ºE

0ºN

30ºN

15ºN

-15ºN

Northern Gondwana province(arid to semi-arid)

3

4

2

1

5

transitional zone

6

7

8

9

Southern Laurasian province(subtropical to warm-temperate)

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.

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.

Odonto

chitin

a a

ncala

Mic

rod

iniu

m s

pp.

Ch

ich

ao

ua

din

ium

spp.

Ce

pa

din

ium

ve

ntr

iosa

Ch

lam

yd

op

ho

rella

spp.

Cle

isto

sp

ha

erid

ium

spp.

Ta

nyo

sp

ha

erid

ium

va

rie

ca

lam

um

Po

lysp

ha

erid

ium

spp.

Dis

silio

din

ium

spp.

Dic

on

od

iniu

m s

pp.

Cyclo

ne

ph

eliu

m p

au

cis

pin

um

Xip

ho

ph

orid

ium

ala

tum

Vo

zzh

en

nik

ovia

spp.

Circu

lod

iniu

m s

pp.

Ch

ich

ao

ua

din

ium

ve

stitu

mA

sco

din

ium

spp.

Cyclo

ne

ph

eliu

m v

an

no

ph

oru

mO

donto

chitin

a s

pp.

Od

on

toch

itin

a im

pa

rilis

Da

psilid

iniu

m w

arr

en

iiE

xo

ch

osp

ha

erid

ium

ph

rag

mite

sF

lore

ntin

a d

ea

ne

iK

aly

pte

a s

pp.

Pin

occh

iod

iniu

m e

rba

eT

rich

od

iniu

m c

asta

ne

a

Su

btilisp

ha

era

ch

eit

+

Pro

toe

llip

so

din

ium

co

rollu

m

++

Olig

osp

ha

erid

ium

to

tum

Olig

osp

ha

erid

ium

pu

lch

err

imu

mD

ino

pte

ryg

ium

cla

do

ide

sD

ap

silid

iniu

m d

efla

nd

rei

Ge

ise

lod

iniu

m s

pp.

Cte

nid

od

iniu

m e

leg

an

tulu

m

Pse

ud

oce

ratiu

m p

ellife

rum

Ca

lla

iosp

ha

erid

ium

spp.

++

Pte

rod

iniu

m s

pp.

Kle

ith

ria

sp

ha

erid

ium

co

rru

ga

tum

Ach

om

osp

ha

era

spp.

+

+

Ba

tia

ca

sp

ha

era

spp.

+

Crib

rop

erid

iniu

m o

rth

oce

ras

+

Exo

ch

osp

ha

erid

ium

spp.

Flo

rentinia

spp.

+

Hystr

ich

od

iniu

m p

ulc

hru

m

+

Hystr

ich

osp

ha

erin

a s

ch

ind

ew

olfii

+

Te

ha

ma

din

ium

spp.

+

Ca

lla

iosp

ha

erid

ium

asym

me

tric

um

+

+

Din

go

din

ium

spp.

He

sle

rto

nia

he

sle

rto

ne

nsis

Hystr

ich

osp

ha

erin

a s

pp.

Mic

rod

iniu

m o

pa

cu

mA

pte

a p

oly

mo

rph

a

+

+

Co

me

tod

iniu

m s

pp.

Mu

de

ron

gia

pa

ria

taP

se

ud

oce

ratiu

m s

ecu

rig

eru

m "

nu

du

m"

Trich

od

iniu

m s

pp.

Su

btilisp

ha

era

spp.

++

++

Sp

inife

rite

s s

pp.

++

Olig

osp

ha

erid

ium

spp.

Olig

osp

ha

erid

ium

co

mp

lex

+

++

Mu

de

ron

gia

spp.

Hystr

ich

od

iniu

m fu

rca

tum

Go

nya

ula

cysta

cre

tace

aC

rib

rop

erid

iniu

m s

pp.

+

Circu

lod

iniu

m b

revis

pin

osu

m

++

++

Su

btilisp

ha

era

pe

rlu

cid

aP

se

ud

oce

ratiu

m s

ecu

rig

eru

mP

ala

eo

pe

rid

iniu

m c

reta

ce

um

+

Olig

osp

ha

erid

ium

aste

rig

eru

m

+

Od

on

toch

itin

a o

pe

rcu

lata

+

+

Mu

de

ron

gia

sta

uro

taK

leith

ria

sp

ha

erid

ium

spp.

Kle

ith

ria

sp

ha

erid

ium

re

ad

ei

Kio

ka

nsiu

m p

oly

pe

s

++

Ellip

so

idic

tyu

m s

pp.

Crib

rop

erid

iniu

m e

dw

ard

sii

Ce

rbia

ta

bu

lata

Ca

lla

iosp

ha

erid

ium

try

ch

eriu

mA

ch

om

osp

ha

era

ne

ptu

ni

Lo

we

r A

lbia

nU

pp

er

Ba

rre

mia

nM

idd

le A

lbia

nLo

wer

Bed

oulie

n

Ro

dis

io

Fm

.

Ga

lé F

m.

Cre

sm

ina

Fm

.

Sta

ge

Fo

rma

tio

n

Ag

ua

Do

ce

Mb

.P

onta

Alta

Co

bre

Mb

.P

dLM

em

be

r

Sa

mp

le

Me

ter

60

40

20

0

80

100

120

140

160

180

L-97

L-91L-88

L-69

L-66

L-60

L-55

L-52

L-48L-43

L-40

L-37L-31

L-19L-16L-13

L-5

L-1

K-3.1

K-2.1

E-1sE-6E-3E-1fE-1c

D-62

D-13

D-7

L-25

200

(A)

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

60

40

20

0

80

100

120

140

160

180

200

220

240

260

Mete

r

+

+

Cle

isto

sphaeridiu

m s

pp.

