The Glacial and Interglacial History of the Black Sea During the Last 134 Ka: About Meltwater Events, Abrupt Temperature Changes

Home , History of the Black Sea, Last Glacial Maximum

The glacial and interglacial history of the Black Sea during the last 134 ka:

About meltwater events, abrupt temperature changes, and sapropels

Inauguraldissertation

zur Erlangung des akademischen Grades

eines Doktors der Naturwissenschaften (Dr. rer. nat.)

der Mathematisch-Naturwissenschaftlichen Fakultät

der Ernst-Moritz-Arndt-Universität Greifswald

vorgelegt von Antje Wegwerth

Greifswald, 05.11.2015

Dekan: Prof. Dr. Klaus Fesser

1. Gutachter: Prof. Dr. Helge W. Arz 2. Gutachter: Prof. Dr. Gerhard Bohrmann

Tag der Promotion: 18.01.2016

TABLE OF CONTENTS

ABSTRACT ...... 1

KURZFASSUNG ...... 3

INTRODUCTION ...... 6

1. Palaeoclimate over the last 150,000 years ...... 6 2. The Black Sea ...... 10 2.1. Geography, oceanography, and modern climate in the Black Sea ...... 10 2.2. Palaeoclimate in the Black Sea region ...... 15

3. Motivation and objectives ...... 18 4. Study area - the ‘Archangelsky Ridge’ ...... 20 5. Materials and methods ...... 21 5.1 Sediment cores and stratigraphy ...... 21 5.2 Inorganic, organic, and ostracod geochemistry ...... 24

6. Manuscript outline and personal contribution ...... 27

MANUSCRIPTS ...... 30

Meltwater events and the Mediterranean reconnection at the Saalian-Eemian transition in the Black Sea ...... 31

Black Sea temperature response to glacial millennial-scale climate variability ...... 74

Northern hemisphere climate control on the environmental dynamics in the glacial Black Sea “Lake” ...... 100

Redox evolution during Eemian and Holocene sapropel formation in the Black Sea .. 129

CONCLUSIONS AND OUTLOOK ...... 155

REFERENCES ...... 159

APPENDIX ...... 179 Related co-author publications and personal contribution ...... 179

Vegetation and environmental changes in Northern Anatolia between 134 and 119 ka recorded in Black Sea sediments ...... 180

Orbital- and millennial-scale environmental changes between 64 and 20 ka BP recorded in Black Sea sediments ...... 182

Environmental and Climate Dynamics During the Last Two Glacial Terminations and Interglacials in the Black Sea/Northern Anatolian Region ...... 184

CURRICULUM VITAE ...... 187

LIST OF PUBLICATIONS AND OTHER SCIENTIFIC ACTIVITIES ...... 188

ACKNOWLEDGEMENTS ...... 191

ABSTRACT

The Black Sea experienced fundamental environmental changes during the last glacial- interglacial transitions. During the last 670,000 years, the Black Sea was at least twelve times connected to Mediterranean Sea, received saltwater via the Bosporus strait, and evolved to a brackish anoxic water body. A lowered global sea level during glacials caused isolation of the basin from the open ocean, and the Black Sea became limnic and well-oxygenated. The last glacial-interglacial history of the Black Sea is relatively well understood and demonstrates the high sensitivity of this basin to global climate and environmental changes. Previous studies particularly focussed on the evolution during the last glacial with meltwater pulses, warming during the glacial-interglacial transition, and the development from a ventilated lake to the present euxinic/brackish water body. Apart from the interglacial warming, the Black Sea sediments clearly recorded short-term abrupt temperature changes associated with cooling during Heinrich events and the Younger Dryas as well the Bølling-Allerød warming, which occurred over the northern hemisphere. However, our knowledge about the Black Sea history before 40,000 BP is comparatively poor even though crucial for understanding hemisphere- wide atmospheric teleconnection patterns and climate mechanisms during older glacials and interglacials.

A multiproxy approach has been applied on three gravity cores and surface sediment from the southeastern Black Sea comprising ostracod geochemistry (Mg/Ca, Sr/Ca, U/Ca, 87Sr/86Sr), major and trace elements (Al, Ca, Fe, K, Ti, Mo, Re, Sr, W, Zr) and organic K’ biomarkers (n-alkanes, alkenones, U 37-palaeotemperatures, glycerol dialkyl glycerol tetraethers, TEX86-palaeotemperatures, BIT-index). The cores cover the last 134,000 a and provide new findings concerning the last and penultimate glacial-interglacial transitions (12,000- 0 a BP; 134,000-120,000 a BP) as well as the abrupt climate changes during the last glacial period (64,000-20,000 a BP).

The major topics of this work are i) the penultimate glacial-interglacial transition (Saalian-Eemian), ii) the environmental conditions in the Black Sea “Lake” during abrupt climate oscillations of the last glacial period, iii) and the comparison of the redox evolution during Eemian and Holocene sapropel formation.

Two meltwater pulses caused a pronounced freshening of the Black Sea “Lake” during the ending penultimate glacial, which originated from the melting Fennoscandian Ice Sheet. Due to unusually high radiogenic Sr-isotope signatures of benthic ostracods, a potential

1 ABSTRACT

Himalayan source communicated via the Caspian Sea is also likely. During the glacial- interglacial transition the temperatures in the Black Sea increased from 9°C to 17°C and the associated global sea-level rise allowed the reconnection between the Mediterranean and Black Seas around 128,000 a BP. Eemian sapropel formation started shortly after the intrusion of saltwater and the water body became gradually euxinic. In comparison with the Holocene sapropel, the Eemian proxy records imply warmer and stronger euxinic conditions and distinctly higher enrichments of redox-sensitive trace elements like e.g. Mo, Re, and W. Because the seawater forms the ultimate source for several trace metals, these enrichments were most likely favoured by the higher salinity due to a ca. 10 m higher sea level and enhanced Mediterranean Sea - Black Sea water exchange.

