Calcareous nannoplankton of the Red Sea - Ruhr-Universität Bochum

Calcareous nannoplankton of the Red Sea:

A proxy for the reconstruction of the paleoenvironmental conditions during

the late Quaternary

Dissertation

zur Erlangung des akademischen Grades

eines Doktors der Naturwissenschaften

an der Fakultät für Geowissenschaften

der Ruhr-Universität Bochum

vorgelegt von

Heiko-Lars Legge

aus Bochum

Bochum,

Oktober 2007

Die vorliegende Arbeit wurde von der Fakultät für Geowissenschaften der Ruhr-Universität

Bochum als Dissertation zur Erlangung des akademischen Grades eines Doktors der

Naturwissenschaften anerkannt.

Erster Gutachter: Prof. Dr. Mutterlose

Zweiter Gutachter: Prof. Dr. Heimhofer

fachfremder Gutachter: Prof. Dr. Zepp

Tag der Disputation: 9. Januar 2008

Table of contents

I

Acknowledgements III

Abstract V

Kurzfassung VII

1 Introduction 1

1.1 Oceanography of the Red Sea 1

1.2 Paleoceanographic background 3

1.3 Calcareous Nannoplankton 4

1.4 Synopsis 7

Explanatory notes 9

References 10

2 Climatic changes in the northern Red Sea during the last 22,000 years as recorded by

calcareous nannofossils 17

Abstract 17

2.1 Introduction 17

2.2 Material and Methods 20

2.3 Results 22

2.4 Discussion 25

2.4.1 Last Glacial Maximum (LGM, circa 22-19 ka) 26

2.4.2 Heinrich Event 1 (H1, circa 18-15.5 ka) 29

2.4.3 H1-Bolling Transition (circa 15.5-14.6 ka) 30

2.4.4 Bolling-Allerod (circa 14.6-13.1 ka) 32

2.4.5 Younger Dryas (YD, 13.1-11.7 ka) 34

2.4.6 Younger Dryas Transition (circa 11.7-11.4 ka) 36

2.4.7 Earliest Holocene Period (circa 11.4-9.5 ka) 37

2.4.8 Early Holocene to Mid-Holocene Humid Period (circa 9.5-6 ka) 38

2.4.9 Mid-Holocene to Late Arid Holocene (circa 6-0.6 ka) 40

2.5 Conclusion 41

Acknowledgements 42

References 43

3 Nannoplankton successions in the northern Red Sea during the last glaciation (60 to

14.5 ka): Reactions to climate change 53

Abstract 53

3.1 Introduction 53

3.2 Material and Methods 55

Table of contents

II

3.3 Results and Discussion 56

3.4 Conclusion 64

Acknowledgements 64

References 65

4 Late Holocene environmental changes in the northern Red Sea region as indicated by

nannoplankton abundance patterns 71

Abstract 71

4.1 Introduction 71

4.2 Methods 73

4.2.1 Material and geological Background 73

4.2.2 Sampling and preparation 74

4.3 Gephyrocapsa spp. abundance patterns 75

4.4 Discussion 77

4.5 Conclusion 83

Acknowledgements 84

References 85

5 Summary 93

6 Outlook 97

Appendix 99

Curriculum Vitae 119

Acknowledgements

III

Acknowledgements

I am grateful to Dr. Jörg Mutterlose (Bochum) for initiating and supervising this thesis.

I would like to thank Dr. Jürgen Pätzold (Bremen), Dr. Helge W. Arz (Potsdam) and Dr.

Ismene A. Seeberg-Elverfeldt (Bremen) for providing sample material and geochemical data.

Varies kind of support were provided by colleagues at the Ruhr-University of Bochum. Dr.

Jens Steffahn (now Universidad Autónoma de Nuevo Léon MÉXICO) and my “fellow-

sufferers” Andreas Rexfort and Petros Hardas are thanked for a constructive and friendly

working atmosphere. I’m indebted to the working group of mineralogy/petrography for

providing access to their laboratory for sample preparation. Dr. Rolf Neuser is thanked for his

help at the SEM.

Special thanks to Dr. André Bornemann (now Leipzig) for varies kind of help and many

fruitful discussions during the last years.

Above all, I am grateful to Sylvia and my family for their continuous interest and moral

support.

This research was supported the by German Research Foundation (DFG Mu 667/23-1, -2 and

Mu 667/26-1, -2) and the Ruhr-University of Bochum (Allgemeines Promotionskolleg).

IV

Abstract

V

Abstract

The objective of this thesis is to understand the calcareous nannoplankton response to past

environmental changes in the northern Red Sea over the last 60,000 years. Small-sized pro-

and eukaryotic phytoplankton dominates the oligotrophic Red Sea environment. Calcareous

nannoplankton is well abundant in the living phytoplankton community and a key component

of the marine sediments. Due to its distinct and fast reaction to environmental changes and its

good preservation potential, it is well suited for paleoceanographic reconstructions. The study

of the calcareous nannoplankton, and the evaluation of its relative and absolute abundances, is

therefore indicative for paleoceanographic changes in the northern Red Sea. Calcareous

nannoplankton findings supplemented by geochemical and sedimentological data were

compared with other marine and terrestrial proxy records and discussed with respect to the

climate history established so far. The thesis presented here is divided into three parts, which

focus on the following aspects of the northern Red Sea history:

1.) Paleoceanographic reconstruction of the northern Red Sea over the last 22,000 years.

2.) Calcareous nannoplankton response to well-documented climate oscillations (long-term

cooling cycles, Heinrich events) of Marine Isotope Stages 3 and 2.

3.) High-resolution analysis of the late Holocene by investigating laminated sediments from a

brine-filled, submarine basin located in the northern Red Sea (Shaban Deep).

It is evident from this study that the calcareous nannoplankton community mirrors

paleoceanographic changes of the northern Red Sea, which are driven by long- and short-

termed climate variations. The paleoceanography of the northern Red Sea shows marked

similarities with the climate and vegetation history of northern Africa and the Near East.

Intervals of enhanced primary production in the marine environment go along with humid

periods on land. Another important result of the thesis is the existence of calcareous

nannoplankton in the northern Red Sea during glacial periods. Furthermore, our results

question the applicability of the geographic isolation of the Red Sea to explain calcareous

nannoplankton community changes during these periods: Due to the restricted water exchange

with the Gulf of Aden-Indian Ocean via the narrow and shallow Strait of Bab-el-Mandeb, the

modern Red Sea is characterized by high saline waters. The restricted water exchange during

glacial times (e.g., Last Glacial Maximum) is a result of the global sea level lowering and has

caused hypersaline conditions. In previous studies it has been suggested that the glacial

progression caused the absence of planktic foraminifera and the occurrence of a monospecific

pteropod population in the northern Red Sea. There is ample evidence from this study that

basic variations of the calcareous nannoplankton community structure, even during full

Abstract

VI

glacial conditions, were driven by regional hydrographic changes under the atmospheric

influence of the extratropics. Changes of surface-water stratification and thereby of the

vertical nutrient distribution within the photic zone seem to be extremely important. The

influence of these processes was probably not restricted to glacial periods; even the

calcareous nannoplankton community structure during the Holocene seems to be affected.

Kurzfassung

VII

Kurzfassung

Im Rahmen der vorliegenden Dissertation sollte die Reaktion des kalkigen Nannoplanktons

auf Veränderungen der Paläoumwelt an Sedimentkernen aus dem nördlichen Roten Meer über

den Zeitraum der letzten 60,000 Jahre untersucht werden. Unter oligotrophen Bedingungen,

wie sie im heutigen Roten Meer vorherrschen, sind die Phytoplanktonvergesellschaftungen

durch kleine Pro- und Eukaryonten gekennzeichnet, wobei das kalkige Nannoplankton die

wichtigste fossil überlieferbare Gruppe repräsentiert. Aufgrund ihrer differenzierten und

schnellen Reaktion auf Veränderungen der Umwelt und ihres guten Erhaltungspotentials

bieten Wechsel in der Zusammensetzung der kalkigen Nannoplanktongemeinschaften eine

ideale Möglichkeit ozeanographische Veränderungen zu rekonstruieren. Durch die Analyse

des kalkigen Nannoplanktons sowie die quantitative und qualitative Erfassung ihrer

Häufigkeiten, können paläoozeanographische Schwankungen im nördlichen Roten Meer

aufgezeigt werden. Die Ergebnisse aus den Untersuchungen des kalkigen Nannoplanktons

wurden, ergänzt durch umfangreiche geochemisch-sedimentologische Daten, mit bereits

publizierten Arbeiten aus dem marinen und terrestrischen Bereich verglichen und vor dem

Hintergrund der derzeit bekannten Klimageschichte diskutiert. Innerhalb der vorliegenden

Dissertation lassen sich drei Untersuchungsschwerpunkte unterscheiden:

1.) Die Rekonstruktion der Paläoozeanographie des nördlichen Roten Meeres während der

letzten 22,000 Jahre.

2.) Die Untersuchung des Einflusses der für die Marinen Isotopenstadien 3 und 2

charakteristischen Klimaschwankungen (Abkühlungszyklen, Heinrichereignisse) auf das

kalkige Nannoplankton.

3.) Eine zeitlich hochauflösende Analyse des späten Holozäns unter Zuhilfenahme laminierter

Sedimente aus einem im nördlichen Roten Meer gelegenen Solebecken (Shaban Tief).

Die durchgeführten Untersuchungen zeigen, dass sich paläoozeanographische Veränderungen

im nördlichen Roten Meer, gesteuert durch lang- und kurzfristige Klimavariationen, in den

Nannoplanktongemeinschaften widerspiegeln. Langfristige Veränderungen der Paläoozeano-

graphie zeigen deutliche Parallelen zur Klima- und Vegetationsgeschichte Nordafrikas und

des Nahen Ostens. So korrelieren beispielsweise Intervalle erhöhter Primärproduktion im

nördlichen Roten Meer meist mit humiden Klimaperioden im terrestrischen Bereich. Ein

weiteres wichtiges Ergebnis der durchgeführten Untersuchungen ist der Nachweis des

kalkigen Nannoplanktons im nördlichen Roten Meer auch während glazialer Zeitabschnitte.

Von besonderer Bedeutung ist in diesem Zusammenhang die Beobachtung, dass

Schwankungen innerhalb der glazialen Nannoplanktongemeinschaft nicht alleine über die

Kurzfassung

VIII

isolierte geographische Lage des Roten Meeres erklärt werden können: Der eingeschränkte

Austausch mit dem Indischen Ozean über die enge und flache Straße von Bab-el-Mandeb

führt zu erhöhten Salzgehalten im heutigen Roten Meer. Bedingt durch die Regression des

globalen Meeresspiegels waren kühlere Phasen in der Erdgeschichte (z.B. Letztes Glaziales

Maximum) durch hypersaline Bedingungen geprägt. Die Verschärfung des eiszeitlichen

Klimas ist im nördlichen Roten Meer durch das Verschwinden planktonischer Foraminiferen

und die Ausbildungen einer nahezu monospezifischen Pteropoden-Gemeinschaft

gekennzeichnet. Die im Rahmen der Dissertation durchgeführten Untersuchungen deuten

darauf hin, dass selbst während der Hochphase der letzten Vereisung wesentliche

Veränderungen innerhalb der kalkigen Nannoplanktongemeinschaft durch andere Faktoren

gesteuert wurden. Veränderungen der regionalen Hydrography, insbesondere Schwankungen

der Intensität und Häufigkeit der winterlichen Durchmischung bedingt durch den variierende

Einfluss außertropischer Luftmassen, scheinen Schlüsselelemente darzustellen. Dieses gilt

nicht nur für das glaziale Rote Meer. Vergleichbare Prozesse scheinen auch während des

Holozäns die Zusammensetzung des kalkigen Nannoplanktons beeinflusst zu haben.

Chapter 1: Introduction

1

1 Introduction

The chapter “Introduction” is subdivided into four sections. First, the modern oceanography

of the Red Sea is briefly summarized in section 1.1. The aim of section 1.2 is to illustrate how

basic changes of the Red Sea environment were connected to variations of the global sea

level. Section 1.3 highlights the importance of calcareous nannoplankton in modern research

and gives some remarks concerning the terminology. In addition, with special reference to the

calcareous nannoplankton, the setting of the study domain is shortly discussed. In section 1.4,

the major goals of the thesis will be summarized. An overview of the performed work will be

given briefly in conjunction with the used sediment material.

1.1 Oceanography of the Red Sea

The Red Sea is a desert-surrounded ocean basin with a restricted seaway to the Indian Ocean-

Gulf of Aden via the shallow (137 m) and narrow (30 km) Strait of Bab-el-Mandeb at its

southern end (Figure 1.1). Encompassing the

Sinai Peninsula, the Gulfs of Suez and

Aqaba mark the northernmost extensions of

the Red Sea. As part of the Red Sea-Gulf of

Aden rift system (with the Afar triple

junction), the Red Sea represents an example

of a relatively young ocean basin (Bosworth

et al., 2005). It is located in the northwest-

trending rift zone where the African-Arabian

plate was broken and spread apart. The

spreading axis extends up the Red Sea main

basin. In addition, due to the movement of

the Arabian plate in northeasterly direction,

a large fracture developed at the passive

plate margin (Almond, 1986; Bosworth et al., 2005). It forms the Gulf of Aqaba and the

trough of the Jordan valley with a number of lake basins (Dead Sea, Sea of Galilee).

The modern climate is arid, with low precipitation (10–200 mm/year) and high rates

(2000 mm/year) of evaporation (Edwards, 1987; Sofianos et al., 2002). No major river system

supplies the Red Sea with freshwater and the influence of temporary run-off (e.g., from

wadis) on the hydrological budget is negligible (Morcos, 1970; Hoelzmann et al., 1998;

Siddall et al., 2003).

Figure 1.1. Map showing the Red Sea area and the Gulf of Aden-Indian Ocean region.


2

Northwesterly winds characterize the northern Red Sea during the entire year (Edwards,

1987). In summer these winds also dominate the southern parts, whereas in winter the

influence of the Indian monsoon system leads here to a reversal of the main wind direction.

The interaction between southeasterly winds over the southern Red Sea and the prevailing

northwesterly winds over the northern parts results in a convergence zone across the central

Red Sea (e.g., Edwards, 1987).

The circulation patterns of the Red Sea waters are anti-estuarine and driven by wind and

especially thermohaline forcing (e.g., Eshel and Naik, 1997). Warm and normal saline waters

enter the Red Sea from the Gulf of Aden. As this water moves northwards along the

longitudinal axis of the basin, it becomes cooler and saltier. Due to biological consumption

(Weikert, 1987) in the southernmost Red Sea, the nutrient gradients along the longitudinal

axis are less pronounced than those for salinity and temperature. Compared with the

oligotrophic central and northern Red Sea, the southern Red Sea is relatively nutrient-rich due

to the inflow from the Gulf of Aden (for details see e.g., Levanon-Spanier et al., 1979; Shaikh

et al., 1986; Weikert, 1987 and references therein). Enhanced productivity in the southern Red

Sea is displayed by an upward shift of the Oxygen Minimum Zone (OMZ) and lower values

of dissolved oxygen compared with the northern part (Weikert, 1987). Bacterial consumption,

associated with the organic decomposition, reduces the oxygen concentration rapidly below

surface waters. In the northern Red Sea lowest values of around 1–1.75 ml O2/l are located at

400–500 m, whereas the oxygen concentration in the southern Red Sea decreases (0.3–0.5 ml

O2/l) at 300–400 m water-depth. Below the OMZ, the concentration increases slowly again.

Values of around 2–3 ml O2/l are found in the deep-water (Edwards, 1987; Weikert, 1987). It

is worth mentioning that the oxygen gradient between the northern and southern Red Sea

reflect also the increasing distance from the area of intermediate and deep-water formation.

The journey of the Red Sea surface waters flowing from south to north eventually ends in

the northern most parts of the Red Sea where it sinks and forms intermediate and deep-water.

The resulting southward deep flow ventilates the Red Sea (e.g., Grasshoff, 1969; Morcos,

1970; Wyrtki, 1971; Manis, 1973; Cember, 1988; Woelk and Quadfasel, 1996; Eshel and

Naik, 1997). Different locations and modes of intermediate and deep-water formation are

identified (cf. Figure 1 in Arz et al., 2003): (1a) Deep convection as well as (1b) direct

injection beneath the pycnocline in the northern Red Sea nearby the Sinai tip (Cember, 1988).

(2) Shallow convection at the collision site of the Western and Eastern boundary current at

around 26–26.5°N (Eshel and Naik, 1997). (3) Deep-water formation in the northern Red Sea


3

initiated by the outflow of dense waters from the shallow (50–70 m) Gulf of Suez (Woelk and

Quadfasel, 1996).

1.2 Paleoceanographic background

It is well accepted that the Red Sea environment and its biological interior suffered extreme

oceanographic changes during the last glacial–interglacial cycles (e.g., Locke and Thunell,

1988; Hemleben et al., 1996; Rohling et al.,

1998; Siddall et al., 2003; Arz et al., 2007). As

stated above, the restricted water exchange with

the Gulf of Aden-Indian Ocean via the

bottleneck at Bab-el-Mandeb controls the

hydrological budget. Hence, variations of the

global sea level dictate the quantity of water

entering the Red Sea. During the Last Glacial

Maximum (LGM), for example (Figure 1.2), the

sea level decreased nearly by 120 m. This drop

reduced the water depth at the Strait of Bab-el-

Mandeb to approximately 17 m. The salinity of

the Red Sea waters increased dramatically due

to the strongly limited exchange (Winter et al., 1983; Locke and Thunell, 1988; Hemleben et

al., 1996; Siddall et al., 2003; Arz et al., 2007). Stable oxygen isotope measurements reveal

increased values in comparison to the global trend (e.g., Hemleben et al., 1996). The

hypersaline conditions (49–50‰ and more; Thunell et al., 1988; Hemleben et al., 1996;

Fenton et al., 2000) reached the tolerance limit for various planktic organisms. Final

ecological consequences of the glacial progression were the absence of planktic foraminifera

(“aplanktic zone”; e.g., Winter et al., 1983; Almogi-Labin et al., 1991; Hemleben et al., 1996)

and the occurrence of a nearly monospecific pteropod (Creseis acicula) community (Chen,

1969; Almogi-Labin, 1982; Winter et al., 1983).

In comparison to foraminifera and pteropoda, little is known about the calcareous

nannoplankton of the Red Sea (McIntyre, 1969; Müller, 1976; Winter, 1982a, 1982b). The

few calcareous nannoplankton studies, which have been carried out yet, suggest a massive

decrease or even the total absence of calcareous nannoplankton during the LGM (McIntyre,

1969; Winter et al., 1983). However, in order to understand past changes of the Red Sea

environment and its impact on the phytoplankton community, continuous and well-dated

Figure 1.2. The Red Sea at present and

during the LGM is schematically shown. For detail see text.


4

marine records are necessary. Until now a detailed reconstruction of the calcareous

nannoplankton abundance patterns is still missing.

1.3 Calcareous Nannoplankton

Like other subtropical pelagic ecosystems, the northern Red Sea is dominated by minute

primary producers (e.g., Winter et al., 1979; Kleijne et al., 1989; Lindell and Post, 1995).

Adapted to oligotrophic environmental conditions, photo-

autotrophic pico- and nan(n)oplankton are the most

common phytoplankton. The latter group comprises as one

important component of the phytocoenosis the calcareous

nannoplankton (Levanon-Spanier et al., 1979; Winter et al.,

1979; Kleijne et al., 1989).

Calcareous nannoplankton, or more precisely their

fossil remains, the calcareous nannofossils, are well suited

as biostratigraphic markers for Mesozoic and Cenozoic

sediments (e.g., Hay and Mohler, 1967; Martini, 1971;

Thierstein, 1971, 1976; Sissingh, 1977; Perch-Nielsen,

1979; Okada and Bukry, 1980; Bown et al., 1998). For a biologist, the term calcareous

nannoplankton covers rather a morphological conglomeration than a coherent phylogenetic

cluster. Based primarily on morphology and crystal

structure, it includes traditionally three major groups:

coccoliths, nannoliths and tiny calcareous dinocysts

(calcispheres). The subjects of the thesis, coccoliths

and nannoliths, are probably taxonomically closely

related, whereas the exceptional position of the

calcareous dinocysts is obvious. Coccoliths, mineral-

ized low Mg-calcite scales (Figure 1.3, Emiliania

huxleyi), are formed by coccolithophorids (or cocco-

lithophores). Coccolithophorids are algal protists,

which belong to the class Prymnesiophyceae (Green

and Jordan, 1994; Billard and Inouye, 2004). As part

of the division Haptophyta, two layers of organic

scales, two golden-brown chloroplasts, two flagella (without any hair-like appendage) and a

haptonema identify them despite their coccoliths (Figure 1.4). It is worth mentioning that a lot

Figure 1.3. E. huxleyi.

Coccolith-bearing species. Bar = 1µm.

Figure 1.4. Coccolithophorid cell;

motile live cycle stage with two

flagella and the haptonema (modified after Billard and Inouye, 2004).


5

of coccolithophorids (respectively their main live cycle stage) are non-motile (Pienaar, 1994).

Furthermore, the haptonema is often strongly reduced or even visual absent (Green and

Jordan, 1994).

Some nannoliths (Figure 1.5, Florisphaera profunda (Nannoliths incertae sedis)) are

probably also produced by coccolithophorids, whereas the origin of some other is still

unknown (cf. Young et al., 2003). Most recent plankton- and micropaleontological studies

make no differentiation between coccolith- and nannolith-bearing species and include both

under the term: coccolithophorids, coccolithophores or calcareous nannoplankton in their

description. To avoid confusion, the presented study followed this schema. The term

calcareous nannoplankton/-fossils comprise here coccoliths and nannoliths only; calcareous

dinocysts have not been considered.

Due to their distinct affinity to autecological factors (e.g., Brand, 1994; Winter et al.,

1994; Ziveri et al. 2004), calcareous nannoplankton is often applied for paleoceanographic

reconstructions (e.g., McIntyre and Bé, 1967; McIntyre,

1970; Okada and Honjo, 1973, 1975; Honjo, 1977; Okada

and McIntyre, 1977, 1979; Reid, 1980; Roth and

Coulbourn, 1982; Samtleben and Schröder, 1992; Molfino

and McIntyre, 1990; Beaufort et al., 1997; de Garidel-

Thoron et al., 2001). Information gained from intensive

biogeographic studies in the Atlantic and Pacific Ocean

allow a separation of five major calcareous nannoplankton

zones: Subartic, temperate, subtropical, tropical and

subantartic (McIntyre and Bé, 1967; Okada and Honjo,

1973, summarized in Winter et al., 1994). The

oceanographic setting of the Red Sea corresponds closely to that of the subtropical zone. In

general there is a permanent stratification of the highly transparent upper waters caused by

high solar irradiance and overall stable environmental conditions. The stratification impeded

the return of recycled inorganic nutrients from deeper waters and caused thereby oligotrophic

surface water conditions (e.g., Li et al., 1998; Post, 2005). Generated by the stratification and

characterized by specialized calcareous nannoplankton communities (Winter et al., 1994, and

references therein), a vertical subdivision of the photic zone in at least three intervals is

possibly: the upper photic zone (0–80 m; characteristic calcareous nannoplankton species e.g.,

Rhabdosphaera clavigera, Discosphaera tubifera), the middle photic zone (80–120 m; less

well defined by specific marker species) and the lower photic zone (120–220 m; characteristic

Figure 1.5. F. profunda.

Nannolith-bearing species. Bar = 1µm.


6

calcareous nannoplankton species e.g., Florisphaera profunda, Gladiolithus flabellatus (=

Thorosphaera flabellate)).

The hydrographic setting and hence the overall oligotrophic environmental conditions, as

typical for the subtropics, are modulated in our study domain during the winter season. The

annual cycle is characterized by a short winter period with lower temperatures. Cooling-

induced overturning of the water masses results in a short interval of high marine

productivity. Detailed insights into the succession patterns of various phytoplankton groups in

the northern Red Sea region are especially

available from the Gulf of Aqaba (e.g.,

Winter et al., 1979; Lindell and Post, 1995;

Post, 2005). Accompanied by distinctive

succession patterns of the phytoplankton,

the environment is characterized by a semi-

annual alternation of water column

conditions (Post, 2005). Surface waters are

characterized during most of the year by a

shallow, but stable stratification. Cold air

incursion in late fall–early winter causes a

density increase of the surface waters

(Figure 1.6). The resulting erosion of the

stratification induces a deep convective

mixing of the waters. Although limited in

strength, in relation to the Gulf of Aqaba,

the overturning of the waters during the

short winter season triggers also the

northern Red Sea environment. As

described in section 1.1, the open northern

Red Sea is the area of intermediate and

deep-water formation. Several studies have

demonstrated the sensitivity of the northern Red Sea hydrography to mid-latitude continental

climate and the influence of temperature (and evaporation) as key factors controlling surface

water stratification (e.g., Eshel et al., 2000; Arz et al., 2003; Rimbu et al., 2006).