+

+

+

+

+

+

Cyclo

nephelium

vannophoru

m

+

+

+

+

+

+

Hystr

ich

osp

ha

erid

ium

arb

orisp

inu

m

+

+

++

Pte

rod

iniu

m s

pp.

+

+

Oligosphaeridiu

m s

pp.

+

+

+

+

+

+

+

+

+

Pro

toe

llip

so

din

ium

co

rollu

m

+

+

+

+

+

+

+

+

+

+

+

+

++

+

+

+

++

+

Co

me

tod

iniu

m s

pp.

+

+

+

+

Ce

rbia

ta

bu

lata

+

Din

go

din

ium

alb

ert

ii

+ +

Flo

ren

tin

a d

ea

ne

i

+

+

+

+

+

Hystr

ich

osp

ha

erin

a s

ch

ind

ew

olfii

+ +

Pa

lae

op

erid

iniu

m c

reta

ce

um

+

+

+

+

+

+

+

+

+

+

Spin

iferite

s s

pp.

+

+

+

+

+

+

+

+

+

Trich

od

iniu

m s

pp.

+

+

+

Subtilisphaera

spp.

+

+

+

+

+

+

+

+

+

+

+

+

+

+

+

+

+

++

Su

btilisp

ha

era

pe

rlu

cid

a

+

+

Rh

yn

ch

od

inio

psis

ap

tia

na

+

+

+

+

Pseudocera

tium

securigeru

m

+

+

Oligosphaeridiu

m c

om

ple

x

+

+

+

+

+

+

+

+

+

+

+

+

+

+

+

Kio

ka

nsiu

m p

oly

pe

s+

+

+

+

+

+

+

+

+

+

Hystr

ich

od

iniu

m p

ulc

hru

m

+

+

+

Cte

nid

od

iniu

m e

leg

an

tulu

m

+

+

Circu

lod

iniu

m b

revis

pin

osu

m+

++

+

+

+

+

+

+

+

+

+

+

+

+

+

Ce

pa

din

ium

ve

ntr

iosa

+

+

++

+

+

+

+

+

+

+

+

+

+

++

Ca

lla

iosp

ha

erid

ium

try

ch

eriu

m

++

Callaio

sphaeridiu

m s

pp.

+

+

+

+

+

+

+

Ca

lla

iosp

ha

erid

ium

asym

me

tric

um

+

+

+

Batiacasphaera

spp.

+

+

+

+

+

+

+

+

+

+

++

+

+

+

+

+

Achom

osphaera

spp.

+

+

+

+

+

+

+

+

+

+

+

+

+

+

+

Hystr

ichosphaeridiu

m s

pp.

+

Co

ron

ife

ra s

pp.

+

Dis

silio

din

ium

spp.

+

Apte

a p

oly

morp

ha

+A

pte

od

iniu

m s

pp.

+

Hete

rosphaeridiu

m s

pp.

+

Olig

osp

ha

erid

ium

aste

rig

eru

m

+

Ce

pa

din

ium

spp.

+

Ovo

idin

ium

spp.

+

Mudero

ngia

sta

uro

ta

+

+

+

Te

ha

ma

din

ium

te

nu

ice

ras

+

+

Odonto

chitin

a o

perc

ula

ta

+

+

Co

ron

ife

ra o

ce

an

ica

+

Exig

uis

phaera

spp.

+

Syste

mato

phora

spp.

+

Su

btilisp

ha

era

pirn

ae

nsis

+

++

Ca

nn

ing

ia s

pp.

+

+

+

Od

on

toch

itin

a s

pp.

+

+

+

+

Chla

mydophore

lla s

pp.

+

+

+

+

+

Od

on

toch

itin

a a

nca

la

Din

go

din

ium

spp.

+

+

Mic

rod

iniu

m s

pp.

+

+

+

+

Cyclo

ne

ph

eliu

m p

au

cim

arg

ina

tum

+

+

+

Su

btilisp

ha

era

se

ne

ga

len

sis

+

+

+

+

+

+

+

Va

lva

eo

din

ium

spp.

+

+

+

+

+

+

+

++

+

+

+

+

Te

ha

ma

din

ium

spp.

+

+

+

+

+

+

+

Crib

rop

erid

iniu

m s

pp.

+

+

+

+

+

+

+

+

Kaly

pte

a s

pp.

+

+

+

+

+

Kle

ithriasphaeridiu

m s

pp.

+

+

+

Crib

rop

erid

iniu

m e

dw

ard

sii

+

Cyclo

ne

ph

eliu

m p

au

cis

pin

um

+

+

++

+

+

Chytr

oeis

phaeridia

spp.

+

Mu

de

ron

gia

spp.

+

Trich

od

iniu

m c

asta

ne

a

+

Oligosphaeridiu

m totu

m

+

+

+

+

+

Exochosphaeridiu

m s

pp.

+

+

Ge

ise

lod

iniu

m s

pp.

+

++

+

+

+

+

Olig

osp

ha

erid

ium

pu

lch

err

imu

m

+

+

+

+

+

Ellip

so

idic

tyu

m s

pp.

+

Ch

ich

ao

ua

din

ium

spp.

+

+

+

+

Mudero

ngia

pariata

+

Subtilisphaera

cheit

+

++

+

+

+

+

+

+

+

+

+

+

+

+

+

+

+

++

Circu

lod

iniu

m s

pp.

+

+

+

+

Din

opte

rygiu

m c

ladoid

es

Low

er

Bedoulia

nU

pp

er

Be

do

ulia

n U

pp

er A

ptia

n L

ow

er A

lbia

n

Port

o d

e M

os F

m.