Based on biomarker analyses, lake surface temperatures could be calculated for the first time for the period between 64,000 and 20,000 a BP, which includes the Marine Isotope Stage (MIS) 3. Abrupt stadial/interstadial temperature changes with amplitudes of up to 4°C in the Black Sea “Lake” clearly resemble the Greenland Dansgaard-Oeschger pattern. However, an exceptional cooling during the so-called Heinrich events is not evident from our cores. This finding agrees with modelling results proposing a deeper penetration of regular Dansgaard- Oeschger cycles into the Eurasian continent when compared with the Heinrich events. During the warm and more humid interstadials, the Black Sea “Lake” became fresher and more productive and the water level probably increased. During the colder and more arid stadials the freshwater supply was decreased and productivity was low. Aridity and stronger westerly winds favoured the input of aeolian transported detritus. The long-term pattern from 64,000 to 20,000 a BP demonstrates a strong influence of orbital-driven changes in the Eurasian ice volume and associated atmospheric circulation patterns over the Black Sea region.

The present multi-proxy study demonstrates that the sediments from the SE Black Sea clearly record not only orbital- but also millennial-scale climate and environmental changes and thus represent an important continental archive for climate change bridging the North Atlantic-Eurasian corridor.

2 KURZFASSUNG

KURZFASSUNG

Das Schwarze Meer war im Zuge der Wechsel von Glazialen zu Interglazialen fundamentalen Umweltveränderungen ausgesetzt. Während der letzten 670.000 Jahre war das Schwarze Meer mindestens zwölfmal mit dem Mittelmeer verbunden, wodurch salzhaltiges Wasser über die Meerenge des Bosporus in das Becken eintrat und ein brackischer, anoxischer Wasserkörper entstand. Ein niedriger Meeresspiegel während der Glaziale bedingte jedoch die Isolation des Beckens vom offenen Ozean und das Schwarze Meer entwickelte sich erneut zu einem limnischen und durchlüftetem System. Die Entwicklung des Schwarzen Meeres während des letzten Glazial-Interglazial Übergangs ist weitgehend bekannt. Es hat sich gezeigt, dass das Schwarze Meer besonders sensibel auf globale Klima- und Umweltveränderungen reagiert. In früheren Arbeiten wurde besonderes Augenmerk auf Schmelzwasserereignisse, den Temperaturanstieg während des Glazial-Interglazial Übergangs sowie die Entwicklung von einem oxischen See zum heutigen euxinischen Brackwasserkörper gerichtet. Abgesehen von der Erwärmung zum Interglazial, haben die Sedimente des Schwarzen Meeres aber auch abrupte, kurzfristige Temperaturveränderungen archiviert, die im Zusammenhang mit denen auf der Nordhemisphäre auftretenden Abkühlungen während der Heinrich Ereignisse und der Jüngeren Dryas aber auch mit der Erwärmung während des Bølling-Allerøds stehen. Jedoch ist unser Kenntnisstand hinsichtlich der Zeit vor 40.000 Jahren für das Schwarze Meer begrenzt. Die Entwicklungsgeschichte während früherer Glaziale und Interglaziale ist jedoch unerlässlich für unser Verständnis der atmosphärischen Fernwirkung und klimatischer Prozesse im globalen Kontext.

An drei Schwerelotkernen und Oberflächensedimenten aus dem südöstlichen Schwarzen Meer wurde im Rahmen dieser Arbeit ein Multiproxy-Ansatz angewandt, der die Analyse der Geochemie von benthischen Ostrakoden (Mg/Ca, Sr/Ca, U/Ca, 87Sr/86Sr), Haupt- und Spurenelemente (Al, Ca, Fe, K, Ti, Mo, Re, Sr, W, Zr) sowie organische Biomarker (n- K’ Alkane, Alkenone, U 37-Paläotemperaturen, Glycerol-Dialkyl-Glycerol-Tetraether, TEX86- Paläotemperaturen, BIT-Index) der Sedimente umfasste. Die Kerne erfassen die letzten 134.000 Jahre und liefern neue Einblicke in die letzten beiden Glazial-Interglazial Übergänge (12.000 J.v.h. bis in die Gegenwart; 134.000 bis 120.000 J.v.h.). Weiterhin wurden abrupte Klimaschwankungen während des letzten Glazials im Zeitraum zwischen 64.000 und 20.000 J.v.h., der somit das Marine Isotopenstadium (MIS) 3 umfasst, detailliert untersucht.

3 KURZFASSUNG

Die primären Inhalte dieser Arbeit umfassen i) den vorletzten Glazial-Interglazial Übergang (Saale-Eem), ii) die Umweltbedingungen des limnischen Schwarzen Meeres während abrupter Klimaschwankungen im MIS 3 und iii) den Vergleich der Entwicklung der Redoxbedingungen während der Sapropelbildung im Eem und Holozän.

Während des ausgehenden vorletzten Glazials führten zwei Schmelzwasserereignisse zu einer erheblichen Süßwasserzufuhr in das damalige limnische Schwarze Meer. Ähnlich des letzten Glazials könnte dieses Schmelzwasser seinen Ursprung im schwindenden Fennoskandischen Eisschild gehabt haben. Die Analyse von Strontium-Isotopen in Schalen benthischer Ostrakoden zeigte jedoch eine erheblich radiogenere Signatur an, die auf einen zusätzlichen Beitrag von Schmelzwasser aus dem Himalaya, das über das Kaspische Meer transportiert wurde, hindeutet.