Figure 1.6. Seasonal succession of three

major phytoplankton groups in the Gulf of

Aqaba (Lindell and Post, 1995). Shown are

cell numbers per 1 m2 for a 600 m depth

water column. Seawater temperatures at

Eilat (4 m water depth) are from Ben-David-Zaslow et al. (1999).


7

1.4 Synopsis

Late Pleistocene paleoecological conditions of the Red Sea waters were reconstructed in the

past mainly by using foraminifera (e.g., Berggren and Boersma; 1969, Geiselhart, 1998;

Fenton et al., 2000) and pteropoda (e.g., Chen et al., 1969; Reiss et al., 1980; Locke and

Thunell; 1988; Almogi-Labin et al., 1998). The current thesis presents the first continuous and

well-dated calcareous nannoplankton record from the northern Red Sea. The objective of this

study was to evaluate the calcareous nannoplankton patterns in response to paleoceanographic

changes over the last 60,000 years (Marine Isotope Stages MIS 1, MIS 2 and MIS 3). The

calcareous nannoplankton results were compared with other marine as well as terrestrial

proxy records to allow an appropriate reconstruction of the northern Red Sea paleoclimate. To

gain relative and absolute abundances, the splitting-filtration technique of Andruleit (1996)

was used for preparation. At least 300 specimens per sample were counted by using a

scanning electronic microscope (LEO 1530 Gemini FESEM). The species identification

followed the taxonomic descriptions of Jordan and Kleijne (1994) and Young et al. (2003).

More details are given in the following chapters and the Appendix.

An overview of the research performed follows. First of all, we take a brief look to the

studied sediment cores GeoB 5844-2 and GeoB 7805-1. Following this, it is shown how the

three sub-studies of the thesis are related and why the performed work focused on specific

aspects of the northern Red Sea paleoceanography.

GeoB 5844-2

Late Glacial-Holocene sediments covering the last 60,000 years

have been studied using core GeoB 5844-2 (Figure 1.7). The

gravity core was taken in 1999 near the tip of the Sinai Peninsula

(27°42´81"N, 34°40´90"E; water depth: 963 m) during RV

Meteor cruise M44/3 (Pätzold et al., 2000).

GeoB 7805-1

Late Holocene sediments spanning the last two millennia have

been analyzed using core GeoB 7805-1 (Figure 1.7). The core

was taken with a multicorer during RV Meteor Cruise M52/3

(Pätzold et al., 2003) from the eastern sub-basin of the Shaban

Deep (26°13!9" N and 35°22!6" E; water depth: 1447 m).

Figure 1.7. Sketch of the

northern Red Sea area

showing locations of the

cores GeoB 5844-2 and

GeoB 7805-1.


8

The calcareous nannoplankton findings obtained from sediment core GeoB 5844-2 are

presented in the form of two manuscripts. The first manuscript deals with the calcareous

nannoplankton variations over the last 22,000 years: Climatic changes in the northern Red

Sea during the last 22,000 years as recorded by calcareous nannofossils. An interesting result

obtained from this study is the existence of calcareous nannoplankton during the hypersaline

period in the northern Red Sea. Though faced with the overwhelming importance of the

glacial salinity increase, some evidence indicates a lasting extratropical atmospheric control

on northern Red Sea environmental changes and calcareous nannoplankton abundance

patterns especially during cold periods like the LGM, Heinrich event 1 and the Younger

Dryas.

To check these findings in more detail, we went further back in time and investigated the

development of the calcareous nannoplankton community since the onset of MIS 3 (around

60,000 years ago) in the second study. The calcareous nannoplankton in sediment core GeoB

5844-2 offered the unique possibility for the examination of millennial-scale environmental

changes during the hypersaline period of the northern Red Sea in response to the longer-term

cooling cycles and Heinrich events. One major aim was to examine whether the calcareous

nannoplankton in the northern Red Sea responses to the well-documented climate oscillations

of the North Atlantic realm or not. The results are presented in the second manuscript:

Nannoplankton successions in the northern Red Sea during the last glaciation (60 to 14.5 ka):

Reactions to climate change. The basic patterns of the North Atlantic climate oscillations

were clearly found in this study. An interesting observation was the recurrent appearance of

E. huxleyi and Gephyrocapsa spp. (G. ericsonii and G. oceanica). The longer-term cooling

cycles are characterized by increasing values of Gephyrocapsa spp. Caused by cold air

incursion from the extratropics, the Gephyrocapsa spp. are thought to indicate an intensified

winter mixing of the waters. The results of the second study lead to the conclusion that the

atmospheric influence from the mid-latitudes via its impact on the northern Red Sea

hydrography was the driving force for abundance variations in calcareous nannoplankton

during the last glacial.

One major objective of the third study undertaken on sediment core GeoB 7805-1 from

the Shaban Deep was to test if the calcareous nannoplankton is also a useful indicator for the

late Holocene climate instabilities. The results are presented in the third manuscript: Late

Holocene environmental changes in the northern Red Sea region as indicated by

nannoplankton abundance patterns. Compared to the last glacial–interglacial cycle, climate

fluctuations during the last two millennia are less pronounced. The Shaban Deep is a high


9

saline, anoxic submarine basin located in the northern Red Sea (Pätzold et al., 2003). The

extreme conditions in the Shaban Deep (e.g., Hartmann et al., 1998) prevent (multi-cellular)

benthic live. The resulting lack of bioturbation allows the sampling and examination of

undisturbed sediments with high resolution. Results presented here suggest that the

Gephyrocapsa spp. serve as a sensitive indicator for the northern Red Sea winter environment

in late Holocene times.

Following this outline, the thesis is organized as follows: First of all we take a closer look

at the paleoceanography of the northern Red Sea over the last 22,000 years (Chapter 2). The

next section deals with the calcareous nannoplankton succession during the late glacial

(Chapter 3). Following the study from the Shaban Deep (Chapter 4), the thesis concludes with

a summary (Chapter 5) and an outlook (Chapter 6) of possibly future research.

Explanatory notes

The author of the thesis has performed the micropaleontological research and has written the

presented manuscripts. The three manuscripts are either published (Climatic changes in the

northern Red Sea during the last 22,000 years as recorded by calcareous nannofossils by

H.L. Legge, J. Mutterlose and H.W. Arz; Paleoceanography, 2006, vol. 21, PA1003 - Chapter

2), submitted (Nannoplankton successions in the northern Red Sea during the last glaciation

(60 to 14.5 ka): Reactions to climate change by H.L. Legge, J. Mutterlose, H.W. Arz and J.

Pätzold; submitted for publication to EPSL - Chapter 3) or ready to submit (Late Holocene

environmental changes in the northern Red Sea region as indicated by nannoplankton

abundance patterns by H.L. Legge and J. Mutterlose - Chapter 4) for publication to an

international journal. The University of Bremen (Dr. Jürgen Pätzold) and the GFZ-Potsdam

(Dr. Helge W. Arz) provided the depth-age models and the complementary geochemical

information for the studied sediment cores. Details are given in the respective manuscripts.

The original manuscripts have been converted to achieve a uniform format of the thesis and

formal errors have been corrected. Sample material and prepared nannoplankton slides are

housed at the Ruhr-Universität Bochum.


10

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16

Chapter 2: Climatic changes in the Red Sea during the last 22,000 years…

17

2 Climatic changes in the northern Red Sea during the last 22,000 years as recorded by

calcareous nannofossils

Abstract

We present a high-resolution record of calcareous nannofossils from the northern Red Sea for

the last 22 kyr. Extreme conditions with enhanced salinities during the Last Glacial Maximum

are characterized by high values of Gephyrocapsa ericsonii. The dominance of Emiliania

huxleyi in Heinrich event 1 indicates a climatic cooling favoring the bloom of opportunistic

species. The calcareous nannofossils record a two-step onset of the postglacial humid climate,

punctuated by the Younger Dryas. Both steps show an early oligotrophic phase dominated by

Florisphaera profunda and Gladiolithus flabellatus and a subsequent fertile phase

characterized by E. huxleyi. The Younger Dryas is described by high values of Gephyrocapsa

oceanica, indicating an increased mixing of the water column. In the late Holocene, repetitive

increases in abundance of F. profunda and G. flabellatus reflect restricted oligotrophic

conditions, caused by the high aridity following the Holocene humid period.

Key words: Red Sea; palaeoceanography; calcareous nannofossils; LGM; H1; termination;

Holocene

2.1 Introduction

The Red Sea is a restricted basin, surrounded by deserts, with a seaway to the Indian Ocean

(Gulf of Aden) via the shallow (137 m) Strait of Bab-el-Mandeb (Figure 2.1). The climate is

arid, with low precipitation (10–200 mm yr-1

) and high rates of evaporation (2000 mm yr-1

).

No major drainage network supplies the Red Sea and the influence of temporary runoff is

negligible (Morcos, 1970; Grasshoff, 1975; Hoelzmann et al., 1998).

Northwesterly winds blow to the south during the entire year over the northern Red Sea

and over the southern Red Sea during summer. In winter, the influence of the Indian monsoon

system leads to a seasonal reversal of the wind direction over the southern and central parts of

the Red Sea where southeasterly winds result in a convergence zone at 20°–25°N. The

circulation patterns are antiestuarine and parallel the longitudinal axis of the Red Sea. Warm

(26°–30°C) normal saline (36–37‰) Gulf of Aden waters enter the Red Sea. As this water

moves northward, it becomes cooler (19°–26°C) and saltier (40–41‰), eventually sinking to


18

form southward flowing intermediate and deep water masses, which ventilate the deep Red

Sea (e.g., Grasshoff, 1969; Morcos, 1970; Wyrtki, 1971; Cember, 1988; Woelk and

Quadfasel, 1996).

Because of biological consumption in the southernmost Red Sea, the nutrient gradients

along the longitudinal axis are less pronounced than those for salinity and temperature

(Weikert, 1987). In general, a unimodal

annual cycle in primary production,

abundance and diversity characterizes

the plankton communities of the

northern and central Red Sea (euphotic

zone ! 100 m). Low primary produc-

tion and low standing crops of the

phytoplankton were observed in the

stratified waters in summer times. In

winter cooling-induced overturning of

the waters results in a short interval of

high productivity. The southern Red

Sea is eutrophic because of the inflow

of waters from the Gulf of Aden during

the entire year (for details, see, e.g.,

Levanon-Spanier et al. (1979), Dowidar

(1983), Shaikh et al. (1986), and

Weikert (1987) and references therein). Plankton, organic detritus and inorganic particles,

especially sediment uptake from the shallow bottom and the narrow shores near the Strait of

Bab-el-Mandeb (Weikert, 1987), increase turbidity thereby reducing the depth of the euphotic

zone (" 60 m).

The higher productivity in the southern Red Sea is also displayed by an upward shift of the

oxygen minimum zone (OMZ) and lower values of dissolved oxygen compared with the north

(Weikert, 1987). In addition, the increasing distance from the area of intermediate and deep

water formation and ventilation (open northern Red Sea and Gulf of Suez (e.g., Cember,

1988; Woelk and Quadfasel, 1996; Eshel and Naik, 1997)) is noteworthy. The surface waters

are well oxygenated in all parts of the Red Sea (around 4.8–4 ml O2/l) with regard to the

oxygen saturation values in warm and saline waters (Edwards, 1987). Oxygen consumption,

due to the organic decomposition, reduces the concentration rapidly below the photic zone

Figure 2.1. Map showing the Red Sea area and the

Gulf of Aden with a detail view of the core location.


19

with a distinctive minimum located between 300 and 500 m water depth. Beneath this depth,

the oxygen concentration increases slowly toward 2–3 ml O2/l in the deep water. In the

northern Red Sea lowest values of around 1–1.75 ml O2/l are located at 400–500 m, whereas

the oxygen concentration in the southern Red Sea decreases up to 0.3–0.5 ml O2/l already

around 300–400 m (Edwards, 1987; Weikert, 1987).

Like other oligotrophic marine ecosystems, the northern Red Sea is dominated by small

phytoplankton (Winter et al., 1979; Lindell and Post, 1995). Picoplankton and calcareous

nannoplankton, the latter consists of coccolithophorids, nannoliths and calcareous

dinoflagellates, are among the most common primary producers (Levanon-Spanier et al.,

1979; Winter et al., 1979; Kleijne et al., 1989). Coccolithophorids are biflagellate algal

protists with mineralized low magnesium-calcite plates or coccoliths, which cover the cell

surface. They belong to the class Prymnesiophyceae (division Haptophyta) and inhabit the

photic zone (0–200 m) of recent oceans (Green and Jordan, 1994). Since the fertility of

coccolithophorids depends on autecological factors (nutrients, temperature, salinity)

coccoliths are often used for paleoceanographic and paleoclimatic reconstructions (e.g.,

McIntyre and Be´, 1967; Okada and McIntyre, 1977; Molfino and McIntyre, 1990; Beaufort

et al., 1997). In contrast to other microfossils (e.g., foraminifera and pteropoda), little is

known about the distribution patterns of calcareous nannofossils (= fossilized calcareous

nannoplankton) in sediments of the Red Sea (but see McIntyre (1969), Müller (1976), and

Winter (1982a, 1982b)).

The Red Sea environment was subjected to large hydrographic changes during the last

glacial interglacial cycles. During the Last Glacial Maximum (LGM), the global sea level

dropped nearly 120 m (Locke and Thunell, 1988), reducing the water depth at the Strait of

Bab-el-Mandeb (now around 137 m) to approximately 17 m. The salinity increased

dramatically because of the reduced water exchange between the Indian Ocean and the Red

Sea. Oxygen isotopes reveal increased values in comparison to the global record (Hemleben

et al., 1996). On the basis of oxygen isotope data (compared to the open ocean record (e.g.,

Thunell et al., 1988; Hemleben et al., 1996)) and microfossil assemblage compositions (e.g.,

upper salinity limit of planktic foraminifera (cf. Hemleben et al., 1989)) the salinity estimates

for the northern Red Sea range around 50‰ and more (Fenton et al., 2000, and references

therein). The hypersaline conditions reached the tolerance limit for various planktic

organisms. Responses to this development include the absence of planktic foraminifera

(‘‘aplanktic zone’’ (e.g., Winter et al., 1983; Almogi-Labin et al., 1991; Hemleben et al.,

1996)) and the occurrence of a monospecific pteropod population (Creseis acicula) in the


20

northern Red Sea (Chen, 1969; Almogi-Labin, 1982; Winter et al., 1983). Previous studies

based on calcareous nannofossils suggest a massive decrease or even their total absence

during the LGM in the Red Sea. McIntyre (1969) argued that high salinities rather than cold

conditions are responsible for an impoverished flora in the central Red Sea. According to

Winter et al. (1983) high salinities caused the absence of calcareous nannofossils in the

northern Red Sea region.

In this study, we present a high-resolution record of calcareous nannofossils from core

GeoB 5844-2 in the northern Red Sea (Figure 2.1). The objective of this study is to analyze

the climatic variations in the northern Red Sea region for the last 22,000 years. The

nannofossil record, including relative and absolute abundances, is presented in combination

with geochemical results published recently for the same core (Arz et al., 2003a, 2003b). The

geochemical data consist of total organic carbon and carbonate, the vertical gradients in stable

oxygen isotopes, alkenone paleothermometry data and reconstructed paleosalinities. Our

findings are discussed in the context of the climatic history of the Red Sea and surrounding

areas for the last 22,000 years.

We want to prove that calcareous nannofossils can record paleoceanographic changes of

the past on a highresolution scale. A second objective of this study is to correlate the marine

findings with the climatic variations described from the surrounding landmasses. We further

want to test whether there is any evidence for synchronous climatic changes in the high

latitudes of the Northern Hemisphere and the northern Red Sea.

2.2 Material and Methods

The investigated gravity core GeoB 5844-2 was taken in 1999 near the tip of the Sinai

Peninsula (Figure 2.1; 27°42´81"N, 34°40´90"E, water depth: 963 m) during R/V Meteor

cruise M44/3 (Pätzold et al., 2000). The stratigraphic framework for core GeoB 5844-2 is

based on 14 calibrated 14

C AMS (Accelerated Mass Spectrometer) datings (Arz et al., 2003a,

2003b). Sedimentation rates are on average 12 cm kyr-1

during the Late Glacial and decrease

to about 6 cm kyr-1

in the Holocene. In order to achieve an average resolution of about 300

yrs, the sampling interval for the nannofossil study was 2 cm in the upper 100 cm and 4 cm

from 102 to 182 cm depth.

Samples were investigated using a scanning electron microscope (SEM). In order to

control SEM preparation and to study general changes in the floral composition light

microscope study was also undertaken using simple smear slides and pipette strew slides (e.g.,

Bown and Young, 1998). The samples were examined with an OLPYMPUS BH-2


21

microscope using polarizing light at a magnification of x1250 and dark field at a

magnification of x150. The samples for the SEM examination were prepared using the

filtration technique of Andruleit (1996). The freeze-dried sediment was weighted on a high-

precision balance, wet separated with a rotary splitter (FRITSCH Laborette 27) and filtered

through polycarbonate filters (pore size 0.25 µm) by means of a vacuum pump. A wedge-

shaped piece was cut out of the dry filter, mounted on an aluminum stub and sputter coated

with gold. High-resolution images were taken from the tip to the margin of the filter wedge on

a SEM (LEO 1530 Gemini FESEM) and subsequently examined on a qualitative and

quantitative basis. At least 300 specimens per sample were identified following the taxonomic

descriptions of Kleijne (1993), Jordan and Kleijne (1994) and Young et al. (2003). The

absolute abundance of the calcareous nannofossils (number of nannofossils per gram dry

sediment) was calculated with the following equation (e.g., Andruleit, 1996):

AA = (F x C x S) / (A x W)

where AA is the absolute abundance (number g-1

), F is the total sediment coated filter area

(mm2), C is the number of counted nannofossils, S is the split factor, A is the investigated

filter area (mm2), and W is the weight of the dry sample (g).

In addition the accumulation rates of calcareous nannofossils were calculated in the

following way:

AR = AA x SR x DD

where AR is the accumulation rate of the nannofossils (number cm-2

kyr-1

), AA is the absolute

abundance (number g-1

), SR is the sedimentation rate (cm kyr-1

), and DD is the dry bulk

density (g cm-3

).

The age-depth model (Figure 2.2) and the geochemical data (Figures 2.2, 2.3, and 2.4) are

taken from Arz et al. (2003a, 2003b). The paleotemperature reconstruction is based on the

alkenone unsaturation index UK 37 as defined by Prahl et al. (1988). Paleosalinities were

estimated from the local Red Sea !18

Ow-salinity relationship using the paleotemperature

equation of Bemis et al. (1998), the alkenone temperatures, and the stable oxygen isotope

measurements of planktic foraminifera and pteropoda to calculate the !18

Ow. Sediment

material and SEM slides are housed at the Department of Geology, Mineralogy and

Geophysics, Ruhr University Bochum.


22

2.3 Results

The absolute abundances of calcareous nannofossils show a clear pattern (Figure 2.3e), with

low numbers during the Glacial (mean: 2.8E+09 number g-1

) and high numbers during the

Glacial-Holocene transition (mean: 1.3E+10 number g-1

) and the Holocene (mean: 2.4E+10

number g-1

). Our interpretations of the environmental conditions over the last 22 kyr are based

on the following species, which represent over 85% of the total nannofossil composition

(Figure 2.4).

Emiliania huxleyi shows relative abundances between 1.4 and 53.1% (mean: 24.8%) and

is the dominant species of the upper photic community since the end of the Last Glacial

Maximum (LGM). The abundance of this species increases during the Heinrich event 1 (H1),

the Bølling-Allerød and the early Holocene to mid-Holocene interval.

Figure 2.2. (a) Age-depth relation, (b) content of total organic carbon (TOC), and (c) carbonate

content of the sediment core GeoB 5844-2 (note reversed axis for TOC). Shaded vertical bars show

the Red Sea sapropels RS1a and RS1b (Arz et al., 2003a).


23

Figure 2.3. Comparison of (a) the stable oxygen isotope record (Stuiver and Grootes, 2000), (b) the

methane concentration of the GISP2 ice core (Blunier and Brook, 2001), (c) the record of warm

water foraminifera from the Aegean Sea (Rohling et al., 2002), (d) the UK’37 temperatures (core

GeoB 5844-2), (e) the absolute abundances (AA) of the nannofossils and the accumulation rates

(AR) of the nannofossils (core GeoB 5844-2), and (f) the sedimentation rates of core GeoB 5844-2.

Major hydrological and climatic instabilities observed in the tropics and subtropics of Africa are

marked with HC1, HC2, and HC3 (Gasse, 2000).


24

Figure 2.4


25

Figure 2.4. Comparison of the relative abundances of the calcareous nannofossils and the

geochemical data (Arz el al., 2003a, 2003b) of core GeoB 5844-2: (a) relative abundance of

Umbilicosphaera spp. (left axis) and H. carteri (right axis), (b) relative abundance of G. ericsonii,

(c) UK’37 temperatures, (d) relative abundance of G. oceanica, (e) vertical gradients in !18O

("!18Op-b, difference between planktic and benthic records; "!18Op-p, difference between shallow

planktic and deep planktic records), (f) relative abundance of E. huxleyi, (g) reconstructed sea

surface salinities, and (h) relative abundance of the lower photic community (F. profunda and G.

flabellatus). Shaded vertical bars show the Bølling-Allerød, the Younger Dryas transition, and the

Holocene humid period. The qualitative distribution of diatoms in the sediments is sketched at the top of the diagram.

Gephyrocapsa ericsonii exhibits relative abundances from 2.1 to 43.0% (mean: 10.7%),

with highest values during the LGM. The values of Gephyrocapsa oceanica fluctuate between

0.0 and 9.6% (mean: 2.6%). A prominent maximum of this species occurs in the Younger

Dryas (YD) interval. The percentages of Florisphaera profunda and Gladiolithus flabellatus

show fluctuations from 16.0 to 55.3% (mean: 37.3%) and 4.0 to 24.7% (mean: 11.4%),

respectively. Both species show maximum values after the H1, in the earliest Holocene and in

the late Holocene.

In addition we included the following less common taxa in our interpretation (Figure

2.2). Umbilicosphaera spp. (U. sibogae and U. foliosa) are rare with values between 0.0 and

4.4% (mean: 0.7%) showing a significant increase in the Holocene. Helicosphaera carteri

fluctuates between 0.0 and 7.4% (mean: 0.8%), with a prominent maximum in the lowermost

part of the investigated interval.

The calcareous nannofossils from core GeoB 5844-2 show an overall good to moderate

preservation during the Glacial-Holocene transition and the Holocene. Very good

preservation was observed in the Bølling-Allerød interval. Slightly overgrown specimens

occur in most of the investigated samples but do not hamper the investigation. In general,

overgrowth becomes more dominant in carbonate rich intervals. Samples with a different or

unusual preservation are discussed below.

2.4 Discussion

Our detailed analysis begins around 22 ka within the latest Glacial, a period of high aridity in

the northern African region (Adamson et al., 1980; Gasse et al., 1990). In the landlocked Red

Sea, the latest Glacial coincides with extreme environmental conditions. A hypersaline and

hostile environment that culminated in the LGM is documented by the amplification of the

oxygen isotope values and the absence of planktic foraminifera (e.g., Locke and Thunell,


26

1988; Hemleben et al., 1996; Rohling et al., 1998; Fenton et al., 2000, and references therein).

Characteristic components of northern Red Sea sediments during the last Glacial are abundant

large central diatoms (Almogi-Labin, 1982; Geiselhart, 1998). In core GeoB 5844-2 such

diatoms are concentrated in the lower part of the investigated interval. Their content decreases

after the LGM becoming absent after the H1 (Figure 2.4). According to Almogi-Labin (1982),

the monospecific Creseis acicula pteropod assemblages and their association with rich diatom

floras in the northern Red Sea indicate hypersaline conditions and high productive waters.