Up

pe

r L

uz M

arls

Lo

we

r L

uz M

arls

Pal

orbi

tolin

aB

eds

Sta

ge

Form

ation

C. d

ecip

iens

Mar

ls

A-201

A-196

A-194A-193

A-188

A-179

A-176

A-172

A-169

A-162

A-154

A-148

A-137A-134

A-121

A-115

A-112

A-108A-106

A-101

A-94

A-79

A-59

A-46A-41

A-33

B-13

B-8

A-125

A-114

A-110

A-97

A-81

A-37

Sa

mp

le(B)

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.

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.

Stage Rey (1992) Berthou and

Leereveld (1990)

This study

Lower

Bedoulian

Upper

BedoulianLower

Aptian

Upper Barremian

Upper Aptian

Lower Albian

Middle Albian

Upper Albian

?

Co

bre

Mb

.

Po

nta

Alta

Mb

.

Pra

ia d

a L

ag

oa

Mb

.

Ro

diz

io F

m.

Ag

ua

Do

ce

Mb

.

Co

bre

Mb

.

Po

nta

Alta

Mb

.

Pra

ia d

a L

ag

oa

Mb

.

Ag

ua

Do

ce

Mb

.

Co

bre

Mb

.

Po

nta

Alta

Mb

.

Pra

ia d

a L

ag

oa

Mb

.

Ro

diz

io F

m.

Ag

ua

Do

ce

?

?

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.

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.

Stage Rey (1986) Berthou and

Leereveld (1990)

This study

Lower

Bedoulian

Upper

BedoulianLower

Aptian

Upper Barremian

Upper Aptian

Lower Albian

Middle Albian

C. depie

ns M

arls

Pa

lorb

iolin

a B

ed

s

Luz M

arls F

m.

Po

rto

de

Mo

s F

m.

C. depie

ns M

arls

Pa

lorb

iolin

a B

ed

s

Luz M

alrs F

m.

Port

o d

e M

os F

m.

? ?

Luz M

alrs F

m. &

Po

rto

de

Mo

s F

m.

?

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.

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.

Po

do

ca

rpid

ite

s s

pp

.

Eu

co

mm

iid

ite

s s

pp

.

Ve

rru

co

sis

po

rite

s s

pp

.C

on

ve

rru

co

sis

po

rite

s s

pp

.

Ha

mu

latisp

orite

s s

pp

.Is

ch

yo

sp

orite

s s

pp

.

Fo

ve

osp

orite

s s

pp

.C

ing

utr

ile

tes s

pp

.E

ch

ina

tisp

oris s

pp

.

Klu

kis

po

rite

s s

pp

.

Re

ticu

latisp

orite

s s

pp

.

Co

nca

vis

po

rite

s s

pp

.C

osta

top

erf

oro

sp

orite

s s

pp

.Im

pa

rde

cis

po

ra s

pp

.

Cic

atr

ico

sis

po

rite

s s

pp

.

Plica

tella

sp

p.

Ru

bin

ella

sp

p.

Trile

te s

po

res in

de

t.

Sta

ge

60

40

20

0

80

100

120

140

160

180

200

Mete

r

Lith

olo

gy

Gymnosperm pollen Trilete sporesAngiosperm

pollen

Ca

llia

lasp

orite

s s

pp

.

Vitre

isp

orite

s p

allid

us

Scia

do

pitysp

olle

nite

s s

pp

.E

ph

ed

rip

ite

s s

pp

.

De

lto

ido

sp

ora

sp

p.

Gle

ich

en

iid

ite

s s

pp

.L

ep

tole

pid

ite

s s

pp

.

Bire

tisp

orite

s s

pp

.C

on

ca

vis

sim

isp

orite

s s

pp

.

Bis

accate

Polle

n

Cla

sso

po

llis s

pp.

Ara

ucariacites s

pp.

% T

rile

te s

po

res

% A

ng

iosp

erm

s

Exe

sip

olle

nte

s s

pp.

Ina

pe

rtu

rop

oll.

spp.

Pe

rin

op

olle

nite

s s

pp.

% A

fro

po

llis s

pp.

0 100 %

barren interval

Sa

mp

le

LP

Z

I

II

III

IV

MU

L-97

L-91L-88

L-69

L-66

L-60

L-55

L-52

L-48

L-43

L-40

L-37

L-31

L-19L-16L-13

L-5

L-1

K-3.1

K-2.1

E-1sE-6E-3E-1fE-1c

D-62

D-13

D-7

L-25

Lo

we

r A

lbia

nU

pp

er

Ba

rre

mia

nM

iddle

Alb

ian

Lo

we

rB

ed

ou

lien

Ro

dis

io

Galé

Fm

.C

resm

ina F

m.

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).

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.

MU

barren interval

monocots and magnoliidsLusitanianBasin

60

40

20

0

80

100

120

140

160

180

200

Lo

we

r A

lbia

nU

pp

er

Ba

rre

mia

nM

idd

le A

lbia

nL

ow

er

Bedoulie

n

Ro

dis

io

Galé

Fm

.C

resm

ina F

m.

Sta

ge

Mete

r

Form

ation

Lith

olo

gy

+

+

+

+

+

+

+

+

Cla

va

tip

olle

nite

s s

pp

.

+

+

+

+

Cla

va

tip

olle

nite

s c

f. h

ug

he

sii

+

+

+

+

Cla

va

tip

olle

nite

s c

f. m

inu

tus

+

+

Re

tim

on

oco

lpite

s s

p. 6

+

+

+

+

+

Aste

rop

ollis

cf. a

ste

roid

es

+

Pe

nn

ipo

llis

sp. 2

+

+

Re

tim

on

oco

lpite

s s

p. 4

+

+

Aste

rop

ollis

sp

. 2

+

+

+

Dic

ha

sto

po

lle

nite

s c

f. g

ha

za

late

nsis

+

+

Ste

lla

top

ollis

sp

p.