Im Verlauf des vorletzten Glazial-Interglazial Übergangs (Termination II) stiegen die Temperaturen im Schwarzen Meer von ca. 9°C auf ca. 17°C an. Der globale Meeresspiegelanstieg führte vor etwa 128.000 Jahren zur Wiederverbindung des Mittelmeeres mit dem Schwarzen Meer. Im Zuge des Salzwassereinstroms entwickelten sich euxinische Bedingungen und die Bildung des sog. Eem-Sapropels setzte ein. Im Vergleich zum Holozän- Sapropel lassen die Analysen des Eem-Sapropels wärmere und stärker euxinische Bedingungen vermuten, die mit erheblich höheren Anreicherungen von redox-sensitiven Spurenelementen wie z.B. Mo, Re oder W einhergehen. Da das Meerwasser die entscheidende Quelle für eine Vielzahl von Spurenmetallen darstellt, wurden diese hohen Anreicherungen insbesondere durch den etwa zehn Meter höheren Meeresspiegel und dem daran gekoppelten höheren Salzgehalt bzw. dem stärkeren Wasseraustausch zwischen Mittelmeer und Schwarzem Meer begünstigt.

Basierend auf Biomarker-Analysen konnten erstmals Oberflächenwassertemperaturen des Schwarzen Meeres für den Zeitraum zwischen 64.000 und 20.000 J.v.h., der somit das das MIS 3 umfasst, berechnet werden. Abrupte Stadial/Interstadial-assoziierte Temperatur- schwankungen mit Amplituden von bis zu 4°C spiegeln eindeutig die sog. Dansgaard- Oeschger Zyklen wider. Außergewöhnliche Abkühlungen während der sogenannten Heinrich- Ereignisse zeichnen sich in den Kernen allerdings nicht ab. Diese Beobachtung deckt sich mit Modellergebnissen, die im Vergleich zu den Heinrich Ereignissen auf eine stärkere Ausdehnung der regulären Dansgaard-Oeschger Zyklen in den Eurasischen Kontinent hindeuten. Die warmen und humiden Interstadiale zeichnen sich zudem durch einen reduzierten Salzgehalt und eine erhöhte Produktivität aus. Es ist außerdem anzunehmen, dass

4 KURZFASSUNG der Seespiegel während der Interstadiale anstieg. Während der kälteren und ariden Stadiale nahmen Süßwasserzufuhr und Produktivität hingegen ab. Trockenere Bedingungen und stärkere Westwinde begünstigten äolischen Eintrag von detrischem Material. Langfristig gesehen war das Schwarze Meer zwischen 64.000 und 20.000 J.v.h. stark durch die orbital gesteuerten Veränderungen des Eurasischen Eisschildes und den verbundenden atmosphärischen Zirkulationsmustern geprägt.

Diese Arbeit belegt, dass die Sedimente des südöstlichen Schwarzen Meeres die Klima- und Umweltveränderungen nicht nur im orbitalen sondern auch im kurzzeitigen Maßstab aufgezeichnet haben. Damit stellt das Schwarze Meer, inmitten des Atlantisch-Eurasischen Sektors, ein überaus wichtiges kontinentales Archiv für Klimaentwicklungen dar.

5 INTRODUCTION

INTRODUCTION

1. Palaeoclimate over the last 150,000 years

During the last two glacials, the Saalian and Weichselian, large parts of the northern hemisphere were covered by continental ice sheets. During the penultimate glacial maximum at ca. 140 ka BP (Lambeck and Nakada, 1992) the Eurasian Ice Sheet extended considerably wider into the continent resulting in colder continental climate conditions (Fig. 1; Svendsen et al., 2004) when compared with the last glacial maximum (LGM) at 22.1 ± 4.3 ka (Shakun and Carlson, 2010). The transitions from glacials to interglacials, the so-called terminations, involved disintegration of the ice sheets and global sea-level rise (Fig. 2 G-H; e.g. Bintanja and van de Wal, 2008; Grant et al., 2012).

Fig. 1: The Eurasian Ice Sheet extent during A) the penultimate glacial >140 ka BP (Saalian) and its evolution during the last glacial (Weichselian) at about B) 60 ka BP, C) 55-45 ka BP, and during D) the last glacial maximum (LGM) at about 20 ka BP (A, B, D from Svendsen et al., 2004; C from Larsen et al., 2006).

The variability of the global sea level in response to shrinking and growing ice sheets is strongly related to periodical changes in the in the earth’s orbital parameters (Milanković, 1969; Hays et al., 1976). The interplay of the temporarily varying obliquity (41 ka period) and precession (22 ka period) of the earth’s axis as well as the eccentricity (100 ka period) of the

6 INTRODUCTION earth’s orbit affect the insolation on earth (Fig. 2 A-C; Milanković, 1969; Hays et al., 1976; Laskar et al., 2004; Ruddiman, 2006).

Fig. 2: Orbital parameters (A) eccentricity, B) obliquity and C) precession) affecting D) insolation and climate on Earth during the last 150 ka (based on La2004; Laskar et al., 2004; JJA: June, July, August; solar constant 1368 W m-2). E) δ18O from NGRIP (North Greenland Ice Core Project; NGRIP members, 2004). F) alkenone-based sea surface temperatures from the mid-latitude North Atlantic (Martrat et al., 2007). E) and F) demonstrate warming during the present and last interglacial, cooling during the last two glacials, and temperature variability during MIS 3 with Greenland interstadials (red numbers, warming) during Dansgaard-Oeschger cycles and Heinrich Events (blue numbers, cooling). G) The relative sea level (RSL) from Grant et al. (2012) and H) the Eurasian Ice Volume from Bintanja and van de Wal (2008) reflect the climatic variability during the last 150 ka. Sea level above the present Bosporus sill depth (dotted line) indicates potential entrance of Mediterranean water into the Black Sea, whereas a sea level below the sill denotes periods of potential isolation. MIS denotes the Marine Isotope Stages 1-6.

7 INTRODUCTION

Because of the maximum continental extent especially the northern summer insolation at 65° (Fig. 2 D) triggers along with varying atmospheric CO2 levels the temperature variability and the build-up/disintegration of ice sheets causing considerable sea-level fluctuations over glacial-interglacial timescales (Fig. 2 G). These variations are strongly related to changes in the oceanic and atmospheric circulation patterns controlling the climate dynamics over the northern hemisphere.