Geiselhart (1998) argued that high !13

C values and mass occurrences of the diatom

Cosinodiscus oculus-iridis indicate high-fertility waters and so upwelling in the northern Red

Sea. A different ecological interpretation of this species was, however, proposed by Kemp et

al. (2000) on the basis of samples from the Gulf of California and the eastern Mediterranean.

They argue that the presence of some large Cosinodiscus diatoms, including C. oculus-iridis,

characterize a stable phytoplankton community under oligotrophic conditions. The occurrence

of Cosinodiscus oculus-iridis is thought to be associated with strongly stratified waters of the

summer to early fall, whereas their sedimentation takes place after the break down of the

stratification in late fall (‘‘fall dump’’; Kemp et al., 2000).

2.4.1 Last Glacial Maximum (LGM, circa 22–19 ka)

In the northern Red Sea the LGM is characterized by high relative abundances of G. ericsonii

along with F. profunda and G. flabellatus, lower photic zone species (Figures 2.4b and 2.4h).

Furthermore higher values of H. carteri are limited to this period (Figure 2.4a).

Emiliania huxleyi, G. ericsonii and G. oceanica are the most common species of the

upper photic community observed in core GeoB 5844-2. In general, calcareous

nannoplankton tend to be dominant under stable, oligotrophic conditions (K selection). Some

taxa, especially E. huxleyi, and to a lesser extent Gephyrocapsa spp., show, however, an

opportunistic character (r selection). In recent oceans, E. huxleyi shows the widest

distribution, while G. ericsonii and G. oceanica have limited ranges. Emiliania huxleyi is a

eurytopic species which tolerates a wider range of environmental conditions than most other

calcareous nannoplankton species. It has the largest temperature range (1°–30°C), grows

under varying nutrient levels (eutrophic to oligotrophic) and reproduces more rapidly

whenever favorable conditions prevail (e.g., Birkenes and Braarud, 1952; Berge, 1962;

McIntyre et al., 1970; Okada and Honjo, 1973; Brand, 1994). This does not imply that G.

oceanica or G. ericsonii are displaced by E. huxleyi in any situation. Generally G. oceanica is

thought to prefer warm surface waters (optimum: 26°–32°C) and fertile conditions; it


27

dominates upwelling areas and marginal seas (Okada and Honjo, 1975; Kleijne, 1993).

Observations of G. ericsonii suggest that this species also prefers productive waters but under

cooler conditions (optimum: 17°–25°C) (Roth and Coulbourn, 1982; Okada and Wells, 1997;

Rogalla, 2002). In the hypersaline Gulf of Aqaba, G. ericsonii is observed in high numbers

within the low diversity living plankton community during the short fertile winter period

(Winter et al., 1979). High relative abundances of G. ericsonii point to fertile conditions

during the LGM, which may characterize the ‘‘winter situation’’. The isolation from the

Indian Ocean and the hypersaline conditions point to an indigenous population adapted to

high salinities that evolved during the glacial in the Red Sea. Once that normal saline

conditions were reestablished in the Red Sea G. ericsonii became less abundant. It is,

however, still common in the hypersaline Gulf of Aqaba (Winter et al., 1979).

In contrast to the Gulf of Aqaba, where no strong vertical differentiation of the

nannoplankton community is developed, the Red Sea shows a clear zonation. This is verified

by the occurrence of F. profunda and G. flabellatus. The lack of a zonation in the Gulf of

Aqaba seems to be closely related to the specific conditions with a weaker stratification and a

more uniform water column (cf. Weikert, 1987; Grossart and Simon, 2002). Florisphaera

profunda and G. flabellatus exclusively inhabit the lower photic zone (between 100–200 m),

which is characterized by low light levels and cold temperatures but a high nutrient content

compared with the upper photic zone (Okada and Honjo, 1973; Young, 1994). Abundance

variations of this ‘‘lower photic community’’ are closely related to changes of the

hydrographic setting. In general, high relative abundances of these species indicate

oligotrophic conditions in the upper photic zone and intense stratification of the water

column. High surface productivity or a deep mixing of the water column inhibits their growth

(e.g., Young, 1994). F. profunda has been successfully used as a proxy for monitoring the

depth and the stability of the nutricline (Molfino and McIntyre, 1990) and variations in

paleoproductivity (Rostek et al., 1997; Beaufort et al., 2001; de Garidel-Thoron et al., 2001).

A strong negative correlation has been proposed between primary production and the relative

abundances of F. profunda (cf. Beaufort et al., 1997). High nutrient levels in the surface

waters favor especially the upper photic community. This scenario is displayed in the

abundance pattern by high values of the upper photic species in relation to F. profunda (high

primary production = low relative abundances of F. profunda). Reduction of nutrients in the

surface waters, on the other hand, inhibits the growth of the upper photic community. Under

theses conditions the ratio of F. profunda is relatively high, favored by the higher nutrient

content in the lower photic zone (low primary production = high relative abundances of F.


28

profunda).

In stratified subtropical waters like the northern Red Sea, which are characterized by

oligotrophic conditions, periods of high primary productivity are closely associated with

seasonally increased wind stress and surface cooling that raises the stratification (e.g., Halim,

1984; Weikert, 1987). The breakdown recycles nutrients into the surface waters. This favors

blooms of species of the upper photic zone with a preference for fertile conditions (e.g.,

Levanon-Spanier et al., 1979). In the discussion we will use the term LPC (lower photic

community) for F. profunda and G. flabellatus, and the term UPC (upper photic community)

for all other nannofossils. We take into consideration that the presence of the LPC with F.

profunda and G. flabellatus in core GeoB 5844-2 is closely related to oligotrophic conditions

in the upper photic zone. This reflects a better stratification of the water column, which

characterizes the ‘‘summer situation’’ during the LGM.

High values of H. carteri during the LGM are a significant feature. Similar observations

are documented from the central Red Sea and the Gulf of Aqaba (McIntyre, 1969; Winter,

1982a). McIntyre (1969) argued that high numbers of H. carteri are a result of its high

preservation potential, on the other hand the appearance of solution-prone species during this

period are noted by the same author. The ecological interpretation of H. carteri is not

unambiguous. Some studies point to an affinity of this species to warmer conditions, whereas

other results emphasize their preference to waters of higher fertility (e.g., McIntyre and Be´,

1967; Roth and Berger, 1975; Brand, 1994; Andruleit and Rogalla, 2002; Rogalla, 2002;

Ziveri et al., 2004). We argue that higher values of H. carteri indicate a higher nutrient

availability during the LGM. This is in accordance with the findings of Seeberg-Elverfeldt et

al. (2004), who studied diatoms in laminated sediments of the Shaban Deep (northern Red

Sea). These authors speculate that a better mixing of the water column or a high input of

terrigenous material increased the productivity during the LGM.

The geochemical record indicates that the mixing of the water column and the

productivity increased only slightly compared with the recent conditions in the northern Red

Sea (Arz et al., 2003a). We argue that the nannofossil record displays pronounced seasonal

cycles for the LGM, with prevailing fertile conditions during the winter. In summer, the water

column was oligotrophic and better stratified. This interpretation explains the co-occurrence

of calcareous nannofossils with different ecological affinities and supports the conclusions

drawn from diatom assemblages (Seeberg-Elverfeldt et al., 2004).


29

2.4.2 Heinrich Event 1 (H1, circa 18–15.5 ka)

The time of massive iceberg discharges of ice-rafted debris into the glacial North Atlantic

named ‘‘Heinrich events’’ (e.g., Heinrich, 1988; Bond et al., 1992; Broecker et al., 1992)

correlates with paleoclimatic anomalies at various sites outside the North Atlantic. A decrease

of sea surface temperature has been recorded from the Mediterranean Sea and the Indian

Ocean suggesting a far-reaching impact of the Heinrich events (e.g., Geraga et al., 2000;

Cacho et al., 2001 Kudrass et al., 2001; Schulte and Müller, 2001). A clear shift in the

abundance toward an E. huxleyi–dominated community can be observed during the H1 (18 to

15.5 ka (Bard et al., 2000)) in the northern Red Sea. Highest values of E. huxleyi occur during

the maximum cooling of 16.4°C at around 16.5 ka (Figures 2.4f and 2.4c). The Sea Surface

Temperature (SST) fell below the temperature optimum or even below the temperature

tolerance of certain subtropical-tropical calcareous nannoplankton species (e.g., McIntyre et

al., 1970; Winter et al., 1994). The most likely explanation for the change in the nannofossil

composition is that the climate encouraged the bloom of opportunistic species (r-selection)

with an ability to grow in moderate to cool temperate waters. We suggest that these conditions

especially favored the dominance of E. huxleyi. During H1 the calcareous nannoplankton

community gained a composition corresponding to that of the temperate climate zone in

present oceans (e.g., McIntyre and Be´, 1967; Winter et al., 1994).

An alternative explanation is that relatively wetter conditions prevailed during the H1.

The Sea Surface Salinity (SSS) record of core GeoB 5844-2 shows a minor drop of the

paleosalinity with a peak close to the maximum cooling (Figure 2.4g). The interpretations of

the climatic conditions in the northern Red Sea region are, in particular, on a short-term

timescale not unequivocal (e.g., Street and Grove, 1979; Street-Perrott and Harrison, 1985;

Fenton et al., 2000; Frumkin et al., 2000; Bartov et al., 2002, 2003). Because of low-

resolution sampling and the dating uncertainties of available paleoclimatic records it is often

impossible to distinguish between the Glacial, the LGM and the H1 period. This hinders a

detailed comparison. Most records from the northern African region display dryer and colder

conditions during the Glacial (e.g., Gasse, 2000, and references therein). The Nile was a

highly seasonal river. The sediment composition of the Nile as well as pollen spectra indicate

more open vegetation, adapted to colder and dryer conditions in the areas of its headwaters

(Adamson et al., 1980). Furthermore, lake levels in northern and eastern Africa declined or

even dried out (Adamson et al., 1980; Gasse, 2000). The climatic constellation of the Near

East is still debated controversially (e.g., COHMAP Members, 1988; Stein et al., 1999;

Bartov et al., 2003; Prasad et al., 2004). Paleoclimatic records from Israel, including lake


30

level records and pollen spectra, hint toward wet regional climate with glacial periods

(COHMAP Members, 1988; Frumkin et al., 2000, and references therein). In general, high

lake levels of the paleo-Dead Sea (Lake Lisan) occur during periods of colder global climate,

low lake levels are associated with warmer periods. During the Heinrich events, however,

lake levels dropped (Bartov et al., 2003; Prasad et al., 2004). Bartov et al. (2003) showed a

close connection between the Heinrich events and major lake level drops of the Lake Lisan. It

has been proposed that the cessation of the moisture supply, which caused the precipitation

over the eastern Mediterranean region and the related high lake levels, was the result of colder

conditions in the Mediterranean during the Heinrich events. Prasad et al. (2004) demonstrated

a close relationship between the deposition of Lake Lisan varved sediments, eastern

Mediterranean aridity and high-latitude cooling during marine oxygen isotope stage 2. This

suggests a close coupling between the climate of the high latitudes of the Northern

Hemisphere and the eastern Mediterranean.

On the basis of the accordance of arid conditions in northern African (Adamson et al.,

1980; Gasse et al., 1990; Gasse, 2000) and similar observations published recently from the

Near East (Bartov et al., 2003; Prasad et al., 2004), we favor temperature for explaining the

shift in the nannofossil composition to E. huxleyi–dominated assemblages. A climatic change

toward colder conditions displayed by the drop of the SST favored E. huxleyi blooms in the

northern Red Sea. At the same time the temperature drop may have caused a decrease of the

precipitation in the eastern Mediterranean. This resulted in the lake level drop of Lake Lisan

(Bartov et al., 2003; Prasad et al., 2004). We infer from the changes seen, a strong coupling

between the northern Red Sea and the climate of the higher latitude of the north. Since the

exchange with the Indian Ocean was restricted by the low sea level, a direct influence through

the Strait of Bab-el-Mandeb in the south seems to be less important for the observed

conditions in the northern Red Sea during H1 (Arz et al., 2003a). The increase of E. huxleyi,

the lowering of the SST, the timing of the changes in relation to the North Atlantic records,

and their rapidity suggest an atmospheric transfer of the cooling signal. We argue that this

mechanism was responsible for the simultaneous onset of the climatic signals observed in

both regions.

2.4.3 ‘‘H1-Bølling Transition’’ (circa 15.5–14.6 ka)

A hard layer of lithified carbonates, which contains encrusted pteropod shells and terrigenous

components, marks the end of H1 (Arz et al., 2003a). This horizon is thought to have been

deposited under rising SST and lower SSS (Milliman et al., 1969). Increased abundances of


31

the LPC correspond to the change from the H1 to the Bølling (Figures 2.4h and 2.5) and

suggest a well-developed water stratification and a decreased primary production. This setting

was limited to an interval here named ‘‘H1-Bølling transition’’. This interval covers the end

of the H1 and the beginning of the Bølling-Allerød, it therefore corresponds to the end of the

last glaciation and the early phase of the global sea level rise (Fairbanks et al., 1992). We

assume that during the H1-Bølling transition the Red Sea was characterized by the continued

inflow of normal marine waters from the Indian Ocean and probably a weaker seasonal

contrast. This idea is supported by the strong dependence between changes of the salinity in

the Red Sea and the global sea level rise (e.g., Rohling et al., 1998; Siddall et al., 2003, and

references therein). A stronger stratification of the Red Sea can be expected for this period.

These physical conditions with high nutrient depletion in the upper water column favored the

LPC (e.g., Beaufort et al., 1997; de Garidel-Thoron et al., 2001). The slowly improving water

Figure 2.5. Relationship between the upper photic community (UPC) and the lower photic

community (LPC) and the trends in primary production as indicated by their abundance patterns

(based on Figure 2.4h; note reversed axis). LPC-I and LPC-II mark the peak levels in abundance of

the LPC. The horizontal dashed line displays the mean value of the last 2 kyr (assumed to

correspond approximately with the modern trend; see section 2.4.9). A conclusion focused on the

two-step (I and II) onset of the postglacial humid climate is shown at the top (based on the

nannofossil record, the geochemical data, and the paleoclimatic records from the surrounding

landmasses). Shaded vertical bars show the Bølling-Allerød, the Younger Dryas transition (YD-T;

for details, see section 2.4.6) and the Holocene humid period. Arrows indicate Red Sea sapropels

RS1a and RS1b.


32

conditions postdating the H1 are reflected by a minor but continuous decrease of the SSS

(Figure 2.4g). The SST did not yet reach Bølling-Allerød levels, but increased after the H1

(Figure 2.4c). The "!18

Op–b gradients increased up to present values, documenting a strong

stratification of the water column (Figure 2.4e). The northern Red Sea region was already

dominated by arid climatic conditions during the H1-Bølling transition (cf. Gasse et al., 1990;

Gasse, 2000). In contrast to the previous periods (see sections 4.1 and 4.2) the calcareous

nannofossils are of oligotrophic affinities as displayed by the high values of the LPC. We

propose that the increased inflow of normal saline waters into the Red Sea and a weaker

seasonality following the H1 cooling are a possible explanation.

2.4.4 Bølling-Allerød (circa 14.6–13.1 ka)

A change in the composition of the nannofossil assemblage takes place around 14.5 ka. A

sharp decrease of the LPC is followed by an interval with high values of the UPC dominated

by E. huxleyi (Figures 2.4f, 2.4h, and 2.5). This is paralleled by an increase of absolute

abundances during this interval (Figure 2.3e). Both, the onset of an E. huxleyi–dominated

community and the increase of total abundance of calcareous nannofossils correspond with

the enhanced content of atmospheric methane (circa 14.7 ka) recorded in the Greenland ice

cores and the onset of warmer conditions (circa 14.5 ka) (e.g., Dansgaard et al., 1993; Blunier

et al., 1995; Brook et al., 2000; Stuiver and Grootes, 2000; Blunier and Brook, 2001, and

references therein). According to Arz et al. (2003a), the !18

O gradient increased up to 2.5‰

and indicates, along with the SSS record, a stratified water column and a freshening of the

surface waters in the northern Red Sea (Figures 2.4e and 2.4g).

The nannofossil composition and the geochemical data of the Bølling-Allerød period

indicate stable water stratification. In contrast to the H1-Bølling transition the stable

stratification goes along with a higher productivity of the surface water as recorded by the

increase of the UPC (Figure 2.5). Because of the global sea level rise, Gulf of Aden waters

flow into the Red Sea and caused a density stratification of the water column. An increased

freshwater input from large desert wadis and a more humid climate due to the intensification

of the southwest monsoon further enhanced the development of the pycnocline in the Red Sea

(Locke and Thunell, 1988). The combination of a stable density stratification and higher

phytoplankton production went along with the deposition of ‘‘Red Sea sapropel 1a’’ (RS1a

(Arz et al., 2003a)) with maximum TOC values of 2.9 wt% (Figure 2.2b). The deposition of

TOC-rich sediments cannot directly be linked to an increase in productivity. Changes in the

deep water formation and ventilation (see section 2.1) are a second factor controlling the


33

depositional environment in the Red Sea (cf. Almogi-Labin et al., 1991; Arz et al., 2003a).

The increase of the UPC reflects a rise in surface water productivity, which coincides with the

deposition of RS1a. Stratification and fertilization seem to be linked to one another during the

Bølling-Allerød, probably due to humid climatic conditions. The increase in primary

production, in combination with the stratified water column, results in higher oxygen

consumption and additionally a pronounced decrease of the oxygen content. The dysoxic

conditions are maintained and further strengthened by reduced ventilation of the deeper

waters.

This relationship is supported by abundance variations between epipelagic pteropods

(restricted to the upper 100 m) and mesopelagic pteropods (diel migratory) in the Red Sea

(Almogi-Labin et al., 1998; Schmelzer, 1999). In general, the dominance of epipelagic

pteropods indicates a highly stratified water column (Almogi-Labin et al., 1991) and probably

an enhanced primary production, associated with a decrease of dissolved oxygen in the

intermediate water (reinforced and vertically extended OMZ (see section 1)). Under these

conditions the epipelagic species are favored by a good food supply. Simultaneously, the

strong stratification and the low contents of dissolved oxygen reduce the mesopelagic

pteropods (Almogi-Labin, 1982; Almogi-Labin et al., 1991, 1998). In the central Red Sea an

abundance maximum of epipelagic pteropods can be correlated with the deposition of TOC-

rich sediments (up to 1.65 wt%) during the Bølling-Allerød. Coherence with the prevailing

humid climate due to the reinforced southwest monsoon is found (Almogi-Labin et al., 1998).

A similar abundance maximum of epipelagic pteropods seems to be evident in the northern

Red Sea (cf. Schmelzer, 1999).

Humid conditions during the last 22 kyr are mainly associated with the ‘‘greening’’ of the

Sahara during the early Holocene to mid-Holocene. Various paleoclimatic records show,

however, that the onset of a more humid climate started earlier (e.g., Pachur and Altmann,

1997; deMenocal et al., 2000, and references therein). The onset of humid conditions in

subtropical northern Africa has been linked to the insolation (increased Northern Hemisphere

summer season insolation) forced intensification of the African Monsoon, associated with the

northward migration of the Intertropical Convergence Zone (ITCZ). Furthermore, the

comparison of the developing of the insolation (during the last 22 kyr, characterized by a

gradual increase toward a peak around 11 ka and a subsequent gradual decrease (cf. Berger,

1978; Berger and Loutre, 1991)) with paleoclimatic records indicate that the insolation alone

cannot explain the observed abrupt onset and termination as well as short-term variations of

the humid climate (e.g., COHMAP Members, 1988; Gasse and Van Campo, 1994; Kutzbach


34

et al., 1996; deMenocal et al., 2000). Various models try to explain the timing and the

intensity of the African Monsoon in response to different amplification mechanisms such as

changes in land surface conditions (e.g., vegetation cover, albedo) or the ocean surface

(deMenocal et al., 2000, and references therein). Apart from local differences the ‘‘African

humid period’’ covers the time interval from the onset of the Bølling-Allerød until the mid-

Holocene (e.g., 14.8 to 5.5 ka (deMenocal et al., 2000)). It is punctuated by an arid period

corresponding to the Younger Dryas. The two-step arid-humid transition is paralleled by

deglacial-warming events of the Greenland ice core, where the first step of the arid-humid

transition coincides with the onset of the Bølling-Allerød (Gasse et al., 1990).

Considering the geographic location of core GeoB 5844-2, which is beyond the direct

influences of the African Monsoon domain, observations from the Mediterranean borderlands

provide evidence for a regional increase in humidity. Records from the northern Sahara and

the eastern Mediterranean suggest high moisture levels expressed by the occurrence of

challenging plant species (warm and humid requirement), high water tables and an increase in

productivity probably provided by precipitation of Mediterranean origin (cf. Gasse et al.,

1990; Rossignol-Strick, 1995; Gasse, 2000, Frumkin et al., 2000). Therefore a Mediterranean

influence must be taken into account for the northern Red Sea at least additionally.

Considering the proximity of core GeoB 5844-2 to land, a nutrient supply caused by

increased precipitation and local wadi runoff may have affected the nannoplankton

community. We assume, that in addition to the former explanation, nutrient enrichment could

be already associated with an increased deepening of the surface mixed layer or enhanced

local upwelling. Apart from the overall seasonal trend in fertilization (see section 1), locally

restricted mixing processes are known to occur sporadically in the present northern Red Sea

during summer (Weikert, 1987). Such eutrophication could have reinforced during the

Bølling-Allerød period. This scenario may have resulted from a higher wind intensity

(consequence of the rapid warming, land–sea gradient) probably related to the air masses,

which caused the precipitation.

2.4.5 Younger Dryas (YD, 13.1 to 11.7 ka)

The transition to the Holocene is punctuated by the YD (Ruddiman and McIntyre, 1981;

Dansgaard et al., 1989). The YD is associated with an increase of aridity in subtropical Africa

(Gasse et al., 1990; Roberts et al., 1993; deMenocal et al., 2000) and in the Near East

(Frumkin et al., 2000). In the northern Red Sea the YD is displayed by a significant increase

of G. oceanica (Figure 2.4d). Within the E. huxleyi–Gephyrocapsa spp. group, a well-mixed


35

water column, that resulted in a higher nutrient content of the surface water, favored

Gephyrocapsa spp. and especially G. oceanica. This feature provides that temperature (see

section 4.4) was not a limiting factor (e.g., Winter et al., 1979; Brand, 1994; Okada and

Wells, 1997).

We argue that higher proportions of G. oceanica point to an increased mixing of the

water column during the YD. A decrease of the oxygen isotope ratios, which suggests a

deepening of the mixed layer, supports this assertion. Figure 2.4e shows the measured vertical

gradients in "!18

O. The reduction of the "!18

O values (especially "!18

Op–p) obviously

corresponds to the increase of G. oceanica. A shift in the nannoplankton community due to a

slight increase in SSS (Figure 2.4g) is rather unlikely since a preference to saltier waters of G.

oceanica was not observed during previous intervals of similar or even higher salinity (LGM,

H1). Having the warm water affinity of G. oceanica in mind, it is noteworthy that the SST

decreased, however, not reaching LGM or even H1 levels. Data from the alkenone

paleothermometry point to a maximum cooling of 23.7°C around 11.8 ka (Figure 2.4c). It is

necessary, however, to take into account that E. huxleyi already dominated the UPC and, that

the LPC provided an important component of the nannofossil assemblage (Figures 2.4f and

2.4h). The present hydrographic conditions of the northern Red Sea show higher annual

variability than other parts. Particularly during the short winter period the Red Sea is strongly

influenced by cold northerly winds (e.g., Morcos, 1970; Edwards, 1987). These conditions

cause a cooling-induced overturning of the waters (Edwards, 1987, Cember, 1988; Woelk and

Quadfasel, 1996) responsible for a fertilization of the photic zone (e.g., Levanon-Spanier et

al., 1979; Winter et al., 1979). It has been demonstrated that this break down of the

stratification had a lasting impact on the plankton communities (Levanon-Spanier et al., 1979;

Winter et al., 1979; Halim, 1984; Weikert, 1987). In the present Gulf of Aqaba the seasonal-

depth distribution of the nutrient concentration is the limiting factor in primary production

during the summer period. With the beginning of ‘‘winter’’ mixing nutrients are supplied to

the surface waters followed by phytoplankton blooms (Levanon-Spanier et al., 1979). Winter

et al. (1979) described important annual differences in the nannoplankton distribution in the

present Gulf of Aqaba and demonstrated that the appearanceof G. oceanica is mainly limited

to the winter period. According to Winter (1982a), a differentiation of the abundance pattern

between G. oceanica and E. huxleyi occurs with respect to the stratification, with G. oceanica

favoring better mixed waters.