+

+

+

Aste

rop

ollis

sp

. 4

+

+

Dic

ha

sto

po

lle

nite

s s

p. 2

+

Re

tim

on

oco

lpite

s c

f. e

xce

lsu

s

+

Re

tim

on

oco

lpite

s c

f. s

p. 11

+

+

+

+

+

+

+

Re

tim

on

oco

lpite

s s

pp

.

Aste

rop

ollis

aste

roid

es

+

Aste

rop

ollis

sp

p.

+

+

Cla

va

tip

olle

nite

s c

f. s

p. A

+

+

+

+

+

Dic

ha

sto

po

lle

nite

s s

p. 1

Ra

ce

mo

no

co

lpite

s c

f. e

xo

ticu

s

+

+

+

Re

tim

on

oco

lpite

s s

p. 3

Re

tim

on

oco

lpite

s s

p. 8

Ste

lla

top

ollis

sp.1

Cla

va

tip

olle

nite

s c

f. t

en

ellis

+

Re

tim

on

oco

lpite

s s

p. 5

+

Dic

ha

sto

po

lle

nite

s d

un

ve

ga

ne

nsis

+

Ste

lla

top

ollis

ba

rgh

oo

rnii

+

Pe

nn

ipo

llis

sp

. 3

+

+

Re

tim

on

oco

lpite

s s

p. 2

+

+

+

Dic

ha

sto

po

lle

nite

s s

p. 4

Re

tim

on

oco

lpite

s s

p. 1

0

+

Re

tim

on

oco

lpite

s s

p. 1

2

+

Re

tim

on

oco

lpite

s s

p. 1

1

+

Cla

va

tip

olle

nite

s s

p. 3

Dic

ha

sto

po

lle

nite

s c

f. s

p. 5

Re

tim

on

oco

lpite

s s

p. 1

+

Re

tim

on

oco

lpite

s s

p. 1

3

+

+

Re

tim

on

oco

lpite

s s

p. 1

6

Dic

ha

sto

po

lle

nite

s s

p. 5

Re

tim

on

oco

lpite

s s

p. 7

+

+

+

+

+

Trico

lpite

s s

pp

.

+

Se

ne

cto

tetr

ad

ite

s s

pp

.

+

+

Str

iato

po

llis t

roch

ue

nsis

+

Ste

ph

an

oco

lpite

s a

ff. fr

ed

ericksb

urg

en

sis

+

Str

iato

po

llis s

pp

.

+

Re

titr

ico

lpite

s a

ff.

ve

rmim

uru

s

Sa

mp

le

eudicots

0 25%0 30

angiosperm

diversity

relative

abundance

Dic

ha

sto

po

lle

nite

s s

p. 6

+

+

+

+

+

+

Dic

ha

sto

po

lle

nite

s s

pp

.

L-97

L-91L-88

L-69

L-66

L-60

L-55

L-52

L-48

L-43

L-40

L-37

L-31

L-19L-16L-13

L-5

L-1

K-3.1

K-2.1

E-1sE-6E-3E-1fE-1c

D-62

D-13

D-7

L-25

LP

Z

I

II

III

IV

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

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

Bis

acca

te P

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us

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lla s

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% A

ng

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Gymnosperm pollen Trilete sporesAngiospermpollen

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100

120

140

160

180

200

220

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260

Mete

r

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ple

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olo

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A-112

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A-33

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A-97

A-81

A-37

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.

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.

Aste

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angiosperm

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relative

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Sa

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A-201

A-196

A-194A-193

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A-172

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A-110

A-97

A-81

A-37

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).

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.

SSW

Lower - Middle

Albian

Low

er

Be

do

ulia

n

Lo

we

r

Ba

rre

mia

n?

Cresmina

section

Torres Vedras

section

NNE

Up

pe

r

Ju

rassic

Caranguejeira

section

Up

pe

r

Ju

rassic

Upper Albian

Cenomanian

Buarcos

section

123Rodizio

Formation

MU

Calvaria

Member

Figueira da Foz

Formation

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 circular-elliptical

columellate-tectate sexine: 0.5 nexine: 1.0

widely spaced length: 0.5

monocolpate well-defined

Pl. 1; Fig. 1-3

C. cf. minutus Brenner (1963) ~20 circular-elliptical

columellate-tectate sexine: 0.6 nexine: 0.4

densely spaced length: 0.5

not visible Pl. 1; Fig. 3-4

C. cf. tenellis Phillips & Felix (1972)

~30 irregular spherical

perforate-tectate sexine: 1.0 nexine: 1.2

very densely spaced length: 1.0

not visible Pl. 1; Fig. 8-9

C. sp. 1 informal species ~15 spherical

columellate-tectate sexine: 1.0 nexine: 1.0

distinct, widely spaced length: 1.0 head Ø: 0.5

monocolpate well-defined

Pl. 1; Fig. 7

C. sp. 2 informal species ~30 circular-elliptical

columellate-tectate sexine: 1.0 nexine: 1.0

very fine barely visible

monocolpate elongate

Pl. 1; Fig. 12-13

C. sp. 3 informal species ~40 circular-elliptical

columellate-tectate sexine: 1.0 nexine: 0.5

densely spaced length: 1.0

monocolpate Pl. 1; Fig. 10-11

C. cf. sp. A Doyle & Robbins (1975)

~22 spherical

microreticulate-tectate sexine: 1.0-1.5 nexine: 1.0

densely spaced length: 1.5-2.0

monocolpate Pl. 1 Fig. 5-6

Genus Asteropollis (HEDLUND & NORRIS) Species Author Size and shape Exine Columellae Aperture Plate