During the last termination (T I), ice sheet melting resulted in meltwater pulses and sea- level rise, but the warming was shortly interrupted during the Younger Dryas cold episode over the northern hemisphere at around 12 ka BP (e.g. Bard et al., 1990; Arz et al., 2003; Siddall et al., 2006). A combination of a weakened Atlantic Meridional Overturning Circulation (AMOC), an anomalous northerly flow of cold polar air masses over Europe, and radiative cooling due to enhanced dust fluxes and reduced atmospheric greenhouse gas concentrations to date best explains the Younger Dryas cooling (Renssen et al., 2015). Whether such a Younger Dryas-like event also occurred during the penultimate termination (T II) still remains controversial (e.g. Sarnthein and Tiedemann, 1990; Cannariato and Kennett, 2005; Risebrobakken et al., 2006; Carlson, 2008). The deglacial warming initiated the present Holocene and the Eemian interglacials. Compared to the Holocene, the Eemian was in all probability warmer and moister and was characterised by a ca. 10 m higher sea level (Fig. 2 F, G; e.g. Pailler and Bard, 2002; Grant et al., 2012; Martrat et al., 2014).

Aside from the long-term climate changes over glacial-interglacial cycles, the last glacial, more precisely the Marine Isotope Stage 3 (MIS 3; ~ 57-29 ka BP), was punctuated by pronounced (multi)millennial- to centennial-scale climate variability expressed by severe alternations in temperature, rainfall and wind intensity (e.g. Heinrich, 1988; Bond et al., 1993; Dansgaard et al., 1993; Huber et al., 2006; Sima et al., 2013). The abrupt climate variations between cold stadials to warmer interstadials are referred to as Dansgaard-Oeschger (DO) cycles and were originally detected in Greenland ice cores (Fig. 2 E, F; e.g. Bond et al., 1993; Dansgaard et al., 1993). While abrupt interstadial warming occurred on a decadal timescale the following cooling periods persisted on centennial- to millennial timescales (e.g. Ganopolski and Rahmstorf, 2001). There is evidence, that the DO-cycles affected the climate over some parts of the northern and southern hemispheres (e.g. Bond et al., 1993; Dansgaard et a., 1993; Arz et al., 1998; Cacho et al., 1999; Wang et al., 2001) but it seems that strongest temperature changes occurred in the North Atlantic region (e.g. Ganopolski and Rahmstorf,

8 INTRODUCTION

2001; Martrat et al., 2007). The DO-climate fluctuations may have also been associated with ice sheet and global sea level fluctuations (Arz et al., 2007).

The most plausible explanation for DO climate variability is the transition in the strength of the thermohaline circulation of the North Atlantic (the ‘ocean conveyor belt’) and the associated change in ocean heat transport, which is triggered by a change in freshwater supply from ice sheets (Broecker et al., 1985; Bond et al., 1993; Ganopolski and Rahmstorf, 2001; Rahmstorf, 2002; Zhang et al., 2014). During a “stable” cold mode (stadial) deep water formation is reduced and located south of Iceland (Fig. 3 A; e.g. Rahmstorf, 2002). Such weak AMOC leads to a limited ocean heat transport from the tropics to higher northern latitudes (Ganopolski and Rahmstorf, 2001; Clement and Peterson, 2008). This situation is explained by elevated freshwater/meltwater supply thereby decreasing surface salinity and deep water formation (e.g. Clement and Peterson, 2008).

Fig. 3: Schemes of proposed modes in the Atlantic Meridional Overturning Circulation (AMOC) causing the three different climate states: A) Stadials (weak AMOC), B) Interstadials (strong AMOC); and C) Heinrich Events (off mode in AMOC). Simulations (upper panel) are from the Potsdam Institute for Climate Impact Research and are slightly modified after Ganopolski and Rahmstorf (2002). Position of the IRD belt adopted from Bard et al. (2000). The scheme of the three different modes of the ocean circulation along with a section of the Atlantic (A-C lower panel) are from Rahmstorf (2002).

9 INTRODUCTION

During interstadials, in turn, deep-water formation increases and shifted northwards into the Nordic Seas involving enhanced ocean heat transport and warming (Fig. 3 B; Ganopolski and Rahmstorf, 2001; Rahmstorf, 2002).

Another feature in the glacial climate system forms pronounced cooling episodes during some stadials occurring on multi-millennial timescales that are referred to as Heinrich Events (Fig. 2 F; e.g. Heinrich, 1988; Bond et al., 1993; Bard et al., 2000). These cooling episodes are associated with massive iceberg discharge from the Laurentide ice sheet through Hudson Strait into the North Atlantic, which is documented by elevated amounts of ice-rafted detritus (IRD) along the so-called IRD-belt (Fig. 3 C; e.g. Ruddiman, 1977; Heinrich, 1988; Bond et al., 1993; Hemming, 2004). Heinrich Events are initiated by considerably increasing freshwater supply and decreasing deep-water formation, which finally results in a collapsing AMOC (e.g. Ganopolski and Rahmstorf, 2001). The associated disruption of the ocean heat transport causes pronounced cooling especially in mid-latitude of the North Atlantic (Fig. 2 F; Heinrich, 1988; Bond et al., 1993; Ganopolski and Rahmstorf, 2001; Hemming, 2004; Martrat et al., 2007).

The (multi)millennial-scale climate variations during DO-cycles and Heinrich events as shortly outlined here, seem to have also affected the southern hemisphere, but in smaller amplitudes and possibly asynchronous to northern hemisphere DO and HE, which is known as the bipolar see-saw effect (e.g. Stocker, 1998; Rahmstorf, 2002; Lamy et al., 2004; Clement and Peterson, 2008). Recent modeling experiments illustrate that the Southern Ocean may have controlled the strength and stability of the AMOC in not unimportant extent and therefore may have contributed to the individual DO cycle characteristics in the North Atlantic region (Buizert and Schmittner, 2015). The interhemispheric climate variability during the last glacial thus represents an important feature in the global climate system.