Considering these observations, we assume that the increase of G. oceanica in core GeoB

5844-2 was associated with a reinforcement of the winter mixing during the YD. This view is


36

supported by diatom assemblage compositions studied by Seeberg-Elverfeldt et al. (2004).

They interpreted the shift from a Rhizosolenia-dominated assemblage to a Chaetoceros-

Bacteriastrum association as a change from the well-stratified waters of the Bølling-Allerød

period to the well-mixed surface waters of the YD. The Chaetoceros-Bacteriastrum

association of the YD displays a higher winter production as a result of an increased mixing

during this season (Seeberg-Elverfeldt et al., 2004).

The reinforcement of the winter mixing during the YD may also explain the already high

values of E. huxleyi alongside with the LPC. Subtropical conditions with a better stratification

of the waters already prevailed during the remaining year.

2.4.6 Younger Dryas Transition (circa 11.7–11.4 ka)

The end of the YD is marked by a decrease of G. oceanica (Figure 2.4d), a subsequent

slightly increase of E. huxleyi (Figure 2.4f) and the occurrence of Umbilicosphaera spp.

(mainly U. sibogae). In the geochemical results, a freshening peak centered around 11.5 ka

and an increase of the oxygen isotope ratios can be observed (Figures 2.4g and 2.4e). The

nannofossil composition and the geochemical results point to a better stratification of the

water column and a good availability of nutrients in the upper photic zone. Furthermore, the

presence of the thermophile U. sibogae, which was an inherent part of the nannofossil

associations in the Red Sea since the Younger Dryas transition (cf. McIntyre, 1969; Winter,

1982a), indicates warm surface waters (e.g., McIntyre and Be´, 1967; Okada and McIntyre,

1979; Roth, 1994). Within the dating uncertainty, we suggest that this short interval

corresponds to the abrupt warming of the ‘‘Holocene–Younger Dryas transition’’ (Dansgaard

et al., 1989; Taylor et al., 1997; Severinghaus et al., 1998). In northern Africa this interval

coincides with the early second step of the arid-humid transition (Gasse et al., 1990). High-

resolution terrestrial records, which cover this interval, are limited. Gasse et al. (1990)

described the onset of a short humid episode of less than 500 years in the Sahelian zone,

which corresponds to the Younger Dryas transition. Taking the scarce information into

account, we postulate a brief humid episode around 11.5 ka. This assumption is supported by

similar shifts of nannofossils and geochemical data, both in the Younger Dryas transition

(11.7 to 11.4 ka) and the Bølling-Allerød (14.5 to 13.1 ka). Higher values of E. huxleyi

(Figure 2.4f), accompanied by an increase of the oxygen isotope gradients (Figure 2.4e) and a

freshening of the surface waters (Figure 2.4g), can be observed during both periods.


37

2.4.7 Earliest Holocene Period (circa 11.4–9.5 ka)

An increase of abundance of the LPC ended the short E. huxleyi–dominated interval of the

Younger Dryas transition (Figure 2.4). The following earliest Holocene period shows a

remarkable coincidence with those pattern already observed for the transition from the H1 to

the Bølling-Allerød (see section 4.3). Therefore we assume that comparable oceanographic

conditions prevailed during both periods (Figure 2.4h). In addition to Figure 2.4h, Figure 2.5

(based on Figure 2.4h; note reversed axis) shows the relationship between the UPC and the

LPC, supplemented by the conclusions drawn from their abundance patterns. Increased

relative abundances of the LPC can be observed since the mid-Holocene times toward the

present (= oligotrophic and poorly mixed environment of the present northern Red Sea (e.g.,

Weikert, 1987)). Despite this trend, highest relative abundances of the LPC occur during the

already discussed H1-Bølling transition (Figure 2.5, LPC-I) and the earliest Holocene period

(Figure 2.5, LPC-II). A clear dominance of the LPC with maximum values centered at around

10.4 ka indicates stratified and low productive waters.

The reduction of the SSS (Figure 2.4g) was driven by the inflow of normal saline water

from the Gulf of Aden. The constant increase of the SST (Figure 2.4c) and of the "!18

O

values (Figure 2.4e) as well as the lack of G. oceanica (Figure 2.4d) in significant numbers

further suggests that a reinforced overturning of the waters (e.g., winter mixing) can be

excluded (see section 4.5). Because of the inflow from the Gulf of Aden and the low

overturning, a strong stratification developed. Under these conditions the surface waters are

usually impoverished in nutrients and the primary production was low.

Near the peak level in abundance of the LPC (Figure 2.5, LPC-II), a short interval of

higher TOC values (1.1 wt%) named ‘‘Red Sea sapropel 1b’’ (RS1b (Arz et al., 2003a))

occurs. Apart from this excursion (centered around 10.6 ka), no significant fluctuations of the

TOC values were observed (Figure 2.2b). As discussed above (see section 4.4), a relationship

between sapropel formation and high primary production is not obligatory. The productivity

during the RS1b was already low, but the initial rise of the UPC shortly after the deposition

indicates a subsequent increase of fertility (Figure 2.5). We assume that the deposition of the

RS1b points to a short period of reduced ventilation, which marks the turning point in the

climatic evolution. The sapropel formation was probably induced by the climate change that

finally became manifested in the later Holocene humid period (see section 4.8). The

termination of the RS1b deposition and the absence of sapropel sediments during the humid

period is best explained by the increase of deep water formation after the flooding of the Gulf

of Suez around 11–10 ka (Arz et al., 2003a). The initiation of this important mode of deep


38

water formation (contribution of up to 50% to the present deep water (cf. Woelk and

Quadfasel, 1996)) probably allowed efficient ventilation independent of the surface water

conditions.

2.4.8 Early Holocene to Mid-Holocene Humid Period (circa 9.5–6 ka)

A shift in the nannofossil community, similar to that of the Bølling-Allerød (see section 4.4),

can be observed during the earliest Holocene period (see section 4.7). The decrease of the

LPC was followed again by an interval composed of E. huxleyi–dominated upper photic

assemblages that culminated between 9.5 and 6 ka (Figures 2.4h, 2.4f, and 2.5). We suggest

that the nannofossil composition of this interval reflects the ‘‘Red Sea Holocene humid

period’’ (Arz et al., 2003b). The paleosalinity record and the !18

O data indicate a reduced

salinity of the surface waters alongside with strong water stratification (Figures 2.4g and

2.4e). The vertical gradients of the oxygen isotopes increased up to 1.8‰ while the salinity

decreased to a minimum of 37‰. There are several lines of evidence suggesting that the

humidity was high during this period of the Holocene. The humid interval as indicated by the

nannofossils and geochemical data also appears in various records from Africa, Arabia and

the Mediterranean (e.g., Adamson et al., 1980; Rossignol-Strick, 1983; Ritchie et al., 1985;

COHMAP Members, 1988; Hoelzmann et al., 1998). The now extremely arid central and

eastern Sahara was dominated by savannah and steppe (Hoelzmann et al., 1998). The

extensive vegetation provided the food source for abundant large herbivorous mammals (e.g.,

antelopes, gazelles, giraffes, elephants). The areas covered by lakes and wetlands increased.

They were inhabited by crocodiles, hippos and numerous species of fishes (Pachur et al.,

1990). The high density of flora and fauna, as well as the availability of water, allowed human

settlement in what are now arid or even hyperarid regions (Gasse, 2000). Similar data were

obtained from the Arabian Peninsula, which was occupied by steppe (Hoelzmann et al.,

1998). In the eastern Mediterranean Sea, the Holocene humid period coincides with the

deposition of sapropel layer S1, which was linked to heavy freshwater input from the river

Nile (Rossignol-Strick et al., 1982; Rossignol-Strick, 1983), the Black Sea or more local

rivers (Lane-Serff et al., 1997). Rainfalls caused by the insolation forced African Monsoon

filled the headwaters of the Nile in Sudan and Ethiopia (Rossignol-Strick, 1983). The regional

humidity of the eastern Mediterranean, the northern Sahara and also the northern Red Sea,

however, is thought to have been driven by local precipitation. The moisture was provided by

a water source of Mediterranean origin (Gasse et al., 1990; Rohling and Hilgen, 1991; Arz et

al., 2003b). Paleoclimatic data (e.g., speleotherm records from Israel (Bar-Matthews et al.,


39

2000), Dead Sea level (Enzel et al., 2003) and snail oxygen isotopes from the Negev desert

(Goodfriend, 1988, 1990)) from the eastern Mediterranean region suggest that inflowing air

masses from the Mediterranean were the source of precipitation during the early Holocene to

mid-Holocene in the northern Red Sea (Arz et al., 2003b).

On the basis of the correlation between the well documented humid period, the !18

O

gradient, the salinity and the shift in the nannofossil assemblages, we argue that an increased

nutrient supply due to higher precipitation in the northern Red Sea region could have initiated

an increase in productivity. For the mid-Holocene slightly higher productivities, in

combination with an increase of humidity in the northern Red Sea region, are proposed

(Almogi-Labin, 1982; Winter, 1982a). Almogi-Labin (1982) assumed a strengthened

upwelling for the northernmost parts of the Gulf of Aqaba and the Red Sea during the mid-

Holocene. Winter (1982a) proposed that a high runoff in the mid-Holocene resulted in a high

nutrient supply for the Gulf of Aqaba. It is noteworthy that these data were obtained from a

core close to a major wadi fan (Wadi Dahab (cf. Winter, 1982a)). In the Holocene, however,

these runoffs must have been relatively low in the Gulf of Aqaba and high in the Red Sea.

During the LGM this trend was reversed (Winter, 1982a). This explanation is supported by

the general glacial/interglacial trends of the Lake Lisan and the paleoclimatic records from

subtropical Africa (see section 4.2). Both, the runoff from local drainage systems and the

precipitation increased (cf. Hoelzmann et al., 1998), supplying nutrients to the Red Sea. In

addition the inflow of the air masses from the Mediterranean may have enforced the mixing

of the surface waters or local upwelling, supporting a high productivity of the UPC (see

section 4.4).

An alternative explanation for the observed increase in primary production is the

increased inflow of Gulf of Aden waters. These added nutrients to the Red Sea (Weikert,

1987). The primary production in the present northern Red Sea is, however, not directly

related to the nutrient input from the Gulf of Aden. Because of the rapid biological

consumption the influence of the fertilization through the inflow is limited to the southern

Red Sea. It does not reach farther north than 19°–21°N. Differences in primary production,

plankton biomass and abundances are relatively small between the northern and central Red

Sea whereas both parts differ distinctively from the southern Red Sea (cf. Khmeleva, 1970;

Weikert, 1987).

The interpretations based on the calcareous nannofossils from GeoB 5844-2 are in

agreement with various terrestrial paleoclimatic records. The Red Sea record of a two-step

onset of the postglacial humid climate, separated by the YD, corresponds with numerous


40

paleohydrological studies (Gasse, 2000, and references therein). It is apparent, however, that

there are some anomalies in the nannofossils records. Bearing in mind the brief humid

interval of the Younger Dryas transition (11.7 to 11.4 ka), the Holocene humid period (9.5 to

6 ka) cannot be interpreted as a direct continuation of this interval because of its separation by

a more oligotrophic/arid period during the earliest Holocene period (11.7 to 9.5 ka).

Direct comparison between marine and terrestrial records on a short timescale is

problematic. Paleoclimatic signals achieved from land can be affected by the interplay of

climate, vegetation and water cycle (e.g., Gasse and Van Campo, 1994). Abrupt climatic

changes in the terrestrial records are often buffered and delayed by the response time of the

water cycle (Gasse and Van Campo, 1994). Evaporation of lakes, swamps and the vegetation

can result in self-stabilizing regional precipitation regimes that compensated periods of

droughts (Pachur and Altmann, 1997). Furthermore, signals of humid periods can be lost by

erosion (Hoelzmann et al., 1998).

The regional effects of the climatic reorientation during the early Holocene are more

complex, but already detectable by the calcareous nannofossils. We assume that the marine

record and the variations of the nannofossil community are a direct response to abrupt climate

changes.

2.4.9 Mid-Holocene to Late Arid Holocene (circa 6–0.6 ka)

The LPC shows a distinct increase in abundance after the humid period (Figure 2.5), with F.

profunda being the most common species in the late Holocene (Figure 2.4h). We argue that

the F. profunda–dominated assemblages in the northern Red Sea reflect geographically

restricted, oligotrophic conditions, caused by an increased aridity following the Holocene

humid period. Enhanced aridity is recorded in the eastern Sahara after 6 ka and in the whole

Sahara after 4.5 ka. Nowadays conditions were established around 2 ka (e.g., Maley, 1997;

Gasse, 2000).

Apart from the long-term trend, minor fluctuations in the nannofossil composition

occurred during the later part of the Holocene, which are still under investigation. Our current

data suggest four intervals of reduced nannofossil abundances, centered around 8.2, 6.4 and

4.0 and after 2.0 ka (Figure 2.3e). Though the number of samples is limited compared to the

frequency of the fluctuations, it seems that these intervals coincide with major cooling events

in the Aegean Sea (Figure 2.3c). These are thought to be related to variations of the

winter/spring intensity of the Siberian High with a periodicity around 2500 years (Figure

2.3c) (Rohling et al., 2002). Hydrological fluctuations, as compiled from various subtropical


41

African lake level records, are centered around 8.5–7.8, 7.0–6.6 and 4.5–3.5 ka (Figure 2.3)

(Gasse, 2000, and references therein). These seem to correspond to an increased aridity and a

stepwise transition to recent conditions in the northern Red Sea region. The 8.2 and the 4.0

events are the most prominent climate events. The 8.2 event can be correlated with a

distinctive !18

O shift and a decrease of the atmospherical methane as observed in the

Greenland ice core. The 4.0 ka event is thought to be related with the collapse of the Old

World societies (e.g., Gasse, 2000; deMenocal, 2001).

2.5 Conclusions

The calcareous nannofossil assemblages of the northern Red Sea mirror the climatic trends of

the last 22,000 years for this region and are closely coupled to the climatic changes of the

northern mid-latitudes and high latitudes. The changes in the nannofossil composition derived

from core GeoB 5844-2 indicate, that the calcareous nannofossil communities are not only

affected by variations of the salinity. Other factors, especially temperature and nutrients,

control the observed distribution patterns of nannoplankton in the northern Red Sea.

The nannofossil record allows differentiation of the Last Glacial Maximum (LGM) and

the Heinrich event 1 (H1) during the Glacial. Gephyrocapsa ericsonii and Helicosphaera

carteri characterize the LGM, whereas an increase of Emiliania huxleyi marks the H1.

Nannofossil assemblages with common G. ericsonii suggest pronounced seasonality during

the LGM. During the H1 the cooler climate encouraged the dominance of E. huxleyi. In the

Younger Dryas higher values of Gephyrocapsa oceanica are closely related to the intensified

‘‘winter’’ mixing of the water column.

The nannofossil assemblages of both the Bølling-Allerød transition and of the Holocene

show two subsequent, distinctive short-term phases. The early phase is dominated by the

lower photic community (LPC). This indicates a strong stratification of the water column and

oligotrophic conditions (H1-Bølling transition and earliest Holocene period). The later phase

shows a significant shift toward an E. huxleyi–dominated upper photic community that points

to more fertile conditions in the surface waters (Bølling-Allerød and Holocene humid period).

The late Holocene is characterized by a continuous increase of the LPC reflecting the

increase of aridity following the Holocene humid period. Apart from the long-term trend

fluctuations, the nannofossil record suggests a close link to minor climatic variations as

detectable during the Holocene.


42

Acknowledgments

Sample material has been supplied by Jürgen Pätzold (Bremen). The work benefited from

discussions with Sylvia Rückheim and André Bornemann. We thank Rolf Neuser for

technical assistance at the SEM. This research was supported by the Deutsche

Forschungsgemeinschaft (Mu 667/23-1, -2). We thank David Watkins and an anonymous

reviewer for useful comments and improvements of the paper. Jeremy Young spent quite

some time on smoothing the English.


43

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Chapter 3: Nannoplankton successions during the last glaciation…

53

3 Nannoplankton successions in the northern Red Sea during the last glaciation (60 to

14.5 ka): Reactions to climate change

Abstract

Due to its restricted connection with the Indian Ocean, the desert-enclosed Red Sea is

extremely sensitive to global sea level changes and thus ideally suited for paleoceanographic

studies of what occurred during the last glaciation. The understanding of its glacial history is,

however, still limited. A serious obstacle to obtain satisfactory paleoecological information

has been the rarity of microfossil proxy species caused by high salinities. Here, we present a

continuous and well-dated calcareous nannoplankton record from the northern Red Sea,

covering the interval from 60–14.5 ka. Our investigation shows that the composition of the

calcareous nannoplankton community varied between ca. 32 ka and 14.5 ka in response to

rapid environmental changes which are closely correlated to climatic fluctuations described

from the North Atlantic region. Heinrich events H3, H2 and H1 are dominated by Emiliania

huxleyi. Gephyrocapsa oceanica and especially Gephyrocapsa ericsonii are abundant

between H3–H2 and H2–H1. A less pronounced response of the calcareous nannoplankton to

the high latitudinal climatic oscillations is documented prior to 32 ka, suggesting that a strong

atmospheric coupling between the northern Red Sea and the North Atlantic realm was

established in the late Marine Isotope Stage 3. In contrast to the previously held view of a sea

level related salinity increase as the major cause for changes of the plankton communities

within the glacial Red Sea, we interpret the documented variations as being caused by local

hydrographic changes under the atmospheric control from the extratropics. Temperature

changes and especially variations of the water stratification appear to be critical selective

factors for the calcareous nannoplankton composition.

Keywords: Red Sea; calcareous nannoplankton; paleoclimate; glacial environment; Heinrich

events

3.1 Introduction

Continuous microfossil records from the northern Red Sea, which cover Marine Isotope

Stages (MIS) 3 and 2 (e.g., Imbrie et al., 1984), are sparse. As a result of the globally low sea

level during the last glacial (up to around 120–125 m), water exchange between the Red Sea


54

and the Gulf of Aden (Indian Ocean) via the shallow (now around 137 m) and narrow (now

around 18 km) Strait of Bab-el-Mandeb was drastically reduced (Siddall et al., 2003; Arz et

al., 2007). Hypersaline conditions, amounting to 50‰ or even above, prevailed during MIS 3

and especially MIS 2 (Thunell et al., 1988; Hemleben et al., 1996; Arz et al., 2003). The

disappearance of various microfossil species prelude the “aplanktic zone”, a period

characterized by the absence of planktic foraminifera and by a pteropod community mainly

consisting of the euryhaline Creseis acicula (Almogi-Labin, 1982; Winter et al., 1983; Fenton

et al., 2000). Recent results indicate the presence of calcareous nannoplankton (simply named

nannoplankton in the ongoing text) during this hypersaline “aplanktic zone” in the northern

Red Sea (Legge et al., 2006).

Figure 3.1. Location of core GeoB 5844-2 in the northern Red Sea and a schematic view of the precipitation regimes at both margins of the subtropical desert belt (north: Mediterranean winter rain,

south: southwest summer monsoon). Dashed line shows position of the Intertropical Convergence Zone (ITCZ); thick black arrow indicates wintertime mean position of the Subtropical Jet (STJ) which roughly mark the southern limit of the extratropical Mediterranean winter climate features (Air Ministry, 1962; Blackmon et al., 1984; Martyn, 1992; Hoskins and Hodges, 2002). The desert belt is shaded in light gray (vegetation-based; Schmithüsen, 1976). A detail view of the northern Red Sea with the Gulf of Aqaba (right) and the Gulf of Suez (left) is shown in the upper right panel.

We present a nannoplankton record from the desert-enclosed northern Red Sea (Figure 3. 1)

covering the period between around 60 ka and the onset of the Bølling-Allerød warm period

at around 14.5 ka (Figure 3.2 and Figure 3.3). Short-term climatic variations, including

Dansgaard-Oeschger (D-O) cycles and Heinrich events, are well documented in the North

Atlantic region during the time period under consideration (Bond et al., 1993; Dansgaard et

al., 1993). D-O cycles, rapid and large amplitude changes observed in the !18O ice core

records from Greenland, reflect variations in air temperature over the northern ice sheets.

Several D-O cycles are grouped into multi-millennial intervals characterized by an overall

progressive cooling (longer-term cooling cycle). Massive iceberg discharges into the North

Atlantic are indicated by extensive layers of coarse-grained debris (Broecker et al., 1992).

These Heinrich events, or more precisely their climatic consequences are documented as


55

pronounced cold periods that occurred at the end of the longer-term cooling cycles (Broecker

et al., 1992; Bond et al., 1993). D-O cycles, longer-term cooling cycles, and the climatic

consequences of Heinrich events (Figure 3.3) are not restricted to the North Atlantic region.

Various records in the mid-latitudes document similar variations and indicate a strong relation

to the climate of the North Atlantic realm (e.g., Voelker, 2002). Given the sensitivity of

nannoplankton to autoecological factors (e.g., Brand, 1994; Ziveri et al., 2004), our record

offers a unique possibility for the examination of millennial-scale environmental changes

during the hypersaline period of the Red Sea in response to the longer-term cooling cycles

and Heinrich events of the North Atlantic realm.

3.2 Material and Methods

The studied sediment core GeoB 5844-2 (Figure 3.1) was recovered in 1999 during Meteor

cruise 44/3 (Pätzold et al., 2000) in the vicinity of the Sinai Peninsula (27°42´81"N,

34°40´90"E, water depth: 963 m). For this study, we investigated fossil nannoplankton from

the time-interval 60–22 ka. To allow a more comprehensive view of the results, we present a

complete record of the 60–14.5 ka time period by including recently published data (Legge et

al., 2006). Sample preparation follows the filtration technique as described by Andruleit

(1996). At least 300 specimens of each sample were examined under scanning electron

microscope. We describe the abundance patterns of the three Noelaerhabdaceae Emiliania

huxleyi (Figure 3.3f), Gephyrocapsa ericsonii (Figure 3.3g) and Gephyrocapsa oceanica

(Figure 3.3h), which dominate the upper-photic

community of the Red Sea. In addition, we figure the

relative abundance of Florisphaera profunda (Figure

3.5d). We integrate F. profunda in our discussion

because this lower-photic species contributes

significantly to the fossil record. Based on AMS

(Accelerated Mass Spectrometer) 14C dating and

additional paleomagnetic dating points, the

stratigraphy framework of core GeoB 5844-2 (Figure

3.2) has been established by Arz et al. (2007). Ages

(on the SFCP timescale; Shackleton et al., 2004) are

announced in calibrated calendar years before

present.

Figure 3.2. Age depth relation of the sediment core GeoB 5844-2 (Arz et al. 2007). The core section/ time interval studied here is gray shaded.


56

3.3 Results and Discussion

The varying composition of nannoplankton assemblages defines two intervals (60–32 ka and

32–14.5 ka). The most striking feature in the younger interval (32–14.5 ka) is the recurrent

appearance of E. huxleyi, G. ericsonii and G. oceanica. Those intervals, which correspond

stratigraphically to Heinrich events H3, H2 and H1 (Figure 3.3), are dominated by E. huxleyi

(Figure 3.3f). The longer-term cooling cycles, or more precisely the intervals between H3–H2

and H2–H1, are characterized by increasing values of G. ericsonii (Figure 3.3g). In addition,

brief increases of G. oceanica are displayed shortly after H3 and especially after H2 (Figure

3.3h). Fluctuations in the abundance of these three species are also documented in the older

part of the record (60–32 ka), there are, however, no analogue abundance patterns as in the

younger interval. A comparable distinct species succession of G. oceanica, G. ericsonii and E.

huxleyi (order in view of the North Atlantic climate/temperature trend) is virtually absent. In

the following discussion we will focus on two aspects: first, the recurrent sequence of G.

oceanica, G. ericsonii, and E. huxleyi (= Noelaerhabdaceae-succession) in the younger

interval, and second, the timing of the onset of the succession at around 32 ka.

Biogeographical distribution patterns emphasize the affinity of G. oceanica to the warm

waters of the tropical and subtropical latitudes (e.g., Brand, 1994; Ziveri et al., 2004).

Gephyrocapsa ericsonii flourishes under cooler conditions, while cold waters also inhibit its

growth (e.g., Gartner, 1988; Okada and Wells, 1997). Emiliania huxleyi tolerates a broad

temperature range as documented by its abundant occurrence from the tropics to the subarctic

(e.g., Gartner, 1988; Brand, 1994).