A. cf. asteroides Hedlund & Norris (1968)

20-25 circular-elliptical

columellate-tectate sexine: 1.5 nexine: < 0.5

densely spaced length: 0.5 club shaped

trichotomocolpate Pl. 3; Fig. 12-13

A. sp. 1 informal species 25-30 irregular spherical

columellate-tectate sexine: 1.0 nexine: 0.5

densely spaced club shaped length: 1.0 head Ø: ~0.5

not visible Pl. 3; Fig. 7-8

A. sp. 2 informal species ~50 irregular spherical

microreticulate-tectate verrucate tectum

very densely spaced barely visible

not visible Pl. 3; Fig. 11

A. sp. 3 informal species ~20 circular-elliptical

perforate-tectate exine: < 1.0

densely spaced barely visible

trichotomocolpate Pl. 3; Fig. 9-10

A. sp. 4 informal species ~25 circular-elliptical

microreticulate-tectate sexine: 1.0 nexine: < 0.3

densely spaced club shaped length: 0.5 - 0.8 head Ø: 0.6

trichotomocolpate Pl. 4; Fig. 2-3

Genus Pennipollis (FRIIS, PEDERSEN & CRANE) Species Author Size and shape Exine Reticulum Muri width Aperture Plate P. sp. 1 informal

species ~20 spherical

reticulate-semitectate sexine: 1.5-2.0 nexine: 0.5

lumina: 1.5-2.5 1.2-1.3 transverse ridges

monocolpate Pl. 4; Fig. 10-11

P. sp. 2 informal species

15 - 20 circular-slightly elliptical

reticulate-semitectate sexine: 0.75 nexine: 0.75

homobrochate lumina: 1.3-2.8

0.5-0.7 verrucate

monocolpate Pl. 4; Fig. 12-13

P. sp. 3 informal species

~20 spherical

reticulate-semitectate sexine: 0.7 nexine: 0.8

homobrochate lumina: 1.0-2.5

0.5-0.8 double-row verrucae

monocolpate Pl. 4; Fig. 14-15

P. sp. 4 informal species

~15 circular-slightly elliptical

reticulate-semitectate sexine: 0.5 nexine: 1.0

heterobrochate lumina: 1.5-4.0

0.5 fine ornamentation

monocolpate Pl. 4; Fig. 16

Genus Dichastopollenites (MAY) Species Author Size and

shape Exine Reticulum Muri width Columellae Aperture Plate

D. cf. ghazalatensis Ibrahim (1996)

28–31 reticulate-semitectate polygonal lumina width: 1.5-4.0

0.8-1.2 widely spaced Ø: 1.1-1.3

zono- aperturate

Pl. 4; Fig. 9

D. dunveganensis Singh (1983)

~45 reticulate-semitectate heterobrochate; polygonal lumina width: 3.5-6.5

1.0-1.3 very widely spaced Ø: 1.2-1.6

zono- aperturate

Pl. 4; Fig. 8

D. sp. 1 informal species

26–35 circular-elliptical

reticulate-semitectate sexine: ~1.5 µm nexine: 0.5 – 0.7 µm

heterobrochate; polygonal lumina width: 2.0-6.0

0.6-0.8 very widely spaced Ø: 0.8-0.9

zono- aperturate

Pl. 5; Fig. 1-2

D. sp. 2 informal species

23–27 circular-elliptical

reticulate-semitectate sexine: 1.5 – 2.0 µm nexine: 0.5 µm

heterobrochate; polygonal lumina width: 1.0-3.0

0.6-0.8 widely spaced Ø: 0.7-0.8 club shaped

zono- aperturate

Pl. 4; Fig. 4-5

D. sp. 4 informal species

45–48 circular-elliptical

reticulate-semitectate sexine: 2.5 – 3.0 µm nexine: 0.5 µm

irregular-heterobrochate incomplete meshes

0.8-1.0 widely spaced Ø: 1.1-1.3 club shaped

zono- aperturate

Pl. 5; Fig. 5-6

D. sp. 5 informal species

~31 elliptical

reticulate-semitectate sexine: 1.3 – 1.5 µm nexine: 0.5 µm

irregular-heterobrochate incomplete meshes

0.5-0.7 triangular profile

widely spaced Ø: 0.8-1.0 spindle shaped

zono- aperturate

Pl. 4; Fig. 6-7

Chapter 5 125

Genus Retimonocolpites (PIERCE) Species Author Size Exine Reticulum Muri width Columellae Aperture Plate R. cf. excelsus

Ward (1986)