2. The Black Sea

2.1. Geography, oceanography, and modern climate in the Black Sea

The Black Sea is a semi-enclosed sea located between south-eastern Europe and western Asia (Fig. 4). The main orogens surrounding the basin are the Carpathians and Balkans in the west, the Pontic Mountains in the south and the Caucasus in the east. Major rivers entering the Black Sea are Danube, Dniester, Dnieper, and Don in the north and

10 INTRODUCTION

Yesilirmak, Kizilirmak, and Sakarya in the south. The Black Sea is connected to the Sea of Asov by the Strait of Kertch and to the Marmara and Mediterranean Seas via the Bosporus and Dardanelles straits (Fig. 4).

Fig. 4: Maps showing the geography of the Black Sea and main rivers entering the basin. Red circles denote the location of the gravity cores studied in the present work. The maps are created with the mapping tool from http://woodshole.er.usgs.gov/mapit/.

The Black Sea covers an area of about 4.2*105 km2 and represents the Earth´s largest brackish water body with a water volume of 5.3*105 km3 (Demaison and Moore, 1980; Özsoy and Ünlüata, 1997). The maximum depth of the basin is about 2200 m with the flat abssyal plain (>2000 m depth) amounting to >60% of the total area (Özsoy and Ünlüata, 1997). Extensive shelf areas characterise the northwestern part of the basin (Fig. 4). While high salinity waters (35.5) from Sea of Marmara enter the basin via the Bosporus strait (35 mbsl), freshwater input occurs by precipitation and rivers at which the Danube contributes 50-70% of the total riverine runoff (Fig. 5 A; Özsoy and Ünlüata, 1997; Nezlin, 2008). The modern Black Sea has a positive water balance because the input of freshwater is higher than the loss of water by evaporation (Özsoy and Ünlüata, 1997).

Water column stratification is caused by a pycnocline separating less denser/saltier surface from deeper waters with a higher density and salinity (Fig. 5 B; Özsoy and Ünlüata, 1997). The pycnocline prevents vertical exchange of water masses and deep water ventilation thereby favouring the formation of a pelagic redoxcline in about 100-150 m water depth (Fig. 5 C; e.g. Özsoy and Ünlüata, 1997; Dellwig et al., 2010). Oxygen consumption and the release of hydrogen sulphide via bacterial sulphate reduction finally results in sulphidic (euxinic) deep waters. This euxinic water body currently amounts to about 87% of the water volume, and therefore the Black Sea is often considered as a modern analogue for ancient

11 INTRODUCTION anoxic basins (e.g. Demaison and Moore, 1980; Glenn and Arthur, 1985; Calvert et al., 1987; Özsoy and Ünlüata, 1997). Even though the inflow of saline water contributes a suite of redox-sensitive trace metals, the volume of intruding waters is insufficient to compensate trace metal fixation in the sulphidic waters finally leading to pronounced depletion of several metals like Mo below the redoxcline (e.g. Algeo and Lyons, 2006). Therefore, sedimentary trace metal signatures from the modern Black Sea sapropel may differ distinctly from ancient sapropels and Black Shales (e.g. Brumsack, 2006).

Fig. 5: A) Schematic cross section from the Mediterranean/Aegean Sea towards the Black Sea showing inflow and outflow of water masses with different salinity (salinities in italic numbers from Özsoy and Ünlüata, 1997). B) Profiles of temperature and salinity across the entire water column (left) and <250m (right) indicate depth of the pycnocline near the study site. C) Profiles of oxygen (O2) and hydrogen sulphide (H2S) across the water column (left: <1500 m, right: <250 m) indicate depth of the redoxcline in the SE Black Sea (data originate from M72/5 Meteor cruise 2007; 21-CTD 1; O. Dellwig, pers. comm.).

A major cyclonically meandering Rim Current and several smaller anticyclonic eddies, between the Rim Current and the coast, drive the general circulation at the surface and intermediate layers (Fig. 6; Oguz et al., 1993). The interior of the basin is dominated by a permanent basin-wide cyclonic current system and either one elongated cell or two main recurrent gyres separating the western from the eastern basin. The deeper circulation is regulated by cyclonic currents (Oguz et al., 1993).

12 INTRODUCTION

Fig. 6: Schematic view of the general circulation patterns in the upper water column of the Black Sea (after Oguz et al., 1993). The base map was created with the GMT mapping tool from http://sfb574.geomar.de/gmt-maps.html.

The climate in the Black Sea region is generally controlled by the northern hemispheric atmospheric circulation pattern over Europe. The main atmospheric features are the mid- to high latitude westerlies, the Siberian and the mid-latitude subtropical high pressure systems, the Polar Front Jet, and the Subtropical Jet (Fig. 7; Wigley and Farmer, 1982; Akçar and Schlüchter, 2005). Although more distant, the East African and Indian monsoons together with the Siberian high pressure system most likely control the dynamics of Mediterranean cyclones (Wigley and Farmer, 1982; Akçar and Schlüchter, 2005). The impact of these dynamic systems on the Black Sea climate varies in dependence of the position of the Polar Front and Subtropical Jet, which shift northwards during summer and southward during winter. Generally, the Black Sea climate is under Mediterranean influence with dry summers and wet winters (Cullen and deMenocal, 2000; Lamy et al., 2006; Göktürk et al., 2011) and is strongly related to the North Atlantic Oscillation (NAO; e.g. Cullen and deMenocal, 2000; Oguz et al., 2006; Capet et al., 2012). The NAO is expressed as the difference between the subtropical Azores high-pressure system and the subpolar Icelandic low-pressure system (Oguz et al., 2006). Strong gradients between these pressure systems during the positive winter NAO cause strong westerly winds carrying cold and dry air to the Black Sea. Negative

13 INTRODUCTION

NAO phases are related to stronger southwesterlies, which result in milder winters and less dry/more wet atmospheric conditions (Oguz et al., 2006).