The alkenone-sea surface temperatures (SSTs) available from core GeoB 5844-2 (Arz et

al., 2007) show that the Heinrich events H3, H2 and H1 are correlated with pronounced

cooling events in the northern Red Sea (Figure 3.3e). A weaker response to the high-

latitudinal cooling is evident in the older part of the alkenone-record, which also corresponds

well to the nannoplankton record. Prior to around 32 ka, fluctuations were of lower intensity

and the SST always remained above 22.1°C. Due to the good correlation of our

nannoplankton findings with the SSTs, it seems reasonable to envisage a temperature-related

explanation for the Noelaerhabdaceae-succession.

The specific temperature preferences may explain the dominance of E. huxleyi during the

maximum cooling intervals (H3, H2, H1) but leaves us with one problem: It remains open

why this eurythermal species was outcompeted by Gephyrocapsa spp. in the transitional

intervals (H3–H2, H2–H1). Corroborated by plankton studies, we speculate on temperature-

driven variations of the water stratification as the key mechanism.


57

Figure 3.3


58

Figure 3.3. Comparison of (a) the SPECMAP !18O record (stacked; Imbrie et al., 1984) and (b) the !

18O GRIP ice-core record (on the SFCP timescale; Shackleton et al., 2004) with results from core GeoB 5844-2 (northern Red Sea): (c) the planktic !18O record (Arz et al., 2007), (d) the recon-structed sea level curve (two temperature end-member scenarios: with and without taking the full range of alkenone SSTs changes into account; for details please see Arz et al., 2007), (e) the alkenone-sea surface temperatures (Arz et al., 2007), (f) the relative abundance of E. huxleyi, (g) the relative abundance of G. ericsonii, (h) the relative abundance G. oceanica and (i) the absolute abundance (AA) and accumulation rates (AR). The qualitative distribution of diatoms (large Cosinodiscus) is sketched in Figure 3.3i above the absolute nannoplankton values. Age control points of core GeoB 5844-2 are indicated by triangles at the bottom (for details see Figure 3.2). The trend of the longer-termed cooling cycles and the position of the Heinrich events (H) are sketched at the top. Interstadials in !18O GRIP ice-core record are numbered (1–17). The extent of the “aplanktic zone” (salinity !49–50‰) is sketched at the bottom of the diagram. Gray shaded vertical bars show intervals of major SST-drops and E. huxleyi-peaks, which correspond to H3, H2 and H1.

The present-day nannoplankton community in the northern Red Sea region, overall

dominated by E. huxleyi (within the Noelaerhabdaceae; Winter, 1982), is controlled by the

stable stratification of the waters resulting in oligotrophic conditions (regenerated

production). Only during the short winter period does the breakdown of stratification allow an

upwards-mixing of nutrient-rich subsurface waters. The result of this winter mixing is an

elevated nutrient level in the photic zone, which favors new production (Levanon-Spanier, et

al., 1979; Weikert, 1987). The hydrographic changes are initiated (i.e. the formation of dense

surface waters) by the invasion of cold air from northerly direction. Data from the Gulf of

Aqaba, the northernmost extension of the Red Sea (Figure 3.1) shows high abundances of G.

ericsonii and G. oceanica associated with the deeply mixed waters of the winter period

(Winter et al., 1979). Independent evidence from plankton studies of other algal groups

indicates that the structure of the phytoplankton community in the Gulf of Aqaba is strongly

affected by the deep (600 m water depth or more) convective mixing in winter (e.g., Levanon-

Spanier, et al., 1979; Lindell and Post, 1995; Post, 2005), whereas generally a year-round

stable stratification (mixing limited to the surface waters) characterizes the open northern Red

Sea. Another useful modern analogue for the G. ericsonii-dominated intervals in core GeoB

5844-2 can be observed in the northern parts of the western Mediterranean Sea, the Gulf of

Lion, and adjoining areas up to the Balearic Islands. The comparison of plankton data

(Knappertsbusch, 1993) and hydrological observations (e.g., Send et al., 1999) reveals a

possible relationship between abundance maxima of G. ericsonii and the overturning of the

waters caused by the strong katabatic mistral-wind in winter.

We argue that a deep convective mixing due to the surface buoyancy loss favors

especially Gephyrocapsa spp. We infer that the increase of G. ericsonii in core GeoB 5844-2


59

reflects significant hydrographic changes driven by the increased strength and influence of

cold northern winds in the northern Red Sea region. Still relative high SSTs during the early

parts of the cooling cycles may account for the brief increases of G. oceanica.

Evidence for the above-described scenario may also comes from planktic diatoms. The

deposition of opal-rich sediments, characterized by abundant large Centrales of the genus

Cosinodiscus (nearly solely Cosinodiscus oculus-iridis) and low absolute values of

nannoplankton (Figure 3.3i), coincide roughly with the Gephyrocapsa spp. dominated

intervals. The mass occurrence of C. oculus-iridis in marine sediments is thought to be the

result of the “fall-dump” (Kemp et al., 2000), a sedimentation-event initiated by the onset of

water mixing in late fall–early winter. It was probably this breakdown of stratification that led

later in the cold season to the increase of the Gephyrocapsa spp.

Though confronted with the unique glacial setting of the Red Sea and the prevailing

“aplanktic zone” (for planktic foraminifera), the nannoplankton abundance patterns are

probably not adequately explained by a salinity increase. The mechanisms described above

provide a plausible explanation for the Noelaerhabdaceae-succession where the glacial

salinity increase is only of secondary importance. Because there is rough accordance between

the absolute abundance patterns (Figure 3.3i) and the long-term trends, documented for

different parameters (e.g., sea level, salinity, temperature), it is difficult to relate the absolute

values to one specific factor. Similarities with the sea level curve (Figure 3.3d; Arz et al.,

2007) could lead to the assumption that the salinity played a crucial role. On the other hand, it

seems to be unlikely that a salinity-related reduction of the absolute values have occurred

without a marked imprint in the species composition. If nannoplankton was affected by the

salinity increase, one could anticipate a gradual disappearance of species (i.e. a non-repeating

species succession somehow similar as documented for planktic foraminifera species). Such

patterns are, however, not recognizable in our record. Instead we observed the reiterated

Noelaerhabdaceae-successions, which are closely correlated with the longer-term cooling

cycles and Heinrich events of the North Atlantic realm. Corroborating evidence against a

salinity-related approach to explain the Noelaerhabdaceae-succession in the northern Red Sea

may come from the Mediterranean. Nannoplankton data from the western Mediterranean

(Colmenero-Hidalgo et al., 2004), although not identical in detail, show comparable patterns

to those from the northern Red Sea. Abundance-peaks of E. huxleyi during Heinrich events

and high values of small Gephyrocapsa between the Heinrich events have been observed

(Colmenero-Hidalgo et al., 2004). It is worth mentioning that even other micro-

paleontological studies undertaken in the northern Red Sea region question a simple salinity-


60

related approach to explain all shifts within the planktic foraminifera and pteropod

communities. Temperature variations, changes of the water stratification and the availability

of food (i.e. productivity of the surface waters) seem to be highly important even in glacial

times (Almogi-Labin, 1982; Fenton et al., 2000).

Moreover, it should be recalled that the trend towards hypersaline conditions mirrors a

diminishing communication with the open ocean (e.g., Siddall et al., 2003; Arz et al., 2007).

Hence, the imprints of climatic signals transmitted from the Gulf of Aden to the Red Sea

should be weaker during full glacial conditions. Due to the great distance from the Strait of

Bab-el-Mandeb (nearly 2000 km), this should be especially true for our study area. The

Noelaerhabdaceae-succession, by contrast, points to a sensitive and fast acting mechanism

linking high latitudinal climate oscillations and environmental changes in the northern Red

Sea region especially after 32 ka.

A plausible mechanism fulfilling the above requirements is the atmospheric

teleconnection with the extratropics and the associated hydrographic changes driven by an, in

relation to early–mid MIS 3 conditions, reinforced influence of cold air from the mid-

latitudes. Data from the North Atlantic (e.g., Shackleton et al., 2000; Figure 3.4a, planktic

!18O record from the Iberian Margin) and the Mediterranean Sea (e.g., Cacho et al., 1999; Fig.

Figure 3.4. Comparison of (a) the North Atlantic planktic !18O record (Shackleton et al., 2000) and (b) the Western Mediterranean alkenone-sea surface temperatures (Cacho et al., 1999) with (c) the alkenone-sea surface temperatures from the northern Red Sea (Arz et al., 2007). The onset/range of the Noelaerhabdaceae-succession is sketched at the bottom of the diagram.


61

3.4b, alkenone record from the Alboran Sea) reveal large and rapid environmental

fluctuations in response to the high latitudinal cooling already before 32 ka. Considering the

onset of the succession at 32 ka and the alkenone-SST variations (Figure 3.4c), a stronger

atmospheric connection between the northern Red Sea and the extratropics was established in

the late MIS 3. The present-day Mediterranean winter climate (Figure 3.1) is affected, in

addition to the westerlies responsible for the characteristic winter cyclonic rain, by cold

continental and altered maritime polar and arctic air reaching the region via southern and

eastern Europe (e.g., Air Ministry, 1962; Martyn, 1992; Hoskins and Hodges, 2002). The

onset of the succession at around 32 ka in the northern Red Sea may reflect an expansion of

the extratropical winter climate features into the subtropical desert belt (a wintertime mean

position of the “Mediterranean front” further south; Air Ministry, 1962). This shift and

accordingly the more direct and more prolonged influence of the extratropical circulation

systems on the northern Red Sea may account for the now undamped atmospheric transfer of

the high latitudinal cooling signals to the northern Red Sea as indicated by the nannoplankton

abundance patterns and the alkenone data.

Terrestrial records from the northern parts of the Saharan desert belt provide a good link

between a climatic change towards cooler and wetter conditions (respectively a positive

precipitation-evaporation balance) and the changes of the nannoplankton community

structure. The onset of a distinctive “pluvial period” (in absolute terms probably semi-arid or

at most semi-humid in large part of the present desert belt) is documented at around 32–30 ka

(e.g., Nicholson and Flohn, 1980; Servant and Servant-Vildary, 1980; Rognon, 1996;

Hoelzmann et al., 2004). The increased atmospheric pressure gradient, associated with the

expansion of the northern ice sheets and the enhanced latitudinal temperature contrast,

modulated the atmospheric circulation. A southward shift of the westerlies that allowed

Atlantic depression to supply winter-rain towards the desert belt has been proposed

(Nicholson and Flohn, 1980; Hoelzmann et al., 2004).

This late Pleistocene pluvial period ended at around 22–20 ka with the onset of arid or

even hyper-arid climatic conditions, which characterize the Last Glacial Maximum (LGM;

22–19 ka) and the subsequent interval up to the onset of the Bølling-Allerød (at around 14.5

ka) in wide areas of Africa and Arabia (e.g., Nicholson and Flohn, 1980; Hamilton and

Taylor, 1991; Gasse, 2000; Hoelzmann et al., 2004). The climatic re-orientation at around 22

ka did not disturb the continuation of the Noelaerhabdaceae-succession (Figure 3.3). On

closer examination, however, it seems to be manifested in the nannoplankton record by the

secular trend of F. profunda (Figure 3.5d). Apart from a moderate decrease at around 21 ka,


62

overall high values of F. profunda since around 22 ka indicate low paleoproductivities (cf.

Beaufort et al., 1997). The highest primary production/lowest F. profunda-abundance is

documented between 32 ka and 22 ka.

The coincidence (in time) between the interval of higher productivity in the northern Red

Sea and the subtropical pluvial period on land (Figure 3.5) suggests a linkage between the

moisture climate and the trophic state of the northern Red Sea waters. Fenton et al. (2000)

propose that the glacial conditions were milder in the Gulf of Aqaba than in the Red Sea.

Unlike modern conditions, a freshwater inflow to the Gulf of Aqaba supplied by

Mediterranean depressions is assumed (Fenton et al. 2000). Higher water levels of the Lake

Lisan (Figure 3.5c), the progenitor of the Dead Sea, may give evidence for higher regional

Figure 3.5. Comparison of the paleoproductivity in the northern Red Sea and the changes in regional precipitation/water balance: (a) variations of the southern desert margin (Reichelt et al., 1992), (b) Chad basin precipitation/evaporation variability (Servant and Servant-Vildary, 1980), (c) Lake level record (in meter below mean sea level) of Lake Lisan (Bartov et al., 2003) and (d) the relative abundance of F. profunda (note reversed axis). Shaded vertical bar indicate the subtropical pluvial (ca. 32 to 22-20 ka) as documented by terrestrial proxy records in the subtropics of Africa (e.g., Nicholson and Flohn, 1980; Rognon, 1996; Hoelzmann et al., 2004). The arrows in Figure 3.5c mark lake level drops of Lake Lisan associated with Heinrich events H5–H1 (Bartov et al., 2003). The paleoflood-record (interval of extreme floods) from the Negev desert (Greenbaum et al., 2006) is sketched in Figure 3.5c together with the Lake Lisan curve.


63

rainfall (Bartov et al., 2002; 2003). There are similarities between fluctuations of the

reconstructed lake level and the F. profunda abundance patterns/productivity changes on

longer time scale, which may indicate an influence of extratropical precipitation on northern

Red Sea fertility. Another possibility is that the African monsoon (Figure 3.1), which affected

the subtropical desert belt (Figure 3.4a) during the late Pleistocene from the south (Rossignol-

Strick, 1983; Reichelt et al., 1992; Rognon, 1996), influenced marine fertility. According to

Almogi-Labin et al. (1998), the late Pleistocene humid period is documented in the central

Red Sea by abundance maxima of epipelagic pteropod species, which should indicate a

reduced ventilation of the surface waters and an increase in productivity caused by a strong

southwest monsoon. After all, any explanation linking Red Sea fertility and extratropical or

tropical rainfall remains debatable. The high productivity period in the northern Red Sea

cannot be linked to the African monsoon without problems; a strong direct influence of the

southwest monsoon should be limited to the southern and central Red Sea. Furthermore,

differences in the spatial rainfall in the Near East (e.g., Horowitz 1979; Amit et al., 2006;

Dayan and Morin, 2006) make it difficult to compare wetter periods in the northern

subhumid/semiarid parts (e.g., as deduced from the Lake Lisan record) with environmental

changes further south (southern Negev-Gulf of Aqaba-northern Red Sea region). Oxygen

isotopic data from the northern Red Sea (Figure 3.3c), overall characterized by a substantial

salinity overprint, reveal no distinctive period of surface freshening between 32 ka and 22 ka.

This does not necessarily exclude a local restricted fresh water supply (e.g., from wadis), but

wind-driven variations of the water column properties were probably more important. Due to

a wetter atmosphere and the related decrease of thermal stability, stronger winds/a higher

frequency of storms, may have caused stronger upwelling/mixing of ocean waters. This

speculative approach offers an alternative mechanism that explains the high productivity

period in accordance with the lack of a distinct freshening signal in the isotopic data.

Evidence of increased storminess under overall still dry regional conditions is documented by

a recently published paleoflood-record from the hyperarid southern Negev desert (Greenbaum

et al., 2006). A conspicuously higher frequency of storms/paleofloods is indicated during the

late MIS 3–early MIS 2 (Figure 3.4). Interestingly, the authors (Greenbaum et al., 2006)

argued for variations of the Red Sea Trough systems (Kahana et al., 2002; for details) as a

possible cause. This implies a profound role of the tropical convection intensity and hence of

the African monsoon system (cf. Kahana et al., 2002; Dayan and Morin, 2006).

The information available so far only allow speculative statements. Although we note a

high productivity period in the northern Red Sea contemporaneous with the pluvial period on


64

land, further research is needed to understand the causal relationship between F. profunda

abundance patterns (and deduced primary production) and wetness documented by terrestrial

proxy records.

3.4 Conclusion

Our record demonstrates that the nannoplankton responded differently to high latitudinal

cooling before and after 32 ka. A strong atmospheric influence from the extratropics is

indicated since the late MIS 3. The climatic evolution of the northern subtropics lends support

to this finding. A reiterated succession, closely correlated with the longer-term cooling cycles

and Heinrich events of the North Atlantic realm, is detected after 32 ka. The three Heinrich

events H3, H2, and H1 are dominated by Emiliania huxleyi. Gephyrocapsa oceanica and

especially Gephyrocapsa ericsonii are abundant between H3 and H2 and between H2 and H1.

The present data led to the assumption that the glacial salinity increase did not cause the

nannoplankton successions. Although the glacial Red Sea represents a unique marine

environment with regard to the impressive sea level driven salinity increase, the documented

nannoplankton changes are probably better explained by a more unspectacular approach. A

plausible mechanism seems to be the atmospheric teleconnection with the extratropics and the

associated hydrographic changes driven by an, in relation to present-day and early–mid MIS 3

conditions, reinforced influence of cold air from northerly direction. Furthermore, we

documented a low F. profunda/high productivity interval contemporaneous with a pluvial

period on land between 32 and 22 ka. This finding raises further questions with respect to the

specific mechanism linking F. profunda abundance patterns, marine productivity and

terrestrial environmental changes. Although we cannot present a final explanation for this

relationship here, it may provide an interesting approach for further studies to investigate past

climatic changes in the northern Red Sea region.

Acknowledgements

We thank Rolf Neuser (Bochum) for technical support at the SEM and Alistair Ruffell

(Belfast) for comments on a previous version of the manuscript. This research was supported

by the Deutsche Forschungsgemeinschaft (Mu 667/23-1, -2).


65

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70

Chapter 4: Late Holocene environmental changes in the northern Red Sea…

71

4 Late Holocene environmental changes in the northern Red Sea region as indicated by

nannoplankton abundance patterns

Abstract

A sediment core retrieved from the Shaban Deep, a brine-filled, submarine basin located in

the northern Red Sea, was studied with respect to its calcareous nannoplankton content. We

focus on the abundance patterns of Gephyrocapsa spp. (Gephyrocapsa oceanica and

Gephyrocapsa ericsonii) and the reaction of this group to environmental changes in late

Holocene times. The changes of the abundance patterns observed for the last two millennia

correlate strongly with climatic fluctuations documented by proxy records from the

extratropics of the Northern Hemisphere. The calcareous nannoplankton record documents

two longer-lasting intervals with high Gephyrocapsa spp. values (A.D. 300–600 and A.D.

1350–1820), which coincide with the so-called “Migration Period Pessimum” and “Little Ice

Age” in Europe. Another short termed Gephyrocapsa spp. maximum has been observed

around A.D. 1050, punctuating the “Mediaeval Warm Period”. The Gephyrocapsa spp.

abundance patterns suggest that during cooler periods, documented particularly for Europe

but also for other areas of the Northern Hemisphere, the northern Red Sea was characterized

by a reinforced winter convective mixing. This in turn was probably caused by a

strengthened/more frequent cold air incursion from northerly direction. These findings are

compared with other proxy records of the late Holocene and discussed with respect to the

atmospheric background conditions potentially being responsible for the changes of the

northern Red Sea winter climate.

Key words: Northern Red Sea; Shaban Deep; nannoplankton; winter climate; late Holocene

4.1 Introduction

The Red Sea (Figure 4.1), located within the Afro-Arabian desert belt, is a long (around 1932

km) and narrow (averaged around 280 km) ocean basin, connected to the Gulf of Aden-Indian

Ocean by the shallow (137 m water depth) Strait of Bab-el-Mandeb at its southern end (e.g.,

Morcos, 1970; Edwards, 1987). Due to the negligible fresh water supply and the arid climate,

evaporation (2000 mm yr-1

) clearly outbalances precipitation (10– max. 200 mm yr-1

), surface

waters are characterized by a distinctive increase in salinity from 36–37‰ close to the Strait


72

of Bab-el-Mandeb to around 40–41‰ in

the northern Red Sea (Edwards, 1987;

Sofianos et al., 2002). The sea surface

temperatures (SST) reveal the opposite

trend; waters are getting continuously

cooler to the north. Following the

longitudinal axis of the basin, the

prevailing winds are basically from

northwest-to-southeast. Only south of

around 19°N the main direction reverses

in winter under the influence of the

Asian monsoon system (Edwards, 1987).

The anti-estuarine circulation patterns of

the Red Sea waters, the northwards

surface flow and the southwards sub-

surface return flow, are driven by wind

and especially thermohaline forcing with

an intermediate and deep-water formation in the northern Red Sea (Cember, 1988; Woelk and

Quadfasel, 1996; Eshel and Naik, 1997).

The northern Red Sea is situated in the northern margin of the subtropical desert belt, in

vicinity of the Mediterranean climate zone. Due to its sensitive hydrography, which is

influenced by extratropical meteorological elements (e.g., Eshel et al., 2000; Arz et al., 2003;

Felis et al., 2004; Lamy et al., 2006), it represents an interesting area to study paleoclimatic

changes. Fluctuations of the water stratification, driven by air temperature and humidity

changes, are influenced by variations of the North Atlantic winter climate and associated

pressure fields. In particular the North Atlantic Oscillation (NAO; Hurrell, 1995) and its

supra-regional counterpart, the Artic Oscillation (AO; Thompson and Wallace, 1998) play a

key role. Owning to a weak stratification of the northern Red Sea waters, relatively cold and

arid regional conditions are linked to positive NAO/AO index winters (e.g., Felis et al., 2000;

Eshel et al. 2000). A high-pressure anomaly, which extends over large parts of the

Mediterranean, favors an anticyclonic flow of surface winds in the eastern Mediterranean and

causes thereby the advection of northerly air from southeastern Europe towards the northern

Red Sea region (Rimbu et al., 2001; 2006; see also Wallace and Gutzler, 1981).

In an attempt to reveal changes of the northern Red Sea environment on a multidecal to

Figure 4.1. Map showing the Red Sea and the

Gulf of Aden. A detailed view of the northern

Red Sea with the Gulf of Aqaba, the Gulf of

Suez and the Shaban Deep is shown in the upper

right panel.


73

multicentennial time-scale for the last two millennia, we investigated variations of the

calcareous nannoplankton (further simply named nannoplankton). Sediment surface samples

as well as plankton studies have shown that the most common nannoplankton species in the

overall well-stratified waters of the northern Red Sea is Emiliania huxleyi (Winter, 1982).

Recent results have indicated that during colder periods in earth climate history different

hydrographic conditions prevailed; high abundances of Gephyrocapsa spp. are thought to

indicate an intensified convective mixing of the waters (Legge et al., 2006). The

Gephyrocapsa spp. abundance patterns can therefore probably be understood as an indicator

for the water column conditions, i.e. the strength and frequency of winter mixing in the

northern Red Sea.

This study aims to test whether Gephyrocapsa spp. (Gephyrocapsa ericsonii and G.

oceanica) is a useful indicator for variations of the environmental conditions in the northern

Red Sea region during the late Holocene. Further on we want to view and discuss potential

variations of Gephyrocapsa spp. abundances in the light of paleoceanographic changes caused

by tropical-extratropical teleconnection.

4.2 Methods

4.2.1 Material and geological background

The studied sediment core GeoB 7805-1 was retrieved from the Shaban Deep, an axial

depression of the northern Red Sea (Figure 4.1 and 4.2). It is one of 25 brine-filled, submarine

basins located in the central and northern

Red Sea (Karbe, 1987; Hartmann et al.,

1998). The Red Sea occupies a relatively

young tectonic basin. It is situated in the

long and narrow northwest-trending rift

zone where the African-Arabian continent

broke and spreads apart (Almond, 1986).

The main contour of the Red Sea

depression was probably achieved during

the Oligocene (Bohannon et al., 1989;

Daradich et al., 2003). In Miocene times,

the precipitation of salts resulted in the

buildup of evaporites on the bottom of the young basin. Today, leaching and diffusion of

these evaporites result in the formation of the brines (Manheim, 1974; Hartmann et al., 1998).

Figure 4.2. Simplified topographic map of the

Shaban Deep with its four sub-basins and side

of core GeoB 7805-1.


74

Two intersected ridges subdivide the Shaban Deep into four smaller sub-basins (Figure

4.2). High-resolution seismic profiles (e.g., Pätzold et al., 2003), in which the top of the

Miocene evaporites is marked by the so-called “S-Reflector”, indicate that the salt crops out

at the basins flanks. Due to its high density, the brine remains “pond-like” in the sub-basins

(brine body thickness around 200 m). The same brine-interface depth (1324.5 ± 0.5 m) and a

similar chemical composition indicate a sub-surface connection of all four sub-basins

(Hartmann et al., 1998). In the current study, we investigated core GeoB 7805-1, taken by a

multicorer during RV Meteor Cruise M52/3 (Pätzold et al., 2003) from the eastern sub-basin

(26°13´9"N and 35°22´6"E; water depth: 1447 m) of the Shaban Deep.