~35 reticulate-semitectate nexine: 0.6-0.7

coarse reticulate irregular, loosely attached

0.6-0.8 widely spaced length: ~1.0

monocolpate Pl. 2; Fig. 11

R. sp. 1 informal species

~15 reticulate-semitectate sexine: 0.6 nexine: 0.7

heterobrochate loosely attached lumina width: 1.0-2.5

0.5 length: ~1.0 monocolpate Pl. 2; Fig. 1-2

R. sp. 2 informal species

~25 reticulate-semitectate sexine: 2.0 nexine: < 0.4

homobrochate; polygonal lumina width: 2.0-5.0

0.8 dispersed verrucae

widely spaced club shaped length: ~2.0 head Ø: 1.0

monocolpate Pl. 2; Fig. 5-6

R. sp. 3 informal species

~25 reticulate-semitectate sexine: 2.0 nexine: 0.5

heterobrochate; irregular lumina width: 2.0-3.5

0.5 few verrucae

widely spaced length: ~2.0 head Ø: ~0.8

monocolpate Pl. 2; Fig. 7-8

R. sp. 4 informal species

~22 reticulate-semitectate sexine: 1.3-1.5 nexine: 0.5

heterobrochate; polygonal lumina width: 1.2-3.0

0.2-0.3 dispersed verrucae

widely spaced length: 1.2-1.5 head Ø: ~0.5

monocolpate Pl. 3; Fig. 3-4

R. sp. 5 informal species

~30 microreticulate-tectate sexine: 1.0 nexine: 0.5

irregular microreticulate lumina width: < 0.5

very dense densely spaced length: 0.5-1.0 head Ø: ~0.5

monocolpate Pl. 2; Fig. 15

R. sp. 6 informal species

30-36 reticulate-semitectate sexine: 0.5 nexine: 0.5

heterobrochate; polygonal microreticulate lumina width: < 1.0

0.2-0.3 densely spaced barley visible

monocolpate elongate colpus

Pl. 4; Fig. 1

R. sp. 7 informal species

30-33 reticulate-semitectate sexine: 1.0-1.5 nexine: 1.0

homobrochate smaller towards colpus lumina width: 1.0-3.0

0.5-1.2 irregular; beaded

densely spaced club shaped length: 1.0 head Ø: 1.0-1.2

monocolpate long colpus

Pl. 2; Fig. 10

R. sp. 8 informal species

~15 reticulate-semitectate very thin sexine nexine: 0.5

heterobrochate loosely attached lumina width: < 1.0

< 0.3 regular

widely spaced length: 0.3

monocolpate Pl. 2; Fig. 9

R. sp. 9 informal species

~24 reticulate-semitectate sexine: 1.5 nexine: 0.8

homobrochate lumina width: 2.0-3.0

0.5 regular

widely spaced club shaped length: 1.5-2.0 head Ø: ~0.5

monocolpate Pl. 2; Fig. 12

R. sp. 10 informal species

~16 reticulate-semitectate sexine: 1.0 nexine: 0.5

extremely heterobrochate irregular lumina width: < 1.5

0.2-0.3 thin length: < 0.5

monocolpate Pl. 2; Fig. 16-17

R. sp. 11 informal species

~21 reticulate-semitectate sexine: 2.0 nexine: 0.5

heterobrochate irregular lumina width: 2.0-5.5

0.5-0.8 triangular profile

widely spaced spindle-shaped, length: < 2.0

monocolpate Pl. 2; Fig. 13-14

R. sp. 12 informal species

~22 reticulate-semitectate sexine: 1.8 nexine: 0.7

homobrochate smaller towards colpus lumina width: 1.0-2.0

0.7 widely spaced club-shaped length: 0.7-1.0

monocolpate elongate colpus

R. sp.13 informal species

~37 reticulate-semitectate sexine: 1.7 nexine: 0.3

homobrochate lumina width: < 2.0

0.3 densely spaced club shaped length: 1.7 head Ø: 0.5-0.7

monocolpate elongate colpus

Pl. 3; Fig. 3-4

R. sp. 15 informal species

~28 reticulate-semitectate sexine: 1.3 nexine: 1.2

heterobrochate loosely attached lumina width: 0.5-1.7

0.3 densely spaced, club shaped length: 1.2

monocolpate Pl. 3; Fig. 5-6

R. sp. 16 informal species

25-30 reticulate-semitectate sexine: 1.2 nexine: 0.6

homobrochate lumina width: < 1.3

< 0.3 dispersed verrucae

densely spaced club shaped length: 1.2

monocolpate elongate colpus

Pl. 3; Fig. 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

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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)

Plate I

1

2

3

4

5

6

7

11

12

9

10

13

14

8

133

Plate II

1

2

3

4

9

5

7

6

8

10

11

12

16

1713 14 15

135

Plate III

1

3

9

7

2

5

6

10 11

4

13

1412

8

137

Plate IV

8 9 16

10

12

14

11

13

15

4 5

6 7

1

2 3

139

Plate V

1 2

3

5 6

4

141

Plate VI

1

2

3

10

11

12

7

8

9

5

6

4

143

1

4

2

3

65

7 8 9

145

Plate VII

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

Sample

Hight (m)

TOC (dry wt %)

TIC (dry wt %)

δ13Cbulk OM (‰ VPDB)

δ13Ccharcoal (‰ VPDB)

δ13Clignite (‰ VPDB)

δ13Cleaves (‰ VPDB)

δ13Ccuticle (‰ 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

Sample

Hight (m)

TOC (dry wt %)

TIC (dry wt %)

δ13Cbulk OM (‰ 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

Sample

Hight (m)

TOC (dry wt %)

TIC (dry wt %)

δ13Cbulk OM (‰ 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

Sample

Hight (m)

TOC (dry wt %)

TIC (dry wt %)

δ13Ccarb (‰ VPDB)

δ13Cbulk OM (‰ 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

Sample

Hight (m)

TOC (dry wt %)

TIC (dry wt %)

δ13Ccarb (‰ VPDB)

δ13Cbulk OM (‰ 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 (m)

δ13C n- C17 stdev

δ13C n- C18 stdev

δ13C n- C23 stdev

δ13C n- C24 stdev

δ13C 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 (cm)