Fig. 7: Idealised scheme of the modern northern hemispheric atmospheric circulation patterns with approx. position of the polar front during January and the Subtropical Jet during July influencing regional climate in the Black Sea region (after Wigley and Farmer, 1982; Akçar and Schlüchter, 2005). The red circle denotes the coring location of the present work. The base map was created with the generic mapping tool provided by woodshole.er.usgs.gov/mapit/index.html.

The polar front, the Siberian anticyclone, cold air masses from the Balkan Peninsula, and the Mediterranean low-pressure area dominate the winter climate (Fig. 7; Wigley and Farmer, 1982; Akçar and Schlüchter, 2005; Deniz et al., 2011). In summer, the northward shift of the polar front, the Azores high-pressure system, and the subtropical Jet cause more maritime conditions (Deniz et al., 2011). The regional atmospheric circulation results in variable climate regimes in the Black Sea (Kosarev et al., 2008). Warmer winters and a considerably higher monthly precipitation are characteristic for the eastern part (Kosarev et al., 2008). NE Anatolia and the eastern Black Sea are the most rain-laden regions during summer due to orographic precipitation in the Pontic and Caucasian Mountains (Türkeş et al., 2009; Deniz et al., 2011). Further moisture sources are the Black Sea itself (Fleitmann et al., 2009; Badertscher et al., 2011; Göktürk et al., 2011) and Mediterranean cyclones during

14 INTRODUCTION winter humidity (Fig. 7; e.g. Wigley and Farmer, 1982; Kwiecien et al., 2009). The mean annual air temperature in the northwestern Black Sea is about 10°C and reaches 14-15°C in the southeast (Kosarev et al., 2008). Annual temperature variations are high in the northwest and lower in the central part and the southeast (Kosarev et al., 2008). The spatial variability is most pronounced during winter, when mean February temperatures are -2°C in the northwest and 7.5°C in the southeast (Kosarev et al., 2008).

2.2. Palaeoclimate in the Black Sea region

The environmental conditions characterising the modern Black Sea with its stratified brackish water body due to intruding Mediterranean waters persist since about 8,000 to 9,000 years (e.g. Degens and Ross, 1972; Ryan et al., 2003; Bahr et al., 2008). There is evidence for at least 12 fundamental changes between brackish and limnic periods during the past 670,000 years because the sea level repeatedly fell below the Bosporus sill depth of about 35 mbsl during glacials (Fig. 2 G; Badertscher et al., 2011). These Black Sea “Lake” periods were considerably longer than the brackish interglacials (Fig. 2; Soulet et al., 2011). During the Last Glacial Maximum (LGM; at 22.1 ±4.3 ka; Shakun and Carlson, 2010) for example, the sea level of the Black Sea was likely up to 110 m lower compared to the present causing a complete restriction from the world`s ocean (Fig. 8; Constantinescu et al., 2015).

Previous studies have shown that the Black Sea reacted very sensitive to global climate changes (Murray, 1900; Ross et al., 1970; Zubakov, 1988; Schrader, 1979; Svitoch et al., 2000). The perhaps most prominent sediment cores recovered during the drilling campaign of DSDP Leg 42B with “Glomar Challenger” in 1975 allowed the investigation of the Quaternary history of the Black Sea (Ross et al., 1978). Several investigations partly cover the entire Pleistocene (e.g. Ross, 1978; Stoffers et al., 1978; Degens and Paluska, 1979; Schrader, 1979; Zubakov, 1988) but detailed studies of the Black Sea history before 40 ka BP are missing. The evolution of the Black Sea during the last 40,000 years is shortly summarised in the following.

The clayey sediments (lutite) deposited during the last glacial form the so-called Unit III (Degens and Ross, 1972). Lake surface temperatures during the ending last glacial were on average 6°C with three periods of strong cooling down to 3°C during North Atlantic Heinrich events (Soulet et al., 2011, Ménot and Bard, 2012).

15 INTRODUCTION

Fig. 8: The present situation (left) with water exchange between the Mediterranean and Black Seas showing salinity differences (numbers in italics) in oxic surface and euxinic deeper waters. The glacial situation (right) under lower sea level (-110m during LGM; Constantinescu et al., 2015) caused freshening, a ventilated oxic lake finally isolated from the Mediterranean Sea. Salinity values are from present day (Özsoy and Ünlüata, 1997) and during the last glacial from Shumilovskikh et al. (2013a). Maps were created with the GMT mapping tool from http://sfb574.geomar.de/gmt-maps.html with the option to set sea level to various depths.

A Northern Anatolian speleothem record documents a signal of cooling and drying during the Heinrich events as well (Fleitmann et al., 2009). In addition to Heinrich cooling, the stalagmite analyses clearly revealed changes in rainfall and temperature during the so- called Dansgaard-Oeschger cycles.

Studies from the northwestern Black Sea could not confirm such climate fluctuations during the Greenland stadial-interstadial transitions so far (Soulet et al., 2011; Ménot and Bard, 2012; Sanchi et al., 2014), although over-regional evidence exists for Dansgaard- Oeschger cycles in Southern and Eastern Europe (e.g. Allen et al., 1999; Tzedakis et al., 2004; Geraga et al., 2005; Fleitmann et al., 2009; Margari et al., 2009; Rousseau et al., 2011; Ehrmann et al., 2013; Sima et al., 2013; Sümegi et al., 2013).