The extreme high salinity (chlorinities 150.7 ± 0.2‰ Cl) and the low oxygen content (<

0.3 mg O2/l) below the brine-seawater interface (Hartmann et al., 1998) have prevented

eukaryotic live and thereby bioturbation, preserving the fine lamination of the brine sediment.

The lack of benthic activity allows a high-resolution paleoceanographic reconstruction. A

detailed description of core GeoB 7805-1 is given by Seeberg-Elverfeldt et al. (2005). Most of

the 48 cm long core shows laminated sediments. From 7.6 cm to 9.5 cm (= HI1) and from

10.9 cm to 18.4 cm (= HI2) core depth two homogenous, non-laminated intervals are

intercalated. Due to the particle size distribution with coarser material at the bottom, both

intervals have been interpreted by Seeberg-Elverfeldt et al. (2005) to be of turbiditic origin.

4.2.2 Sampling and preparation

We tried to balance the sample density between resolution and totally covered time span.

Therefore, core GeoB 7805-1 was consecutively sampled every 0.5 cm. In order to gain

relative and absolute data of the nannoplankton, the filtration technique of Andruleit (1996)

was used for preparation. The freeze-dried sediments were weighted on a high-precision

balance, wet separated with a rotary splitter (FRITSCH Laborette 27) and filtered by means of

a vacuum pump on polycarbonate filters with 0.4 µm pore size. After drying at 40°C, a

wedge-shaped piece of the filter was mounted on an aluminum stub with self-adhesive tape,

outlined with a silver colloidal suspension and coated with gold in argon atmosphere using a

sputter coater (BioRad SC-500). At least 300 specimens per sample were identified and

counted by use of a scanning electronic microscope (LEO 1530 Gemini FESEM).

We investigated the complete core (also the homogeneous intervals/turbidties HI1 and

HI2). The nannoplankton results are prepared in two forms. (1) Primary data set: The absolute

and relative abundance patterns of Gephyrocapsa spp. and the Gephyrocapsa spp./E. huxleyi

ratios are shown in Figure 4.3 on the depth scale (all data are figured, intervals HI1 and HI2


75

are included). (2) Corrected for turbities: The relative abundance patterns of Gephyrocapsa

spp. and the Gephyrocapsa spp./Emiliania huxleyi ratio are shown in Figure 4.4 and 4.5 on

the age scale (intervals HI1 and HI2 are excluded; details will be given below).

4.3 Gephyrocapsa spp. abundance patterns

The abundance patterns of Gephyrocapsa spp. are shown in Figure 4.3 on the depth scale.

Two extensive intervals (around 40–45 cm and around 3–20 cm) of higher Gephyrocapsa spp.

values are documented. Throughout the reminder of the core, the abundances of

Gephyrocapsa spp. are, despite a minor peak at around 26 cm, characterized by overall low

values. Figure 4.4 and 4.5 shows the nannoplankton data on the age scale. The depth-age

model (corrected for turbidites) is taken from Seeberg-Elverfeldt et al. (submitted). Core

GeoB 7805-1 covers the time span from A.D. 160 to 1820 (1790 to 130 B.P.). The turbidite

Figure 4.3. Counted (a) absolute and (b, c) relative abundances of Gephyrocapsa spp. as a function

of the core depth. The relative abundances of Gephyrocapsa spp. are shown as their % proportion

on the total nannoplankton community (b; Gephyrocapsa spp. – TOTAL) and as their % propotion

in relation to Emiliania huxleyi (c; Gephyrocapsa. spp – Emiliania huxleyi). The homogeneous

intervals/turbidities HI1 (7,6–9,5 cm) and HI2 (10,9–18,4 cm) are outlines by gray shaded bars.


76

HI1 removed no previously deposited sediments, whereas turbidite HI2 eroded sediments

presenting a time interval of around one century. Based on this depth-age model, the two

long-termed Gephyrocapsa spp. maxima occur between A.D. 300 and 600 (1650 to 1350

B.P.) and between A.D. 1350 and 1820 (600 to 130 B.P.), respectively. The short-termed

maximum at 26 cm is dated A.D. 1050 (900 B.P.).

Figure 4.4. Comparison of (a) changes in alpine glacier extent (Holzhauser et al., 2005), (b)

variations of petrologic tracers of ice-rafted debris in the eastern North Atlantic (Bond et al., 2001),

(c) reconstructed NH temperature anomalies (Mann and Jones, 2003), (d) stalagmite growth rates

from the Uamh-an-Tartair cave in northwest Scotland (Proctor et al., 2000; 2002) and (e)

reconstructed solar irradiance changes (Bard et al., 2000; with different scaling factors: 0.25%,

0.4%, 0.55% and 0.65%) with the relative abundances of Gephyrocapsa spp. (f and g). Age control

points for core GeoB 7805-1 are indicated by triangles at the bottom.


77

4.4 Discussion

A general observation is that the core intervals, which are characterized by high values of

Gephyrocapsa spp. coincide in time with periods of overall colder conditions, documented in

most parts of Europe but also elsewhere in the Northern Hemisphere (NH). Changes in alpine

glacier extent (Holzhauser et al., 2005) and variations of petrologic tracers of ice-rafted debris

in the eastern North Atlantic (Bond et al., 2001) as well as reconstructed NH temperature

anomalies (Mann and Jones, 2003), documented overall similar trends (Figure 4.4). Periods of

glacier advance (Figure 4.4a), increased drift ice (Figure 4.4b) and lower NH temperatures

(Figure 4.4c) co-occur with higher values of Gephyrocapsa spp. in the northern Red Sea

(Figure 4.4f and g). The most prominent features in core GeoB 7805-1 are the two long-

lasting Gephyrocapsa spp. maxima (A.D. 300–600 and A.D. 1350–1820), which seem to

correlate with the “Migration Period Pessimum” (roughly between A.D. 300 and 700;

sometimes also “Dark Ages Cold Period”, e.g., Keigwin and Pickart, 1999) and the “Little Ice

Age” (LIA, roughly between A.D. 1350 and 1850). In addition, the Gephyrocapsa spp.

maximum at A.D. 1050, although less well reflected in the above records, probably matches a

short-lasting cool interval within the “Mediaeval Warm Period” (MWP) and the Oort solar

minimum (cf. Pfister et al., 1998; Damon and Peristykh, 2000; Pla and Catalan, 2005;

Ammann et al., 2007).

In order to integrate the observed nannoplankton changes into a more regional context,

we compare our data with records from the eastern Mediterranean (Figure 4.5). Based on the

interpretation of proxy records and archaeological evidence, Issar (1998) created an integrated

climate scheme for the eastern Mediterranean, which shows alternating cold-humid (Roman

Period = 2200–1700 B.P., Byzantine Period = 1800–1500 B.P., Crusader Period = 1000–800

BP, LIA = 500–100 B.P.) and warm-dry (Moslem-Arab Period = 1300–1000 BP, Moslem-

Turkish Period = 800–500 B.P.) phases for the late Holocene (Figure 4.5a). The long-lasting

Gephyrocapsa spp. maxima (Figure 4.5f and g) probably coincide with the Byzantine Period

and the LIA of Issar (1998). The short-lasting Gephyrocapsa spp. maximum at A.D. 1050

may fall within the Crusader Period. A detailed correlation of single proxy records from the

eastern Mediterranean with the nannoplankton abundance patterns of core GeoB 7805-2 is

more difficult. The Dead Sea lake level record for the last two millennia is shown in Figure

4.5b (Enzel et al., 2003; Bookman et al., 2004). Lake level high stands occurred at the end of

the Byzantine and Crusader period as well as from the end of the 19th century to the onset of

the 20th century. With regard to the proposed precipitation patterns, a deviation from the

scheme of Issar (1998) is provided by terrestrial and marine !18

O and !13

C records (!13

C not


78

further discussed here) from the southeastern Mediterranean (Figure 4.5c and d). Oxygen

isotope variations of the foraminifera Globigerinoides ruber) from the southeastern

Mediterranean Sea (Figure 4.5c; Schilman et al., 2001) and of the Israeli Soreq cave

speleothem (Figure 4.5d; Bar-Matthews et al., 1997; Bar-Matthews and Ayalon, 2004) are

linked to changes of temperature and especially to variations of the regional hydrological

balance. The information provided by both, the foraminiferal and the speleothem record, are

summarized by Schilman et al. (2002), which propose two dry (and/or cold) “events” (Event I

Figure 4.5. Comparison of paleoclimatic records from the eastern Mediterranean region with the

nannoplankton data: (a) simplified illustration of the cold-humid/warm-dry series of Issar (1998),

(b) Dead Sea lake level (Enzel et al., 2003), (c) foraminiferal (Schilman et al., 2002) and (d)

speleothem !18O (Bar-Matthews et al., 1997; Bar-Matthews and Ayalon, 2004) data from the south

eastern Mediterranean and (e) a schematic view of the precipitation/evaporation changes in the Nar

Gölü crater-lake in central Turkey (Jones et al., 2006). The displayed nannoplankton record (f and g) is identical with Figure 4.4f and 4.4g. Age control points are indicated as in Figure 4.4.


79

= 300 B.P., Event III = 900 B.P.) and two humid (and/or warm) “events” (Event II = 700

B.P., Event IV = 1300 B.P.) during the last two millennia.

Evidence, at least partly, for the basic NH climate trends can be also seen in the eastern

Mediterranean. However, even if we take dating uncertainties as well as different approaches

to interpret the impact of temperature and rainfall on the paleoclimatic archives for granted,

the overall impression from the above records is that they reflect also the complexity of

Mediterranean climate caused by orography, land-sea contrasts and the diversity of climate

regimes affecting the region (e.g., Kutiel et al., 1996; Kutiel and Benaroch, 2002; Alpert et

al., 2005; Ziv et al., 2006).

Our approach to explain the nannoplankton variations in core GeoB 7805-1 is based on

the intimate relationship between the physical state of the water column and the structure of

the phytocoenosis. Vertical mixing of the water column controls the nutrient distribution

within the photic zone and triggers the composition and annual succession of the

phytoplankton (e.g., Lindell and Post, 1995). A preference for higher levels of inorganic

macronutrients and an adaptation to relatively unstable environmental conditions is typical of

Gephyrocapsa spp. Gephyrocapsa oceanica and G. ericsonii are abundant in upwelling areas

and nutrient-enriched coastal regions, as well as in overall oligotrophic waters in response to

seasonal mixing events (e.g., Winter et al., 1979; Okada and Wells; 1997, Ziveri et al., 2004).

Since no nutrient-rich fresh waters enter the Red Sea (Morcos, 1970; Hoelzmann et al., 1998;

Siddall et al., 2003) and the relatively eutrophic waters from the Gulf of Aden-Indian Ocean

effect only the south (Weikert, 1987), higher trophic levels in the open northern Red Sea

photic zone mainly result from the upwards convective mixing of the nutrient-richer

subsurface waters in winter. Convection can penetrate to sizeable depths only if an adequate

increase in density results from cooling and evaporation of the surface waters. Reinforced

winter mixing, most likely induced by a strengthened/more frequent penetration of cold

northerly air into the northern Red Sea region during periods of higher Gephyrocapsa spp.

values, is therefore assumed.

Detailed insights into the succession patterns of various phytoplankton groups in the

northern Red Sea region are especially available from the Gulf of Aqaba (e.g., Levanon-

Spanier et al., 1979; Winter et al., 1979; Lindell and Post, 1995). The environmental

conditions in the Gulf are characterized by an annual cycle of water stratification and mixing,

accompanied by distinctive succession patterns of the phytoplankton community. Surface

waters are characterized during most of the year by a shallow, but stable thermocline. Air

temperature lowering in late fall causes a rapid erosion of the stratification and induces a deep


80

convective mixing up to 600 m water depth in winter (December, January, February).

Levanon-Spanier et al. (1979) investigated the larger sized phytoplankton (! 65 "m). The

community is dominated by dinoflagellates followed by diatoms and cyanobacteria

(Trichodesmium) in summer. In contrast, diatoms nearly solely dominate (within the spectrum

of investigated species) the winter period. Winter et al. (1979) studied the calcareous

nannoplankton. Emiliania huxleyi is the major component of the assemblage composition

during most time of the year. In winter, G. ericsonii, along with G. oceanica, increases in

abundance and characterizes the nannoplankton community in addition to E. huxleyi.

Furthermore, G. ericsonii is able to displace E. huxleyi as the dominating species during this

season. Lindell and Post (1995) compared the abundance patterns of three small sized

phytoplankton-groups: small eukaryotic algae (2–8 "m), the cyanobacteria Synechococcus (1–

2 "m) and the prochlorophyte Prochlorococcus (0.6–1 "m). The small eukaryotic algae

dominate in late winter, stimulated by the deep convective mixing and the associated injection

of nutrients into the photic zone. With the onset of stratification, eukaryotic algae decline and

Synechococcus increases in abundance. Prochlorococcus, absent in winter and rare during the

first two months after the onset of stratification, dominates the stratified and nutrient poor

waters in summer. In sum, the change of the surface water conditions, caused by the winter

convection, seems to be the most important factor affecting the composition of the

phytoplankton in the Gulf of Aqaba (Lindell and Post, 1995; Post, 2005). Previous studies

portray the influence of low temperature and evaporation as a key factor controlling the water

column stability in the northern Red Sea (see 4.1). Though limited in strength and depth, in

relation to the Gulf of Aqaba, the mixing of the waters during the short winter season triggers

the northern Red Sea environment and its biological interior (Wyrtki, 1974; Weikert, 1987;

Eshel et al., 2000).

As stated in the Introduction section, severe conditions in the northern Red Sea region are

thought to be favored today during positive NAO/AO index winters. Hence, the

Gephyrocapsa spp. maxima may be explained in term of prolonged positive NAO/AO index

winters. Taking the modern phase opposition between the northern Red Sea region and the

European sector as a response to NAO/AO for granted, such approach would probably be in

conflict with paleoclimatic information provided from other study sides (e.g., Proctor et al.,

2000; Shindell et al., 2001; Noren et al., 2002). The NAO index is based on the normalized

winter sea level atmospheric pressure differences between the Azores high and the Icelandic

low. Atmospheric circulation patterns, which cause stormy but relatively warm conditions in

northwestern Europe are favored today during positive NAO index winters (strong westerlies


81

over the eastern North Atlantic), whereas a negative index (weak westerlies) coincide with

cold and dry conditions (Hurrell, 1995). Enhanced European winter cooling, driven by strong

advection of northeasterly air under the influence of continental anticyclones during the LIA,

or at least during their severest stages (Maunder Minimum, Late Maunder Minimum), is

probably rather consistent with a negative than with a positive NAO (e.g., Luterbacher et al.,

2001; Shindell et al., 2001). Growth rate variations of stalagmites from the Uamh-an-Tartair

cave in northwest Scotland are interpreted to reflect variations of the local precipitation,

strongly controlled by the winter NAO (Figure 4.4d). High precipitation throughout the MWP

is linked to persistently positive index winters, whereas predominantly negative indexes mark

the LIA (Proctor et al., 2000; 2002). It has to be noted, that not all records agree on a

persistently negative state of the NAO during the LIA (Luterbacher et al., 2001; Cook et al.,

2002; Brückner and Mackensen; 2006). In fact, multiproxy reconstructions illustrate large

variability. For example, comparable periods of high positive NAO index winters, as

documented during the twentieth century, seem to have occurred from the mid 15th century

towards the mid 16th century (Cook et al., 2002). Irrespective by whether the NAO/AO

phenomenon exhibited a long lasting preferred mode or was restricted to an interannual and

interdecal timescale, the Gephyrocapsa spp. abundance patterns documented in core GeoB

7805-1 are difficult to ascribe to NAO/AO on the basis of its modern atmospheric flow

characteristics (cf. Keigwin and Pickart, 1999; Bond et al., 2001).

Other teleconnections than NAO/AO may have played a role in our study domain.

Prominent but less well-studied extratropical teleconnections detected in the eastern

Mediterranean region are the zonal oriented East Atlantic-West Russia Pattern (Barnston and

Livezey, 1987; Krichak and Alpert, 2005) and the comparable North Sea-Caspian Pattern

(NCP; Kutiel and Benaroch, 2002). A possible relationship between the NCP and the eastern

Mediterranean winter climate was recently suggested on the basis of a proxy record (for

precipitation) from the Nar Gölü crater-lake (Cappadocia region of Turkey; Jones et al.,

2006). Associated with an increased northeasterly atmospheric circulation across the

Balkan/Black Sea-region (Kutiel et al., 2002), below-normal temperature and precipitation is

documented in the Cappadocia region today during positive NCP episodes. In the Nar Gölü

crater-lake record (schematically illustrated in Figure 4.5e) overall dryer conditions are

documented from A.D. 300 to 500, around A.D. 800–900 and from A.D. 1400 to 1950.

Wetter conditions, on the other hand, prevailed from A.D. 560 to 750 and from A.D. 1000 to

1350. Though the NCP may play a significant role in the northeastern Mediterranean, the

conclusion that it also influenced the northern Red Sea winter environment is highly


82

speculative at this moment. More detailed records are essential to decipher the regional

importance of the NCP/East Atlantic-West Russia Pattern during the late Holocene.

Another approach to explain the Gephyrocapsa spp. maxima is that the linkage between

the northern Red Sea and the mid-latitudes changed more generally during the NH cold

periods. The reaction of the nannoplankton in the subtropical northern Red Sea to

extratropical forcing should be finally critical to the poleward development of the Hadley cell

and hence to strength and position of the associated upper-tropospheric subtropical jet

(located today at around 30°N in winter; Held and Hou, 1980; Chang, 1995). Due to the

exposed position of our study domain, even a subtle equatorward shift/weakening of the

Hadley cell/subtropical jet may increase the likelihood of extratropical air to affect the

northern Red Sea setting. Similar dynamics of the subtropical jet are observed during the last

decades in the opposite direction caused by the global warming (Fu et al., 2006) and may

account, at least partly, for the strong correlation between winter rainfall in Israel and ENSO

(El Niño-Southern Oscillation) events since the mid-1970s (Price et al., 1998; Alpert et al.,

2006). The signature of ENSO is also documented in a 245-year !18

O coral record from the

northern Red Sea (Felis et al., 2000). The authors consider variations in the strength of

NAO/ENSO atmospheric teleconnection (i.e., changing regional importance of extratropical

and tropical teleconnection modes) to explain the relationship between northern Red Sea

SSTs and coral !18

O variations on a multi-decadal time scale.

There is reason to assume that the influence from the extratropics could have been

intensified during the NH cold periods beyond those changes associated with phenomena like

the NAO/AO or other extratropical teleconnections. Following this line, the Gephyrocapsa

spp. abundance variations documented in core GeoB 7805-1 seem to reflect first and foremost

changes of the tropical-extratropical connection which may have been caused by a more or

less latitudinal-parallel shift/modification of the polewards branch of the Hadley cell. In

addition, a modulation of the regional characteristic of the Hadley cell/subtropical jet stream

patterns coupled with variations of the tropical circulation should be considered. The intensity

of tropical diabatic heating and associated tropical atmospheric circulation strength is seen to

modulate the location and shape of the ridges within the subtropical jet (Krishnamurti, 1979;

Yang and Webster, 1990). A weaker diabatic heating in central equatorial Africa may have a

lasting influence on the Mediterranean ridge of the jet. A recently published proxy record (Mg

% in authigenic calcite) from Lake Edwards provides evidence for droughts/reduced tropical

convection at A.D. 540–890, 1000–1200 and 1400–1750 in equatorial Africa (Russel and

Johnson, 2007).


83

Finally there is also the possibility that the late Holocene climatic changes are related to

variations of solar output. With regard to the mechanism treated above (i.e., Hadley

cell/Subtropical jet stream pattern), the effect of irradiance changes on the thermally structure

of the stratosphere and on the tropospheric winter climate is of particular interest. Model

studies (e.g., Haigh, 1996; Shindell et al., 1999) of the eleven-year sunspot cycle indicate a

broadening (retreating) of the Hadley circulation and poleward (equatorward) shifts of the

subtropical jets during periods of solar maximum (solar minimum). To the extent that it is

appropriate to apply the above model results to the late Holocene epoch, it seems possible that

solar variability played a crucial role on the observed nannoplankton changes by modifying

the poleward branch of the Hadley cell and thereby the tropical-extratropical connection.

Figure 4.4e shows the solar irradiance changes for the last 1200 years as indicated by

fluctuations of the cosmogenic nuclides 10-Beryllium (10

Be) and 14-Carbon (14

C) (Bard et al.,

2000). In comparison with the nannoplankton record (Figure 4.4f and g), similarities between

the Gephyrocapsa spp. maxima and intervals of reduced solar output are indicated.

It is worth mentioning that the relevance of solar variations in forcing late Holocene

climate is still the subject of much discussion (Lean and Rind, 1998; Bond et al., 2001;

Ammann et al., 2007) and that we cannot present here a final key for the understanding of the

sun-climate interaction. Whether solar activity was the dominating external forcing factor or

not, variations of the subtropical jet/of the Hadley cell are probably an important approach

that could help to explain the distinctive nannoplankton changes documented during the last

two millennia in the northern Red Sea and the coincidence of Gephyrocapsa spp. maxima and

lower NH temperatures.

4.5 Conclusions

The calcareous nannoplankton from the Shaban Deep supplies crucial information for a better

understanding of the Red Sea paleoenvironment. Throughout the last two millennia, the

northern Red Sea region was affected by climatic fluctuations, which left a clear imprint in

the nannoplankton record. The record revealed three intervals of higher Gephyrocapsa spp.

values, which seem to coincide in time with periods of overall lower temperatures.

Convective mixing, prefaced by the buoyancy loss of the surface waters, probably triggered

the increase of Gephyrocapsa spp. The changes of the regional hydrography can be explained

by a reinforced wintertime penetration of cold northerly air. Our observations underscore that

the linkage between the northern Red Sea and the mid-latitudes is complicated and varies

through time. The indicated changes of the northern Red Sea hydrographic are probably not


84

adequately explained by NAO/AO (supposing modern patterns). Suggestions are made,

linking the documented calcareous nannoplankton changes with other phenomena. Overall, it

remains important to pay attention to the exposed position of the northern Red Sea near the

tropical-extratropical atmospheric “boundary”. Thermally driven changes of the Hadley cell

in winter and of the associated subtropical jet may have contributed to the observed

nannoplankton changes by modulating the tropical-extratropical connection. Variations of

solar output are perhaps the primary trigger; further research is, however, needed to

understand its influence on the climate, especially with regard to the conditions in the North

Atlantic realm and associated atmospheric circulation patterns.

Acknowledgements

This study was supported by the Deutsche Forschungsgemeinschaft (Mu 667/26-1, -2).

Sample material has been supplied by Ismene A. Seeberg-Elverfeldt and Jürgen Pätzold

(Bremen). Special thanks go to Rolf Neuser for technical support at the SEM and to Sylvia

Rückheim for helpful comments and discussions.


85

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91





92

Chapter 5: Summary

93

5 Summary

Following the outline of the thesis, the basic findings are summarized in the order of the three

manuscripts. Finally, some general remarks concerning the calcareous nannoplankton

abundance patterns and the responsible background mechanisms are given.

1.) Covering the last 22,000 years, the calcareous nannoplankton record of core GeoB 5844-2

allows the identification of the Last Glacial Maximum, the Heinrich Event 1, the

Bølling/Allerød warm period, the Younger Dryas and the Holocene in the northern Red Sea.

Furthermore, the subdivision of the Holocene and the identification of two distinctive short-

term phases subsequent to the Bølling/Allerød warm period and the “Holocene Humid

period” are possible.

Extreme conditions during the hypersaline “aplanktic zone” are characterized by

Gephyrocapsa ericsonii and Emiliania huxleyi. High abundances of Gephyrocapsa ericsonii

during the Last Glacial Maximum point to moderately cool and fertile conditions, whereas the

dominance of Emiliania huxleyi during Heinrich Event 1 goes ahead with a further climatic

cooling favoring the bloom of opportunistic species. Punctuated by the Younger Dryas with

high values of Gephyrocapsa oceanica, a two-step onset of the postglacial humid climate

characterized the calcareous nannoplankton community. Both steps show an early

oligotrophic phase dominated by Florisphaera profunda and Gladiolithus flabellatus, and a

subsequent fertile phase characterized by Emiliania huxleyi. Following the Emiliania huxleyi-

dominated “Holocene Humid period”, the repetitive increases in abundance of F. profunda

and G. flabellatus reflect modern oligotrophic conditions accompanied by high aridity on

land.