Phytoclasts trans. < 25

Phytoclasts trans. >25

Phytoclasts opaque <25

Phytoclasts opaque >25

Membranes

Cuticle

Sporomorphs

Dinocysts

other cysts

Forams

AOM

Sum

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

H-1

H-2H-3H-4H-5H-6H-7H-8

H-9H-10

H-11

H-12

H-13

H-14

H-15H-16

H-25

H-24

H-23H-22H-21

H-20

H-19

H-18

H-17

H-26

H-27

H-28

H-29

H-30

H-31H-32

H-33

H-34

H-44

H-43H-42H-41

H-40

H-39

H-38

H-37H-36H-35

H-53

H-52

H-51

H-50

H-49

H-48H-47H-46

H-45

H-54

H-55

H-56

H-63

H-62H-61H-60

H-59H-58

H-57

H-64

H-65

H-70

H-69

H-66

H-67

H-68

H-71

H-72

H-73

H-74

H-75

H-76H-77

H-80

H-79

H-78

H-83

H-82

H-81

H-85

H-84

H-86

H-87

H-88

H-89

H-90

H-91

H-103

H-102

H-101H-100H-99

H-98

H-97

H-96

H-95

H-94

H-93

H-92

K-1

K-2

K-3

K-2.1

10

20

15

5

25

30

35

40

45

50

Ro

dis

io F

orm

atio

nC

resm

ina

Fo

rma

tio

n

Fo

rma

tio

n d

e R

ibe

ira

de

Ilh

as

Qz

Qz

Qz

Qz

Qz

Qz

Qz Qz

Qz

Qz

Qz

Qz

Qz

Qz

Qz

Qz

Qz

Qz

QzQz

Qz

Qz

Qz

Qz

Qz

K

K

K

Re

ga

tão

Fo

rma

tio

n

65

75

70

60

80

85

90

95

100

55

500

Me

ters

Fo

rma

tio

n

Lith

olo

gy

Sa

mp

les

Fo

ssils

Se

dim

en

tolo

gy

Me

ters

Fo

rma

tio

n

Lith

olo

gy

Sa

mp

les

Fo

ssils

Se

dim

en

tolo

gy

Appendix

157

A8: Ericeira section, upper part

K-4

K-5

K-6

K-7

K-14

K-13

K-12

K-11

K-10

K-9

K-8

K-18K-17

K-16

K-15

K-25

K-24

K-23

K-22

K-21

K-20

K-19

K-31

K-30

K-29

K-28K-27

K-26

K-38

K-37

K-32

K-33

K-34K-35K-36

K-3.1

120

130

125

115

135

140

105

110

Galé

Form

ation

Qz

Qz

Qz

Qz

Qz

100

Me

ters

Fo

rma

tio

n

Lith

olo

gy

Sa

mp

les

Fo

ssils

Se

dim

en

tolo

gy

Appendix

158

A9: Cresmina section , lower part

Qz

Qz

Qz

Qz

Qz

Qz

Qz

Qz

Qz

Qz

Qz

Qz

Qz

Qz

Qz

Qz

QzQz

Qz

Qz

Qz

QzQz

Qz

Qz

Qz

D D

K

K K

D 1D 2

D 3

D 4

D 5

D 6

D 7

D 8

D 9

D 10

D 11

D 12

D 13

D 14

D 15

D 16

D 17D 18

D 19

D 20/21D 22

D 23

D 24D 25D 26D 27

D 28D 29D 30D 31D 32D 33D 34

D 35D 36D 37

D 38

D 39

D 40

D 41

D 42D 43

D 44

D 45

D 46

D 47

D 49

D 50D 51

D 52

D 53

D 54

D 55

D 56

D 57D 58

D 59

D 60D 61

D 63

D 64

D 65

D 66

D 67

D 62

D 48

Cre

sm

ina F

orm

ation

Co

bre

Me

mb

er

D 87

D 88

D 89

D 90

D 91

D 92

D 93

L1

L2

Qz

Rodis

io F

orm

ation

Galé

Form

ation

Ag

ua

Do

ce

Me

mb

er

Qz

K

K

K

80

90

100

50

60

70

D 67

D 68

D 69

D 70

D 71

D 72

D 73

D 74

D 75

D 76

D 77

D 78

D 79

D 80

D 81

D 82

D 83

D 84

D 85

D 86

Cre

sm

ina F

orm

ation

Ponta

Alta M

em

ber

Pra

ia d

a L

agoa

10

20

30

40

50

0

Mete

rs

Form

ation

Lith

olo

gy

Sa

mp

les

Fo

ssils

Se

dim

en

tolo

gy

Mete

rs

Form

ation

Lith

olo

gy

Sa

mp

les

Fo

ssils

Se

dim

en

tolo

gy

Appendix

159

A9: Cremina section , upper part

L47L48

L49

L50

L51

Qz K

L52

L53

L54

L55

L56

L57

L58

L59

L60

L61L62L63

L64

L65

L66

L67

L68

L69

L70

L71

L77

L78

L79

L80

L81

L82

L83

L84

L85

L86

L87

L88

L89

L90

L91

L92

L93

L94

L95

L97

L98

L99

L100L101

L102

L103

L96

L74

L73

L75L76

L72

K

K

Qz

K

K

K

K

Qz

L22L23

L24

L25

L26

L27

L28

L30L31

L34

L35

L36

L37L38

L39

L40L41

L42

L43

L44

L46

L45

L32L33

L29

L3

L4

L5L6L7

L8

L9

L10

L11

L12

L13

L14

L15

L16

L17L18L19L20L21

Qz

Qz

Qz

K

K

K

K

K

K

K K

Ga

lé F

orm

atio

nA

gu

a D

oce

Me

mb

er

Galé

Form

ation

Ag

ua

Do

ce

Me

mb

er

100

110

120

130

140

150

150

160

170

180

190

200

Mete

rs

Form

ation

Lith

olo

gy

Sa

mp

les

Fo

ssils

Se

dim

en

tolo

gy

Mete

rs

Form

ation

Lith

olo

gy

Sa

mp

les

Fo

ssils

Se

dim

en

tolo

gy

Qz

Qz

Qz

Qz Qz

Qz

Qz

Qz

Qz

Qz

Qz

Qz

Qz

Qz

Qz

Qz

Appendix

160

A10: Luz section, lower part

B 1

B 2

B 3

B 4

B 5

B 6

B 7

B 8

B 9

B 10

B 11

B 12

B 13B 14

B 15

B 16

B 17

B 18

B 19

B 20

B 21

B 22

B 23

Qz

Qz10

20

30

40

50

70

80

90

100

Pa

lorb

ito

lina

Be

ds

G.