16 INTRODUCTION

During Termination I, the last glacial-interglacial transition, four prominent Red Layer intervals (18-15.5 ka BP) are considered to reflect meltwater events (e.g. Bahr et al., 2005; Soulet et al., 2013). Previous studies suggested a meltwater origin from a potential spill-out from the Caspian Sea (fed by river Volga) via the Manych depression (Fig. 4) and/or from the Alpine Ice Caps by the Danube River (e.g. Major et al., 2002; Ryan et al., 2003; Bahr et al., 2005; Kwiecien et al., 2009; Badertscher et al., 2011). Detailed investigations on the Red Layers suggest, to date, that the meltwater was transported by the Dnieper River and originated from proglacial Lake Disna that formed during Heinrich Stadial 1 due to the disintegrating Fennoscandian Ice sheet (Fig. 2 H; Soulet et al., 2013).

Subsequent to the meltwater events, the Black Sea “Lake” surface temperatures increased up to 11°C during the Bølling-Allerød interstadial warming and decreased again to 5°C during the Younger Dryas cooling episode (Soulet et al., 2011). Thereafter, gradual warming at the beginning of the Holocene interglacial and the coupled global sea-level rise allowed the Mediterranean-Black Sea reconnection.

The last glacial termination also involved a transition to more humid conditions in the Black Sea region, which favoured the spreading of forests and the reduction of steppe vegetation (Shumilovskikh et al., 2012). The inflow of saline water led to the formation of a pycnocline in the Black Sea and restricted deep-water ventilation favoured the formation of the redoxcline. The gradually redoxcline rise was interrupted at ca. 5.3 ka BP (Eckert et al., 2013) and reached the present situation with sapropel formation under euxinic conditions. The main sapropel layer, deposited between 7 and 2.7 ka BP, forms the so-called Unit II (Degens and Ross, 1972). After a redoxcline drop until 2.7 ka BP, the Black Sea reached its present state (Unit I) as the world’s largest anoxic brackish water body. The development to more saline conditions implicated the disappearance of limnic organisms, such as ostracods and limnic dinoflagellates (Bahr et al., 2008; Shumilovskikh et al., 2012). Nevertheless, productivity increased considerably due to the invasion of marine organisms like the coccolithophoride Emiliania huxleyi, which is responsible for the well-known coccolith ooze layers introducing Unit I (Bukry, 1970; Degens and Ross, 1972; van der Meer et al., 2008; Coolen et al., 2009) since about 2.76 ka BP (Lamy et al., 2006).

17 INTRODUCTION

3. Motivation and objectives

The climatic and environmental history of the Black Sea is well documented for the last 40 ka and demonstrates high sensitivity of the basin in response to northern hemisphere climate dynamics. When considering the colder Saalian glacial as well as the warmer and wetter Eemian interglacial (Fig. 2; e.g. Seidenkrantz, 1993, Martrat et al., 2007; NEEM members, 2013), it is of particular interest to learn how the Black Sea responded to the changing climate during the penultimate glacial-interglacial transition (Termination II). Furthermore, during a 10-20 m higher Eemian sea level (Fig. 2 G; e.g. Siddall et al., 2006; Grant et al., 2012) shelf areas were larger (Fig. 9) and temperatures higher most likely leading to considerable ecological and biogeochemical differences compared to the Holocene.

Fig. 9: Maps showing the spatial extent of the Black Sea (“Lake”) for the present, the last glacial (sea level -110 m; Constantinescu et al., 2015), and two Eemian situations (smoothed sea level +10 m and maximum sea level +20 m; Grant et al., 2012). Maps were created with the GMT mapping tool from http://sfb574.geomar.de/gmt-maps.html with the option to set sea level to various depths.

18 INTRODUCTION

During R/V “Meteor” cruise M72/5 in 2007, gravity cores were obtained on the Archangelsky Ridge in the SE Black Sea, which document the Black Sea history back to about 134 ka BP. With this perspective, the following two major objectives are addressed:

1. Unravelling the climatic and environmental response of the Black Sea (“Lake”) to the penultimate deglaciation and last interglacial warming with special focus on:

 the magnitude and temporal development of the deglacial warming  meltwater events and their pathways as consequence of retreating continental ice sheets  associated changes in the basins hydrology, salinity, and productivity  the expression of the penultimate Mediterranean-Black Sea reconnection  the redox evolution during the Eemian sapropel formation

The results derived from these foci will be further used to compare, for the first time, the Black Sea evolution during the penultimate and last glacial-interglacial transitions. When considering present-day global warming and sea-level rise, knowledge about the environmental changes during Eemian warming may also open a window into the future evolution of the Black Sea.

2. Investigation of the Black Sea climate response to Dansgaard-Oeschger cycles and Heinrich events during Marine Isotope Stage 3. This objective involves:

 Identification of potential (multi)millennial-scale climate fluctuations associated with Greenland Dansgaard-Oeschger cycles (DO) and North Atlantic Heinrich Events (HE), and if so:  Estimation of temperature amplitudes during abrupt climate changes  Evaluation of the spatial temperature patterns in the North Atlantic-Eurasian region for estimation of the far-field impact of North Atlantic climate dynamics and inland propagation of temperature variability by atmospheric teleconnection  Assessment of the impact of DO and HE and associated temperature, rainfall, and wind patterns on the aquatic and adjacent terrestrial environment in the SE Black Sea  Long-term trends in the observed climatic and environmental changes across the Marine Isotope Stages 2, 3, and 4

19 INTRODUCTION

With these objectives, the present study aims to unravel the climatic and environmental history and the response of regional and global/northern hemisphere climate variability during lacustrine and marine phases of the Black Sea. The new results obtained in this study may improve our process understanding of possible long-term and abrupt climate transitions and potential “tipping points” in the climate system that may occur in the future.

4. Study area - the ‘Archangelsky Ridge’

This study bases on three gravity cores (22GC-3, 22GC-8, and 25GC-1) and surface sediment samples from 22MUC-1 retrieved during R/V “Meteor” cruise M72/5 (2007) on the Archangelsky Ridge in the south-eastern Black Sea in water depths at 418 m (station 25) and around 847 m (station 22; Fig. 10, Table 1). The Archangelsky Ridge represents an uplifted fault block in the south-eastern continental slope region within a tectonically still active region. This so-called horst block is surrounded by several extensional faults (Meredith and Egan, 2002).