2.) The glacial northern Red Sea represents a unique marine environment with regard to the

hypersaline “aplanktic zone”. The findings summarized above, however, indicate the presence

of calcareous nannoplankton during this period. Northern Red Sea environmental changes

were analyzed in the second study with special reference to the longer-term cooling cycles

and Heinrich events of the North Atlantic realm. The calcaroues nannoplankton record

obtained from core GeoB 5844-2 allows a precise and detailed correlation with the North

Atlantic climate trends. It is shown that the calcareous nannoplankton responded differently to

high latitudinal cooling before and after ca. 32 ka. A reiterated succession, closely correlated

with the longer-term cooling cycles and Heinrich events, is detected after 32 ka. The three

Heinrich events (H3, H2, and H1) are dominated by Emiliania huxleyi. Gephyrocapsa

oceanica and especially Gephyrocapsa ericsonii are abundant between H3 and H2 and

Chapter 5: Summary

94

between H2 and H1. Fluctuations in the abundance of Emiliania huxleyi, Gephyrocapsa

oceanica and Gephyrocapsa ericsonii are also documented before 32 ka. However, no

analogue succession as in the younger interval was found. The record presented here, and the

comparison with other microfossil data, led to the assumption that the glacial salinity increase

did not cause the documented calcareous nannoplankton succession. The calcareous

nannoplankton record points to a sensitive and fast acting mechanism linking high latitudinal

climate oscillations and environmental changes in the northern Red Sea region, especially

after 32 ka. The key mechanism seems to be the atmospheric teleconnection with the

extratropics and the associated hydrographic changes driven by a reinforced influence of cold

air. A strong influence from the mid-latitudes is evident since the late Marine Isotope Stage 3.

This approach to explain the calcareous nannoplankton abundance patterns is in good

agreement with climatic evolution of the northern subtropical desert belt.

3.) The third study focuses on the abundance patterns of Gephyrocapsa spp. (Gephyrocapsa

oceanica and Gephyrocapsa ericsonii). The response of Gephyrocapsa spp. to environmental

changes over the course of the last two millennia was analyzed. The documentation of

calcareous nannoplankton changes in late Holocene times corroborates the findings from core

GeoB 5844-2. Overall, the record from the Shaban Deep provides further evidence for the

linkage between the northern Red Sea and the mid-latitudes. The calcareous nannoplankton

record documents three intervals with high Gephyrocapsa spp. values in the northern Red

Sea. The Gephyrocapsa spp. abundance peaks seem to coincide with cold periods in Europe.

Reinforced convective mixing has probably triggered their increase. The changes of the

regional hydrography can be explained by a strengthened penetration of cold air from

northerly direction. The Gephyrocapsa spp. seem to be a sensitive indicator for the

hydrography of the northern Red Sea winter environment. Similar conditions, as documented

during the Gephyrocapsa spp.-maxima, are thought to be associated today with positive index

winters of the North Atlantic Oscillation/Artic Oscillation (NAO/AO). However, the

Gephyrocapsa spp. maxima in late Holocene times are probably inadequately explained on

the basis of the modern NAO/AO patterns. It is tentatively assumed here that the

Gephyrocapsa spp. record may reflect fluctuations in the strength and/or position of the

subtropical jet/development of the poleward branch of the Hadley cell.

In summary, this thesis represents the first well-dated, comprehensive calcareous

nannoplankton study from the northern Red Sea region. As remarkable as the glacial salinity

increase seems, the changes documented in core GeoB 5844-2 indicate that the calcareous

nannoplankton were less affected by variations of the salinity than other fossil groups. Other

Chapter 5: Summary

95

factors explain the observed distribution patterns probably better. Environmental factors

affecting phytoplankton distribution can interact in complex ways. The data from core GeoB

5844-2 point to variations of the water stratification and its influence on the nutrient

distribution within the photic zone as a crucial factor for the structure of calcareous

nannoplankton community. Closely coupled to the climate of the northern mid and high

latitudes, basic changes documented in the fossil assemblage compositions seem to be driven

by local hydrographic changes under the atmospheric influence from the extratropics. These

findings are consistent with modern observations, which illustrate the sensitivity of the

northern Red Sea hydrography to mid-latitude continental climate. The investigation of core

GeoB 7805-1 may reveal more about the atmospheric elements involved. One of the most

distinctive features of the northern Red Sea is its close position to the tropical-extratropical

atmospheric “boundary”. The results from core GeoB 7805-1 seem to indicate a crucial role

of the position and/or strength of the subtropical jet/the development of the poleward branch

of the Hadley cell on environmental changes in late Holocene times.

96

Chapter 6: Outlook

97

6 Outlook

The Near East-North African region represents a key area of human civilization where

important changes took place in prehistoric and historic time. Controlling the vegetation cover

and the water storage, temperature and especially precipitation changes in this dry part of the

world were key factors for human settlement and cultural development. Furthermore, the

availability of fresh water (Figure 6.1) is

seen as a crucial factor for the regional

development and the prevention of social

and military crises in the near future. Since

observational data cover only a restricted

time span, the study of paleoclimatic

archives is crucial for the understanding of

the climate system. Continuous and well-

dated records are needed for this attempt.

Unfortunately the arid environment with its

unprotected land surface limited the

availability of such records from the

terrestrial setting. One answer to this

problem is the use of records from nearby

marine sites. The study of microfossils and

certain geochemical-sedimentological data

can provide insights into the past oceanographic setting and hence into the climate history.

The calcareous nannoplankton community, and the study of its composition, represents one

fruitful attempt. It has been unequivocally shown here that this fossil group mirrors the

climatic trends of the last 60,000 years for the northern Red Sea region. The calcareous

nannoplankton data extended and improved past studies undertaken in this region: Fortunately

the calcareous nannoplankton seems to be less affected by the glacial salinity increase than

other marine organisms; even during hypersaline periods a detailed reconstruction of the

paleoceanographic conditions is possible. In order to update the analysis of past

environmental changes, more detailed plankton studies are needed. Better information on the

seasonal succession patterns of calcareous nannoplankton species, together with detailed

physical and chemical measurements, are probably an important requirement to refine the

explanatory power of the fossil records. Paleoceanographic studies of Red Sea sediments

older than Marine Isotope Stage 3 may help to gain deeper insights into the climatic evolution

Figure 6.1. Map showing the Red Sea area

and the water availability in 2000 (based on

information provided by the United Nations

Environment Programme).

Chapter 6: Outlook

98

of the African-Arabia desert belt. In addition, the accompanying investigation of marine

records from the Mediterranean Sea could be an important approach to get more detailed

information about the extratropical influence on northern Red Sea environmental changes and

the role of the North Atlantic Oscillation/Artic Oscillation (NAO/AO). The hypothesized role

of the subtropical jet in modulating marine environmental conditions in the northern Red Sea

region raises further questions with respect to its development and specific regional impact.

There is no final answer to the question of how variations of the sun are included. Direct

observations of how solar variations modulate the atmospheric structure are limited to short-

term effects. Furthermore, our knowledge about the specific role of the El Niño-Southern

Oscillation (ENSO) and the NAO/AO in influencing jet stream patterns on a longer time scale

is far from being complete.

99

Appendix

100

Methods

101

All nannoplankton samples were investigated using a

scanning electron microscope (SEM; LEO 1530 Gemini

FESEM). The filtration technique of Andruleit (Andruleit,

H., 1996, A filtration technique for quantitative studies of

coccoliths, Micropaleontol., 42, 403–406) was applied for

preparation. The freeze-dried sediment was weighted on a

high-precision balance, wet-separated with a rotary splitter

(FRITSCH Laborette 27; picture on the right hand) and

filtered by means of a vacuum pump on polycarbonate

filters. After drying at 40°C (2-3 hours), a wedge-shaped

piece of the filter was mounted on an aluminum stub with

self-adhesive tape, outlined with a silver colloidal suspension and coated with gold in argon

atmosphere using a sputter coater (BioRad SC-500). At least 300 specimens per sample were

identified and counted.

The absolute abundance was calculated with the following equation:

AA = (F x C x S)/(A x W)

AA = absolute abundance (number g-1

)

F = total sediment coated filter area (mm2)

C = number of counted coccoliths/nannoliths

S = split factor

A = investigated filter area (mm2)

W = weight of the dry sample (g)

The accumulation rates were calculated with the following equation:

AR = AA x SR x DD

AR = accumulation rate (number cm-2

kyr-1

)

SR = sedimentation rate (cm kyr-1

)

DD = dry bulk density (g cm-3

)

Methods

102

In addition, in order to control SEM preparation and to study general changes in the sediment

composition, all sample were studied with an OLPYMPUS BH-2 light microscope using

polarizing light at a magnification of x1250 and dark field at a magnification of x150. For

slide preparation please see Bown et al. (Bown, P. R., and J. R. Young, 1998, Techniques, in

Calcareous Nannofossil Biostratigraphy, edited by P. R. Bown, pp 16–28, Chapman and Hall,

London).

Taxonomy

103

Species identification followed the description of Kleijne (Kleijne, A., 1993, Morphology,

taxonomy and distribution of extant coccolithophorids (calcareous nannoplankton), Ph.D.

thesis, 320 pp., Vrije Universiteit, Amsterdam), Jordan and Kleijne (Jordan, R. W., and A.

Kleijne, 1994, A classification system for living coccolithophores, in Coccolithophores,

edited by A. Winter and W. G. Siesser, pp. 83–105, Cambridge Univ. Press, Cambridge) and

Young et al. (Young, J. R., M. Geisen, L. Cros, A. Kleijne, I. Probert, C. Sprengel, and J. B.

Ostergaard, 2003, A guide to extant coccolithophore taxonomy, J. Nannopl. Res. Spec. Iss., 1,

124 pp.).

Below an overview (up to the genus level) of the classification schema of Young et al.

(2003) is given:

Classification system

Order ISOCHRYSIDALES

Family Isochrysidaceae (non-calcifying)

Isochrysis

Chrysotila

Family Noelaerhabdaceae

Emiliania

Gephyrocapsa

Reticulofenestra

Order COCCOSPHAERALES

Family Coccolithaceae

Coccolithus

Cruciplacolithus

Family Calcidiscaceae

Calcidiscus

Hayaster

Oolithotus

Umbilicosphaera

Family Pleurochrysidaceae

Pleurochrysis

Jomonlithus

Family Hymenomonadaceae

Hymenomonas

Ochrosphaera

Taxonomy

104

Order ZYGODISCALES

Family Helicosphaeraceae

Helicosphaera

Family Pontosphaeraceae

Pontosphaera

Scyphosphaera

Order SYRACOSPHAERALES

Family Syracosphaeraceae

Calciopappus

Michaelsarsia

Ophiaster

Coronosphaera

Syracosphaera

Family Calciosoleniaceae

Calciosolenia

Alveosphaera

Family Rhabdosphaeraceae

Acanthoica

Algirosphaera

Anacanthoica

Cyrtosphaera

Discosphaera

Palusphaera

Rhabdosphaera

Saturnulus

HETEROCOCCOLITHS INCERTAE SEDIS, Heterococcolith families and genera incertae sedis

Family Alisphaeraceae

Alisphaera

Canistrolithus

(Polycrater)

Family Umbellosphaeraceae

Umbellosphaera

Narrow-rimmed placoliths

“Calyptrosphaera”

Tetralithoides

Turrilithus

Placorhombus

Family Papposphaeraceae

Papposphaera

Pappomonas

Taxonomy

105

Narrow-rimmed muroliths, Genera incertae sedis aff. Papposphaeraceae

Picarola

Vexillarius

Wigwamma

NANNOLITHS

Family Braarudosphaeraceae

Braarudosphaera

Family Ceratolithaceae

Ceratolithus

(Neosphaera)

Nannoliths incertae sedis

Florisphaera

Gladiolithus

Ericiolus

HOLOCOCCOLITHS [Calyptrosphaeraceae]

Stratigraphic framework

106

Age model of core GeoB 5844-2 used in the first manuscript/Chapter 2 (Arz, H. W., J.

Pätzold, P. J. Müller, and M. O. Moammar, 2003, Influence of Northern Hemisphere climate

and global sea level rise on the restricted Red Sea marine environment during termination I,

Paleoceanography, 18(2), 1053, doi:10.1029/2002PA000864):

Core Depth 14

C AMS Age Calibrated Age

(cm) (years BP) (calendar years BP)

0 1015 ±35 516

10 2490 ±80 2002

20 3920 ±45 3764

30 5325 ±50 5584

40 6320 ±100 6658

55 8525 ±55 8922

70 9990 ±50 10630

85 12030 ±55 13430

100 13120 ±60 14497

125 14960 ±120 17205

150 17030 ±90 19587

175 18580 ±100 21371

200 20540 ±120 23626

225 22090 ±160 25433


107

Age model of GeoB 5844-2 used in the second manuscript/Chapter 3 (Arz, H. W., F., Lamy,

A. Ganopolski, N. Nowaczyk, and J. Pätzold, 2007, Dominant Northern Hemisphere climate

control over millennial-scale glacial sea-level variability, Quat. Sci. Rev., 26, 312–321):

Core Depth 14

C AMS Age Calibrated Age (calendar years BP)

(cm) (years BP) CalPal Fairbanks

0 1015 ±35 510 ±30 505 ±19

10 2490 ±80 1870 ±100 1866 ±90

20 3920 ±45 3590 ±70 3589 ±60

30 5325 ±50 5470 ±90 5506 ±76

40 6320 ±100 6560 ±110 6546 ±118

55 8525 ±55 8820 ±120 8814 ±131

70 9990 ±50 10650 ±60 10653 ±67

85 12030 ±55 13360 ±80 13369 ±110

100 13120 ±60 14840 ±150 14745 ±130

125 14960 ±120 17680 ±160 17438 ±215

150 17030 ±90 19670 ±140 19581 ±116

175 18580 ±100 21700 ±130 21478 ±144

200 20540 ±120 23870 ±170 23927 ±151

225 22090 ±160 25720 ±310 25921 ±239

250 25110 ±210 29580 ±370 29179 ±325

275 26360 ±240 30660 ±220 30952 ±202

300 28890 ±290 32890 ±710 33005 ±313

325 32060 ±400 36500 ±400 36468 ±407

Paleomagnetic excursions

361 Laschamp event 41800 calendar years BP

417 – 49400 calendar years BP

517 Norwegian-Greenland-Sea event 66100 calendar years BP


108

Age model for the core GeoB 7805-1 (Shaban Deep) used in the third manuscript/Chapter 4

(Seeberg-Elverfeldt, I. A., J. Pätzold, H. W. Arz, and J.-B. W. Stuut, 2007, Late Holocene

climate variability in the northern Red Sea; submitted to Paleoceanography):

Core Core Depth 14

C AMS Age Cal. Age (cal. years)

Orig./trans. (cm) (years BP) BP AD

GeoB 7805-1 0,5/0,5 *725 ±25 175 1775

10/8,2 935 ±25 385 1565

19/9,7 1115 ±25 550 1400

25/15,7 1535 ±30 880 1070

35/25,7 1825 ±30 1200 750

45/35,7 2245 ±30 1630 320

GeoB 7803-2_2 10,5/10,5 *705 ±30 150 1800

38/16,7 1295 ±30 785 1165

52,2/29,7 1865 ±30 1230 720

GeoB 7804-4_1 7,5/7,5 870 ±25 345 1605

38/18,2 1570 ±25 945 1005

*Combined since both data points lie within their respective error range

Note: Since cores from the Shaban Deep show distinct lamination that can be used for cross-

correlation, 14

C-ages were transferred from other cores (GeoB 7803-2, GeoB 7804-4) to core

GeoB 7805-1 to improve the age model.

Chapter 2

Figure 2.3

Core

dep

th (

cm)

Cal

endar

age

(B.P

.)

Abso

lute

abundan

ces

(no.g

-1 s

ed.)

Acc

um

ult

ion r

ates

(no. cm

-2 k

a-1)

Core

dep

th (

cm)

Cal

endar

age

(B.P

.)

Abso

lute

abundan

ces

(no.g

-1 s

ed.)

Acc

um

ult

ion r

ates

(no. cm

-2 k

a-1)

1 665 1,99E+10 1,54E+11 90 13786 1,67E+10 1,57E+11

2 813 2,47E+10 1,91E+11 92 13928 1,46E+10 1,22E+11

4 1110 2,38E+10 1,84E+11 94 14070 1,35E+10 1,20E+11

6 1408 2,30E+10 1,88E+11 96 14212 1,26E+10 1,12E+11

8 1705 2,85E+10 2,44E+11 98 14355 1,13E+10 1,74E+11

10 2002 2,62E+10 2,05E+11 100 14497 6,63E+09 1,09E+11

12 2354 3,34E+10 2,05E+11 102 14714 6,89E+09 7,61E+10

14 2707 2,84E+10 1,70E+11 106 15147 6,57E+09 6,39E+10

16 3059 1,90E+10 1,17E+11 110 15580 6,30E+09 4,34E+10

18 3412 1,58E+10 1,03E+11 114 16013 6,69E+09 4,30E+10

20 3764 2,26E+10 1,44E+11 118 16447 5,94E+09 3,36E+10

22 4128 1,82E+10 1,14E+11 122 16880 3,76E+09 1,62E+10

24 4492 2,22E+10 1,40E+11 126 17300 3,26E+09 1,85E+10

26 4856 2,89E+10 1,88E+11 130 17681 2,15E+09 1,24E+10

28 5220 2,07E+10 1,35E+11 134 18063 2,45E+09 1,47E+10

30 5584 2,13E+10 1,37E+11 138 18444 2,24E+09 1,43E+10

32 5799 2,19E+10 2,47E+11 142 18825 2,59E+09 1,61E+10

34 6014 2,02E+10 2,09E+11 145 19111 2,50E+09 1,51E+10

36 6228 1,49E+10 1,58E+11 150 19587 3,06E+09 1,87E+10

38 6443 1,64E+10 1,73E+11 154 19872 2,62E+09 2,00E+10

40 6658 2,57E+10 2,65E+11 158 20158 2,16E+09 1,70E+10

42 6960 2,63E+10 1,95E+11 162 20443 2,46E+09 1,97E+10

44 7262 3,01E+10 2,18E+11 166 20729 2,40E+09 1,92E+10

46 7413 2,09E+10 1,56E+11 170 21014 2,12E+09 1,92E+10

48 7865 2,71E+10 1,99E+11 174 21300 2,60E+09 2,11E+10

50 8167 2,46E+10 1,68E+11 178 21642 3,06E+09 1,86E+10

52 8469 3,35E+10 2,36E+11 182 22002 2,27E+09 1,47E+10

54 8771 3,48E+10 2,35E+11

56 9036 2,84E+10 2,58E+11

58 9264 3,51E+10 3,14E+11

60 9492 2,44E+10 2,22E+11

62 9720 2,36E+10 2,06E+11

64 9948 2,65E+10 2,45E+11

66 10176 2,25E+10 2,02E+11

68 10404 2,63E+10 2,25E+11

70 10632 2,23E+10 1,69E+11

72 11005 1,53E+10 7,79E+10

74 11378 1,40E+10 7,03E+10

76 11751 1,58E+10 7,86E+10

78 12124 1,46E+10 6,66E+10

80 12497 1,43E+10 6,67E+10

82 12870 1,85E+10 8,08E+10

84 13243 1,71E+10 7,21E+10

86 13501 1,80E+10 1,74E+11

88 13643 1,94E+10 1,79E+11

109

Chapter 2

Figure 2.4 and 2.5C

ore

dep

th (

cm)

Cal

endar

age

(B.P

.)

Um

bil

icosp

haer

e sp

p. (%

)

Hel

icosp

haer

a c

art

eri

(%)

Gep

hyr

oca

psa

eri

csonii

(%

)

Gep

hyr

oca

psa

oce

anic

a (

%)

Em

ilia

nia

huxl

eyi

(%)

Flo

risp

haer

a p

rofu

nda (

%)

Gla

dio

lith

us

flabel

latu

s (%

)

Core

dep

th (

cm)

Cal

endar

age

(B.P

.)

Um

bil

icosp

haer

e sp

p. (%

)

Hel

icosp

haer

a c

art

eri

(%)

Gep

hyr

oca

psa

eri

csonii

(%

)

Gep

hyr

oca

psa

oce

anic

a (

%)

Em

ilia

nia

huxl

eyi

(%)

Flo

risp

haer

a p

rofu

nda (

%)

Gla

dio

lith

us

flabel

latu

s (%

)

1 665 1,5 0,3 8,9 1,2 24,3 49,2 5,5 90 13786 0,0 0,7 7,5 2,4 47,3 17,0 12,9

2 813 1,1 0,3 5,7 4,2 25,2 40,8 11,6 92 13928 0,0 0,9 9,4 2,5 43,9 25,4 8,2

4 1110 2,9 0,3 5,1 1,6 19,9 52,1 8,7 94 14070 0,0 0,9 4,0 3,4 52,2 16,0 9,2

6 1408 2,5 0,6 8,4 1,2 23,0 50,6 5,3 96 14212 0,0 0,0 3,0 2,0 53,1 24,5 4,0

8 1705 1,0 0,0 9,9 3,0 19,1 46,7 9,2 98 14355 0,0 0,4 7,3 1,1 15,8 29,0 23,5

10 2002 1,5 0,0 8,2 2,9 23,1 40,9 15,5 100 14497 0,0 0,7 10,5 1,4 17,3 48,0 15,5

12 2354 0,6 0,0 8,9 4,6 19,6 41,3 18,3 102 14714 0,3 0,0 5,8 1,0 19,9 54,0 16,2

14 2707 2,2 0,0 8,8 4,4 21,3 40,0 16,9 106 15147 0,0 0,0 6,0 0,4 27,3 55,3 4,6

16 3059 1,3 0,0 6,9 3,3 24,3 39,0 18,4 110 15580 0,0 0,0 4,8 0,7 43,5 35,3 5,8

18 3412 1,0 0,3 6,6 2,8 27,3 39,9 14,3 114 16013 0,0 0,3 7,7 1,0 40,8 37,5 5,4

20 3764 1,9 0,0 7,0 5,7 23,5 48,3 4,8 118 16447 0,3 0,7 7,4 0,3 42,6 35,5 7,1

22 4128 1,2 0,3 7,8 3,4 27,0 38,8 11,2 122 16880 0,0 0,3 10,3 0,7 29,0 36,0 11,0

24 4492 2,8 0,0 6,5 1,8 25,5 40,0 10,5 126 17300 0,0 0,7 14,7 0,7 20,3 36,0 15,0

26 4856 1,2 0,3 7,2 3,2 21,4 44,6 10,4 130 17681 0,0 0,7 43,0 2,7 7,6 29,9 9,3

28 5220 1,7 0,0 6,6 3,6 29,1 40,7 8,6 134 18063 0,0 0,7 31,5 2,7 4,8 37,7 7,2

30 5584 1,9 0,3 5,2 1,6 28,5 48,3 6,2 138 18444 0,0 0,7 33,2 1,0 4,7 40,0 11,9

32 5799 1,9 0,0 6,9 1,2 30,6 41,5 8,7 142 18825 0,0 0,6 33,3 0,0 2,9 37,9 14,6

34 6014 1,9 0,0 5,1 1,9 29,4 44,3 11,1 145 19111 0,0 1,7 32,4 0,3 3,7 43,5 9,7

36 6228 1,9 0,6 4,5 3,2 35,2 42,6 6,5 150 19587 0,0 2,4 34,9 1,0 4,1 32,1 9,1

38 6443 1,6 0,3 4,5 1,9 33,9 40,3 5,8 154 19872 0,0 2,0 31,0 1,0 3,3 29,7 7,9

40 6658 1,2 0,0 9,2 4,0 21,0 37,9 13,6 158 20158 0,0 2,8 23,4 0,7 5,9 25,2 7,3

42 6960 1,5 0,0 10,2 1,5 26,5 29,3 14,5 162 20443 0,0 1,0 23,7 1,0 1,4 32,5 8,8

44 7262 0,4 0,4 7,0 3,2 25,7 28,5 17,2 166 20729 0,0 6,1 24,5 3,2 3,5 27,4 5,4

46 7413 1,6 0,3 7,5 3,3 32,1 34,0 12,8 170 21014 0,0 5,7 20,0 1,3 5,0 40,0 6,7

48 7865 2,1 0,7 8,3 3,1 28,1 37,2 9,0 174 21300 0,0 6,4 23,8 1,0 5,7 40,3 5,4

50 8167 0,7 0,3 6,2 1,3 35,8 37,8 10,4 178 21642 0,0 3,6 22,0 2,0 3,6 44,4 7,2

52 8469 1,3 0,3 6,9 4,6 33,0 33,3 10,4 182 22002 0,0 7,4 22,8 2,3 4,4 30,5 7,7

54 8771 0,3 0,0 7,6 2,9 32,6 33,9 8,9

56 9036 2,4 0,3 5,4 1,4 30,1 38,2 13,2

58 9264 0,0 0,0 8,3 3,3 24,3 37,3 16,3

60 9492 1,3 0,0 2,7 2,0 28,9 38,0 18,5

62 9720 1,7 0,0 4,8 3,1 25,9 39,7 19,0

64 9948 1,4 0,0 4,5 2,4 27,7 38,8 20,1

66 10176 0,0 0,3 3,7 1,7 26,9 41,9 16,6

68 10404 0,6 0,0 3,5 1,9 18,7 41,8 24,7

70 10632 0,7 0,0 3,7 1,7 26,5 38,9 22,1

72 11005 0,3 0,6 4,6 0,9 38,7 35,3 12,7

74 11378 0,4 0,0 3,9 5,0 40,2 30,2 12,9

76 11751 0,0 2,4 2,8 4,5 30,0 34,9 13,3

78 12124 0,0 0,0 4,3 8,9 34,2 35,2 10,9

80 12497 0,3 0,7 2,1 9,6 27,8 31,9 18,5

82 12870 0,0 0,4 4,2 7,1 37,0 20,5 9,9

84 13243 0,0 1,3 2,4 6,4 28,3 35,4 14,5

86 13501 0,0 0,7 6,9 2,3 43,2 30,4 5,2

88 13643 0,0 1,0 4,0 3,7 40,8 32,0 9,8

110

Chapter 3

Figure 3.3

Core

dep

th (

cm)

Cal

endar

age

(B.P

.)