tro

ch

ilisco

ide

s F

m

50

60

0

Lo

we

r L

uz M

arls

A10

A11

A13

A15

A16A17A19

A20

A21

A22

A23

A24

A25

A26

A27

A28

A29

A30

A31

A32

A33

A35

A37

A38

A39

A40

A41A43A45A46A47

A49A51

A52A53

A54

A55A56

A57

A58

A59A60

A61

Qz

Qz

Qz

Qz

Qz

Qz

Qz

Qz

Qz

Qz

Qz

Qz

K K K K

Fe

Fe

Fe

K

K

K

K

K

K

K

K

K K

A7A8A9

Up

pe

r L

uz M

arls

Mete

rs

Form

ation

Lith

olo

gy

Sa

mp

les

Mete

rs

Form

ation

Lith

olo

gy

Sa

mp

les

Se

dim

en

tolo

gy

Se

dim

en

tolo

gy

Fossils

Fossils

Appendix

161

A10: Luz section, middle part

110

120

130

140

150

100

A133

A62

A63

A64

A65

A66

A67A68

A69

A70

A71

A72A73

A74aA74A75

A78A76

A79A80A81A82

A83

A84

A85

A86

A87

A88A89

A90

A91

A92

A93

A94A95

A96

A97

A98

A99

A100

A101A102A103

A104

A105

A106

A107

A108

A109

A110

A111

A112

A113

A114

A115

A116

A117

A118

A119

A120

A121

A122

A123

A124A125

A126

A127

A128

A129

A130

A131a

A131A132

A134

A135

A136

A137

A138

A139

A140

A141

A142

A143

A144

A145A146

A147

A148

A149A150A151

A152

A153

A154

A155

A156

A157A158

A159

A135a

Qz

Qz

KK

K

K

K

K

KK

K K

K K

160

170

180

190

200

150

Up

pe

r L

uz M

arls

Up

pe

r L

uz M

arls

Me

ters

Fo

rma

tio

n

Lith

olo

gy

Sa

mp

les

Se

dim

en

tolo

gy

Fo

ssils

Me

ters

Fo

rma

tio

n

Lith

olo

gy

Sa

mp

les

Se

dim

en

tolo

gy

Fo

ssils

Po

rto

de

Mo

s F

m.

Appendix

162

A10: Luz section, upper part

Me

ters

Fo

rma

tio

n

Lith

olo

gy

Sa

mp

les

Se

dim

en

tolo

gy

Fo

ssils

Me

ters

Fo

rma

tio

n

Lith

olo

gy

Sa

mp

les

Se

dim

en

tolo

gy

Fo

ssils

A 200

A 201

A 202

A 203

A 204

A 205

A 206

A 207

A 208

A 209

A 210A 211

A 212

A 213

A 214A 215

A 216

A 217

A 218

A 219

A 220

A 221

A 222

A 223

A196

A197

A198

A199

250

260

270

280

290

A160

A161

A162A163

A164

A165

A166

A167

A168

A169

A170

A171

A172

A173

A174

A175

A176

A177

A178

A179

A180

A181

A182

A183A184A185A186

A187

A188

A189

A190

A191

A192

A193

A194

A195

K

?

200

210

220

230

240

250

Po

rto

de

Mo

s F

m.

Po

rto

de

Mo

s F

m.

Appendix

163

A11: Burgau section, lower part

Qz

Qz

Qz

K

Qz

Qz

Qz

QzQz

G1/48

G2/49

G3/50

G4

G5

G6

G10

G11

G12

G13

G14

G15

G16

G17

G18

G19

G7/8/9

G41

G42

G43

G44

G45

G46

G47

10

20

30

40

50

Pa

lorb

itolin

a B

ed

sG

. tr

och

ilisc

oid

es

Fm

.

Qz

K K

K

K K

C 1

C 2

C 3

C 4

60

70

80

90

100

500

Lo

we

r L

uz

Ma

rls

Lo

we

r L

uz

Ma

rls

Up

pe

r L

uz

Ma

rls

Mete

rs

Form

ation

Lith

olo

gy

Sam

ple

s

Fossils

Se

dim

en

tolo

gy

Mete

rs

Form

ation

Lith

olo

gy

Sam

ple

s

Fossils

Se

dim

en

tolo

gy

Appendix

164

A11: Burgau section, upper part

Echinoderms

Sponges

Cyclamminidae

Agglutinating benthic foram.

Miliolinids

Undifferentiated benthic foram.

Corals

Serpulids

Bryozoans

Bivalves

Brachiopods

Gastropods

Ostracods

Green algae

Charpophytes

Fish debris

Nodular bedding

Cross bedding

Birdseyes

Burrow

Quarz grainsQz

Teepe structure

Small wavy stratification

Current ripples

Bioturbation

Wave ripples

Bioclasts (undifferentiated)

Fossil wood

Orbitolinids

Lithoclasts

Fining upward

Oysters

Graded bedding

Calcareous nodulesK

Dolomitic nodulesD

Trough cross bedding

Tabular cross bedding

Heringbone cross bedding

Flaser bedding

Load cast

Channel

Rudists

Stromatoporids

G26

G27

G28

G29

G30

G31

G32

G33

G34

G35

G36

G39

G38

G37

G40

G25

G20

G21

G23

G22

G24

Up

pe

r L

uz

Ma

rls

130

140

150

110

120

100

Me

ters

Fo

rma

tio

n

Lith

olo

gy

Sa

mp

les

Fo

ssils

Se

dim

en

tolo

gy

Hardground

Sedimentary structures

Fossil content

Lithology

Limestone

Calcareous marl

Claystone

Siltstone

Sandstone

Conglomerate

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