Em

ilia

nia

huxl

eyi

(%)

Gep

hyr

oca

psa

eri

csonii

(%

)

Gep

hyr

oca

psa

oce

anic

a (

%)

Abso

lute

abundan

ces

(no.g

-1 s

ed.)

Acc

um

ula

tion r

ates

(no.c

m-2 k

a-1)

Core

dep

th (

cm)

Cal

endar

age

(B.P

.)

Em

ilia

nia

huxl

eyi

(%)

Gep

hyr

oca

psa

eri

csonii

(%

)

Gep

hyr

oca

psa

oce

anic

a (

%)

Abso

lute

abundan

ces

(no.g

-1 s

ed.)

Acc

um

ula

tion r

ates

(no.c

m-2 k

a-1)

96 14335 53,1 3,0 2,0 1,26E+10 8,11E+10 270 30384 44,3 13,8 0,9 9,45E+09 1,68E+11

98 14533 15,8 7,3 1,1 1,13E+10 1,26E+11 274 30557 53,9 12,7 1,0 7,94E+09 1,25E+11

100 14730 17,3 10,5 1,4 6,63E+09 7,86E+10 278 30889 61,9 8,8 1,0 9,99E+09 8,18E+10

102 14967 19,9 5,8 1,0 6,89E+09 7,26E+10 282 31275 75,2 4,1 2,5 5,69E+09 4,73E+10

106 15440 27,3 6,0 0,4 6,57E+09 6,09E+10 286 31660 52,1 7,8 7,8 5,44E+09 4,84E+10

110 15914 43,5 4,8 0,7 6,30E+09 4,14E+10 290 32046 49,3 6,6 15,6 6,01E+09 5,24E+10

114 16388 40,8 7,7 1,0 6,69E+09 4,10E+10 294 32432 20,5 19,1 21,1 5,49E+09 4,31E+10

118 16861 42,6 7,4 0,3 5,94E+09 3,20E+10 298 32817 40,4 9,9 6,0 6,21E+09 5,00E+10

122 17335 29,0 10,3 0,7 3,76E+09 1,55E+10 302 33299 37,5 12,2 2,2 7,55E+09 4,26E+10

126 17776 20,3 14,7 0,7 3,26E+09 2,21E+10 306 33876 16,9 4,8 16,9 3,96E+08 2,27E+09

130 18118 7,6 43,0 2,7 2,15E+09 1,48E+10 310 34454 28,4 8,4 6,9 7,60E+09 4,49E+10

134 18460 4,8 31,5 2,7 2,45E+09 1,76E+10 314 35032 35,7 15,4 7,2 7,68E+09 4,50E+10

138 18803 4,7 33,2 1,0 2,24E+09 1,71E+10 318 35609 43,4 9,9 4,3 7,84E+09 4,59E+10

142 19145 2,9 33,3 0,0 2,59E+09 1,92E+10 322 36187 40,5 7,2 8,9 7,37E+09 4,44E+10

145 19402 3,7 32,4 0,3 2,50E+09 1,81E+10 326 36764 48,9 5,9 12,1 6,12E+09 3,95E+10

150 19830 4,1 34,9 1,0 3,06E+09 2,20E+10 330 37339 36,1 6,5 20,6 6,97E+09 4,50E+10

154 20147 3,3 31,0 1,0 2,62E+09 1,76E+10 334 37915 30,7 14,1 15,7 8,64E+09 5,34E+10

158 20464 5,9 23,4 0,7 2,16E+09 1,49E+10 338 38491 37,7 8,9 10,1 1,03E+10 6,23E+10

162 20780 1,4 23,7 1,0 2,46E+09 1,73E+10 342 39066 32,8 10,3 3,2 8,09E+09 4,99E+10

166 21097 3,5 24,5 3,2 2,40E+09 1,69E+10 346 39642 28,3 10,4 8,1 1,00E+10 6,29E+10

170 21414 5,0 20,0 1,3 2,12E+09 1,69E+10 350 40217 40,3 7,9 12,1 6,20E+09 3,44E+10

174 21731 5,7 23,8 1,0 2,60E+09 1,85E+10 354 40793 29,2 14,9 13,2 7,19E+09 3,88E+10

178 22063 3,6 22,0 2,0 3,06E+09 1,93E+10 358 41368 31,3 12,4 21,5 6,81E+09 4,35E+10

182 22401 4,4 22,8 2,3 2,27E+09 1,53E+10 362 41936 39,0 6,5 9,4 6,07E+09 3,97E+10

186 22738 4,9 27,8 4,9 2,48E+09 1,50E+10 366 42479 32,4 10,2 15,2 6,71E+09 5,10E+10

190 23076 5,5 30,7 9,3 3,07E+09 2,09E+10 370 43021 26,9 12,8 20,2 5,52E+09 4,15E+10

194 23414 3,2 25,0 18,8 3,11E+09 2,12E+10 374 43564 22,4 19,8 9,4 8,57E+09 6,39E+10

198 23751 7,2 18,8 21,6 3,11E+09 2,02E+10 378 44107 21,7 14,0 13,7 1,12E+10 8,41E+10

202 24071 6,1 28,2 4,4 3,10E+09 2,69E+10 382 44650 12,5 9,9 18,6 7,02E+09 5,68E+10

206 24374 10,7 21,3 2,7 3,44E+09 3,31E+10 386 45193 11,9 12,2 19,2 1,14E+10 8,55E+10

210 24676 12,6 30,1 2,0 3,01E+09 2,36E+10 390 45736 19,3 8,0 17,6 8,27E+09 6,22E+10

214 24978 22,1 24,8 0,7 2,67E+09 1,98E+10 394 46279 26,1 12,0 4,9 1,37E+10 7,94E+10

218 25281 31,3 16,3 0,7 4,91E+09 3,83E+10 398 46821 19,6 11,9 23,1 1,10E+10 8,39E+10

222 25583 43,8 19,0 0,7 4,85E+09 3,76E+10 402 47364 19,9 10,4 21,5 1,07E+10 8,59E+10

226 25958 43,4 8,4 1,0 4,35E+09 1,85E+10 406 47907 23,6 6,5 17,5 5,81E+09 4,76E+10

230 26552 37,3 9,3 1,7 3,80E+09 1,32E+10 410 48450 14,9 10,5 15,9 1,03E+10 8,30E+10

234 27146 8,9 43,4 1,4 3,17E+09 1,12E+10 414 48993 13,5 19,3 10,8 1,14E+10 8,95E+10

238 27739 14,4 33,4 1,6 3,60E+09 1,41E+10 418 49568 22,9 7,6 12,7 1,17E+10 7,73E+10

242 28333 15,9 28,0 2,4 4,20E+09 1,78E+10 422 50238 25,7 13,1 14,7 1,20E+10 7,60E+10

245 28778 29,3 29,0 3,8 3,74E+09 1,62E+10 426 50908 34,1 10,1 10,1 1,29E+10 8,20E+10

250 29520 31,2 18,8 10,5 5,97E+09 2,69E+10 430 51578 25,9 11,6 11,0 1,00E+10 6,46E+10

254 29693 34,8 18,5 2,3 5,70E+09 9,54E+10 434 52248 17,8 27,9 8,7 1,32E+10 7,87E+10

258 29866 40,7 14,3 5,5 5,95E+09 9,02E+10 438 52918 19,3 17,8 13,2 1,12E+10 7,25E+10

262 30038 40,9 9,7 3,5 7,27E+09 8,34E+10 442 53588 21,5 19,6 11,7 9,24E+09 5,93E+10

266 30211 33,7 14,4 2,2 6,96E+09 7,93E+10 446 54258 18,4 16,8 10,8 9,48E+09 6,40E+10

111

Chapter 3

Figure 3.3

Core

dep

th (

cm)

Cal

endar

age

(ka

B.P

.)

Em

ilia

nia

huxl

eyi

(%)

Gep

hyr

oca

psa

eri

csonii

(%

)

Gep

hyr

oca

psa

oce

anic

a (

%)

Abso

lute

abundan

ces

(no.g

-1 s

ed.)

Acc

um

ula

tion r

ates

(no.c

m-2 k

a-1)

450 54928 29,0 14,4 10,9 9,68E+09 6,00E+10

454 55598 20,3 30,5 8,5 1,36E+10 9,00E+10

458 56268 20,9 23,2 12,2 1,02E+10 6,45E+10

462 56938 30,6 12,3 11,7 1,06E+10 6,49E+10

466 57608 31,9 12,6 8,8 1,11E+10 7,14E+10

470 58278 29,1 12,9 12,3 1,01E+10 6,45E+10

474 58948 31,6 26,5 5,8 1,10E+10 6,88E+10

478 59618 33,1 24,8 5,1 7,42E+09 4,66E+10

112

Chapter 3

Figure 3.5

Core

dep

th (

cm)

Cal

endar

age

(B.P

.)

Flo

risp

haer

a p

rofu

nda

(%

)

Core

dep

th (

cm)

Cal

endar

age

(B.P

.)

Em

ilia

nia

huxl

eyi

(%)

96 14335 24,5 270 30384 27,4

98 14533 29,0 274 30557 13,6

100 14730 48,0 278 30889 18,9

102 14967 54,0 282 31275 10,2

106 15440 55,3 286 31660 24,6

110 15914 35,3 290 32046 21,2

114 16388 37,5 294 32432 24,2

118 16861 35,5 298 32817 30,5

122 17335 36,0 302 33299 34,1

126 17776 36,0 306 33876 22,9

130 18118 29,9 310 34454 42,2

134 18460 37,7 314 35032 25,9

138 18803 40,0 318 35609 32,9

142 19145 37,9 322 36187 34,2

145 19402 43,5 326 36764 22,3

150 19830 32,1 330 37339 21,3

154 20147 29,7 334 37915 26,6

158 20464 25,2 338 38491 27,8

162 20780 32,5 342 39066 38,3

166 21097 27,4 346 39642 40,7

170 21414 40,0 350 40217 31,1

174 21731 40,3 354 40793 22,4

178 22063 44,4 358 41368 23,8

182 22401 30,5 362 41936 36,1

186 22738 28,8 366 42479 29,2

190 23076 24,1 370 43021 28,3

194 23414 18,5 374 43564 36,7

198 23751 17,1 378 44107 30,4

202 24071 22,4 382 44650 44,6

206 24374 27,1 386 45193 41,3

210 24676 35,4 390 45736 41,2

214 24978 20,7 394 46279 40,2

218 25281 19,2 398 46821 27,2

222 25583 14,7 402 47364 31,3

226 25958 25,6 406 47907 40,5

230 26552 32,3 410 48450 43,8

234 27146 16,7 414 48993 30,4

238 27739 28,9 418 49568 25,4

242 28333 23,8 422 50238 30,9

245 28778 15,6 426 50908 24,3

250 29520 18,5 430 51578 30,8

254 29693 14,9 434 52248 26,5

258 29866 26,7 438 52918 32,8

262 30038 31,8 442 53588 30,9

266 30211 34,6 446 54258 29,4

113

Chapter 3

Figure 3.5

Core

dep

th (

cm)

Cal

endar

age

(ka

B.P

.)

Flo

risp

haer

a p

rofu

nda

(%

)

450 54928 26,8

454 55598 21,6

458 56268 21,9

462 56938 25,3

466 57608 30,0

470 58278 27,0

474 58948 24,5

478 59618 20,9

114

Chapter 4

Figure 4.3, 4.4 and 4.5

Core

dep

th (

cm)

Cal

endar

age

(A.D

.)

Cal

endar

age

(B.P

.)

Gep

hyr

oca

psa

spp. ab

solu

te

abundan

ces

(no.g

-1 s

ed.)

Gep

hyr

oca

psa

eri

csonii

(%

)

Gep

hyr

oca

psa

oce

anic

a (

%)

G. sp

p.

— T

OT

AL

(%

)

Em

ilia

nia

huxl

eyi

(%)

G. sp

p.

— E

. huxl

eyi

(%)

0,25 1820 130 9,66E+08 1,9 4,7 6,6 48,1 12,0

0,75 1804 146 1,65E+09 4,4 6,2 10,6 39,6 21,1

1,25 1788 162 5,76E+08 0,3 2,8 3,1 31,5 8,9

1,75 1773 178 8,18E+08 0,5 4,3 4,8 31,2 13,4

2,25 1757 193 8,64E+08 1,4 3,4 4,9 33,5 12,7

2,75 1741 209 5,74E+08 1,0 3,8 4,8 35,6 11,8

3,25 1725 225 8,25E+08 2,3 3,6 6,0 30,5 16,4

3,75 1709 241 1,15E+09 3,6 4,5 8,1 36,7 18,1

4,25 1694 257 1,44E+09 2,9 5,1 7,9 35,9 18,1

4,75 1678 272 7,79E+08 1,5 5,1 6,6 26,0 20,2

5,25 1662 288 1,30E+09 2,6 3,9 6,5 35,6 15,4

5,75 1646 304 9,00E+08 3,2 5,7 8,8 30,6 22,4

6,25 1630 320 9,95E+08 1,6 6,9 8,5 24,3 26,0

6,75 1614 336 7,04E+08 2,8 3,4 6,2 34,2 15,3

7,25 1598 353 2,74E+09 6,0 3,5 9,4 38,4 19,7

7,75 1593 358 9,09E+08 2,0 5,9 7,9 36,0 18,0

8,25 1588 363 1,23E+09 4,6 5,2 9,9 25,3 28,1

8,75 1583 368 1,19E+09 6,8 4,8 11,6 32,3 26,5

9,25 1578 373 5,35E+08 2,0 2,7 4,7 29,2 13,7

9,75 1573 378 1,89E+09 4,8 5,1 9,9 30,7 24,4

10,25 1559 391 7,27E+09 10,8 5,4 16,2 37,6 30,1

10,75 1550 400 3,68E+09 7,8 3,7 11,5 38,5 23,0

11,25 1542 408 6,63E+08 3,4 5,6 9,1 46,9 16,2

11,75 1534 416 2,63E+09 7,2 7,2 14,4 28,1 33,8

12,25 1525 425 2,35E+09 8,6 2,4 11,0 36,1 23,4

12,75 1517 433 2,46E+09 7,4 4,8 12,2 30,9 28,3

13,25 1509 441 1,86E+09 5,3 3,9 9,3 32,0 22,4

13,75 1500 450 2,61E+09 10,0 4,4 14,3 25,9 35,7

14,25 1492 458 1,53E+09 7,1 3,8 10,9 31,2 25,9

14,75 1484 466 2,55E+09 7,7 3,8 11,5 35,5 24,5

15,25 1475 475 8,02E+08 1,6 3,6 5,2 36,4 12,5

15,75 1467 483 1,49E+09 3,8 4,4 8,1 37,2 17,9

16,25 1459 491 1,42E+09 3,3 5,1 8,4 25,4 24,8

16,75 1451 499 1,71E+09 4,0 4,6 8,6 33,0 20,7

17,25 1442 508 2,55E+09 9,6 3,5 13,2 28,9 31,3

17,75 1434 516 1,49E+09 3,2 3,4 6,6 34,2 16,2

18,25 1426 524 1,28E+09 3,9 4,5 8,3 24,4 25,5

18,75 1417 533 2,48E+09 8,2 3,4 11,6 29,0 28,6

19,25 1388 562 1,37E+09 3,0 3,0 5,9 29,3 16,8

19,75 1358 592 8,16E+08 2,3 2,3 4,6 39,3 10,5

20,25 1329 621 2,08E+09 3,6 5,2 8,8 34,9 20,1

20,75 1299 651 9,10E+08 2,0 2,6 4,6 37,4 10,9

21,25 1270 680 1,37E+09 3,1 2,5 5,6 34,4 14,1

21,75 1240 710 1,30E+09 2,8 2,8 5,6 34,5 14,1

22,25 1211 739 1,10E+09 3,3 3,3 6,5 38,7 14,5

115

Chapter 4

Figure 4.3, 4.4 and 4.5

Core

dep

th (

cm)

Cal

endar

age

(A.D

.)

Cal

endar

age

(B.P

.)

Gep

hyr

oca

psa

spp. ab

solu

te

abundan

ces

(no.g

-1 s

ed.)

Gep

hyr

oca

psa

eri

csonii

(%

)

Gep

hyr

oca

psa

oce

anic

a (

%)

G. sp

p.

— T

OT

AL

(%

)

Em

ilia

nia

huxl

eyi

(%)

G. sp

p.

— E

. huxl

eyi

(%)

22,75 1181 769 1,42E+09 2,6 5,5 8,1 37,5 17,8

23,25 1154 796 1,15E+09 3,0 4,5 7,5 36,1 17,2

23,75 1130 820 1,97E+09 3,2 4,4 7,6 31,6 19,4

24,25 1105 845 1,44E+09 3,2 3,8 6,9 31,2 18,2

24,75 1081 869 5,08E+08 0,7 2,6 3,3 26,3 11,1

25,25 1065 885 9,95E+08 3,1 3,7 6,9 32,4 17,5

25,75 1058 892 1,67E+09 4,0 5,0 9,0 31,6 22,1

26,25 1051 899 3,21E+09 10,9 3,4 14,3 27,2 34,5

26,75 1044 906 1,31E+09 5,6 3,3 8,9 24,7 26,4

27,25 1037 913 7,99E+08 2,4 4,0 6,4 36,2 15,0

27,75 1030 920 8,51E+08 1,7 2,6 4,3 27,4 13,5

28,25 1023 927 1,45E+09 2,5 4,3 6,7 24,2 21,8

28,75 1016 934 7,65E+08 2,1 3,0 5,2 27,3 15,9

29,25 1009 941 9,68E+08 4,3 2,0 6,2 29,2 17,6

29,75 994 956 2,53E+09 6,0 2,6 8,7 34,6 20,0

30,25 970 980 1,48E+09 4,5 4,8 9,3 36,1 20,4

30,75 947 1003 2,30E+09 5,6 1,8 7,4 34,0 17,9

31,25 923 1027 8,12E+08 3,1 1,8 4,9 35,9 12,0

31,75 900 1050 1,41E+09 2,9 5,1 8,0 36,0 18,2

32,25 876 1074 1,67E+09 6,3 2,8 9,1 34,9 20,7

32,75 853 1097 1,17E+09 4,2 4,9 9,2 48,4 15,9

33,25 829 1121 1,66E+09 5,1 4,1 9,2 37,5 19,7

33,75 806 1144 1,39E+09 2,2 3,8 6,1 36,9 14,2

34,25 782 1168 1,24E+09 1,9 4,5 6,4 30,5 17,4

34,75 759 1191 1,06E+09 1,6 3,8 5,4 35,2 13,3

35,25 745 1205 1,08E+09 1,8 4,2 6,0 29,5 16,9

35,75 740 1210 1,38E+09 3,8 1,6 5,4 29,8 15,5

36,25 736 1214 1,99E+09 3,3 3,3 6,5 32,4 16,8

36,75 731 1219 1,44E+09 5,0 2,2 7,2 32,1 18,4

37,25 727 1224 7,11E+08 3,4 1,2 4,6 25,6 15,3

37,75 722 1228 2,35E+09 8,0 2,4 10,3 33,6 23,5

38,25 718 1233 1,08E+09 3,4 4,0 7,3 36,6 16,7

38,75 691 1260 9,57E+08 1,0 6,7 7,7 33,0 18,9

39,25 661 1289 1,50E+09 5,6 2,7 8,3 31,8 20,7

39,75 632 1318 1,61E+09 5,9 2,2 8,1 33,1 19,7

40,25 602 1348 1,94E+09 5,2 2,6 7,8 38,0 17,0

40,75 573 1377 3,19E+09 12,1 2,1 14,2 26,3 35,0

41,25 543 1407 3,27E+09 7,2 5,7 12,9 19,8 39,4

41,75 514 1436 1,32E+09 5,1 2,9 8,0 31,4 20,3

42,25 484 1466 1,27E+09 7,4 4,2 11,6 27,3 29,8

42,75 455 1495 2,74E+09 11,7 5,8 17,5 27,8 38,6

43,25 426 1525 3,32E+09 11,7 4,2 16,0 25,1 38,9

43,75 396 1554 2,55E+09 10,3 3,0 13,3 23,9 35,8

44,25 367 1583 1,90E+09 5,8 3,7 9,5 25,5 27,2

44,75 337 1613 1,22E+09 5,9 2,6 8,5 37,7 18,4

116

Chapter 4

Figure 4.3, 4.4 and 4.5

Core

dep

th (

cm)

Cal

endar

age

(A.D

.)

Cal

endar

age

(B.P

.)

Gep

hyr

oca

psa

spp. ab

solu

te

abundan

ces

(no.g

-1 s

ed.)

Gep

hyr

oca

psa

eri

csonii

(%

)

Gep

hyr

oca

psa

oce

anic

a (

%)

G. sp

p.

— T

OT

AL

(%

)

Em

ilia

nia

huxl

eyi

(%)

G. sp

p.

— E

. huxl

eyi

(%)

45,25 308 1642 1,94E+09 5,9 6,3 12,2 28,1 30,2

45,75 278 1672 1,72E+09 7,9 2,6 10,5 31,5 25,0

46,25 249 1701 1,93E+09 5,2 4,3 9,6 37,3 20,4

46,75 220 1731 1,69E+09 5,5 5,5 10,9 34,1 24,3

47,25 190 1760 2,60E+09 7,4 6,5 13,9 27,2 33,8

47,75 161 1789 2,66E+09 8,0 6,8 14,8 23,8 38,3

117

118

119

Curriculum Vitae

Personal data

Name: Heiko-Lars Legge

Date of birth: 26. January 1973

Place of birth: Bochum

Marital status: single

Nationality: German

School education

1979 – 1983 Gemeinschaftsgrundschule, Alte Wittener Str. 19, Bochum

1983 – 1992 Erich Kästner-Schule, Markstr. 189, Bochum

“Compulsory community service” (Ziviler Ersatzdienst)

1992 – 1993 Behinderten-Schule, Bochum-Langendreer

University education and scientific activity

1994 – 2002 Study of biology at the Department of Biology, Ruhr-University

Bochum

Diploma-Thesis (“Die Coccolithophoriden des Roten Meeres:

Ökologie und Biogeographie der letzten 24.000 Jahre”) at the

department of Geology, Mineralogy and Geophysics, Ruhr-

University Bochum.

2002 Participant Meteor cruise M52/3 (Red Sea)

2003 – 2007 Ph.D. student and research assistant at the department of

Geology, Mineralogy and Geophysics, Ruhr-University Bochum

Documents

Calcareous nannoplankton of the Red Sea - Ruhr-Universität Bochum