The thickness of Quaternary sediments can reach up to 2,000 m (Fig. 10). Due to relatively low sedimentation rates in the SE Black Sea, deposits covering the penultimate glacial-interglacial transition occur in depth of only ~10 m below the seafloor, whereas the penultimate glacial deposits in the western Black Sea are found in depths >100 m (e.g. Quan et al., 2013).

The Archangelsky Ridge separates together with the Andrusov Ridge the eastern basin from the western Black Sea (Zonenshain and Pichon, 1986; Shillington et al., 2008; Nikishin et al., 2015). The Black Sea itself represents a remnant of a Mesozoic marginal sea, the Para- Tethys, which once developed as a back-arc basin (Zonenshain and Pichon, 1986).

20 INTRODUCTION

Fig. 10: Study area showing bathymetry and seismic profiles of the Archangelsky Ridge in the SE Black Sea and locations of cores 22GC-3, 22GC-8, and 25GC-1. The map was created with the GMT mapping tool from http://sfb574.geomar.de/gmt-maps.html. Seismic profiles 1.-4. generated during R/V “Meteor” cruise M72/5 in 2007 show the morphology across the ridge and were provided by Helge W. Arz. SSW-NNE section (5.) across the eastern Black Sea shows the Archangelsky Ridge under influence of compressional and extensional tectonics as well as the thickness of Quaternary sediments (slightly modified from Meredith and Egan, 2002).

5. Materials and methods

5.1 Sediment cores and stratigraphy

The present work bases on three gravity cores covering different periods of the Black Sea (“Lake”) history (Fig. 10; Table 1). The upper ~60 cm of core 22GC-3 encompass the last 12,000 years and therefore the late part of the last glacial-interglacial transition (Termination I (18-11 ka BP; MIS 2/1). This core was used for comparison between the penultimate and last

21 INTRODUCTION terminations at the same site and water depth. The core contains the stratigraphic Units I-III (Fig. 11; Degens and Ross, 1972) and comprises of limnic clays (Unit III), the Holocene sapropel (Unit II), and the coccolith ooze (Unit I). The Red Layers frequently observed in Unit III in cores from the western Black Sea (chapter 2) are also slightly observable in core 22GC-3. Additionally to this core, the uppermost surface sediment sample (0-1 cm) of the multicorer 22MUC-1 was analysed. The age model of core 22GC-3 relies on preliminary studies and bases on lithological correlation with the radiocarbon and Santorini ash dated core GeoB 7622-2 from the SW Black Sea for Units I-II (Lamy et al., 2006; Shumilovskikh et al., 2012). The ages for the glacial part (Unit III) were imported from the radiocarbon dated core 24GC-3 from the SE Black Sea and correlated with core 22GC-3 by means of high-resolution magnetic susceptibility and XRF Ca intensity (Shumilovskikh et al., 2012). Reservoir age correction was adopted from Kwiecien et al. (2008).

Table 1: Overview of the three gravity cores analysed in the present study.

core latitude longitude water core length age*2 [ka BP] time interval*2 depth [m] [cm] 22GC-3 42°13.53'N 36°29.55'E 838 839 < 12 Termination I (< 60)*1 (MIS 2/1) 25GC-1 42°06.2’N 36°37.4’E 418 952 20 - 64 MIS 2 - MIS 4

(307-952)*1

22GC-8 42°13.53'N 36°29.59'E 847 945 122 - 134 Termination II (785–945)*1 (MIS 6/5e) *1 corresponds to core section analysed in this study; *2 corresponds to periods investigated in this study

The lowermost 645 cm of core 25GC-1 were examined for the last glacial period between 20 and 64 ka BP. This sediment core comprises of limnic clays and intercalated authigenic carbonates (Fig. 11). The age model of this core (Nowaczyk et al., 2012) bases, for the investigated period, on eight AMS (accelerator mass spectrometry) radiocarbon datings on juvenile shells of the bivalve Dreissena rostriformis from the parallel core 24GC-3 (208 m water depth), the Campanian Ignimbrite tephra (Y5; 39.28 ± 0.11 ka; De Vivo et al., 2001), and the Laschamp geomagnetic excursion (40.70 ± 0.95 ka BP; Singer et al., 2009). The related palaeomagnetic records, including ChRM (characteristic remanent magnetisation) inclination and declination as well as relative palaeointensity served for further improvement of the chronology (Nowaczyk et al., 2012). Finally, coastal ice-rafted detritus (IRDC), carbonate content, Ca-XRF (X-ray fluorescence) counts, and high-resolution magnetic

22 INTRODUCTION susceptibility of the core were tuned to the oxygen isotope record from the North Greenland Ice core Project on the GICC05 timescale (NGRIP; Svensson et al., 2006, 2008). Tuning was done with the xtc-program (extended tool for correlation) allowing an interactive wiggle- matching routine (Nowaczyk et al., 2012).

Fig. 11: Age depth-models and stratigraphy of the gravity cores studied in the present work based on the stratigraphy presented in Nowaczyk et al. (2012), Shumilovskikh et al. (2012, 2013a, 2013b) and Wegwerth et al. (2014).

23 INTRODUCTION

To investigate the penultimate glacial-interglacial transition (Termination II), the lowermost 160 cm of the sediment core 22GC-8 were used. Similar to the cores covering Unit I-III, this section also comprise of limnic clays, the incomplete Eemian sapropel, and coccolith ooze (Fig. 11). The age model for the glacial part of core 22GC-8 bases on the 18 230 18 correlation between the δ Oostracods of core 22GC-8 and the Th dated δ O speleothem record of the Northern Anatolian Sofular Cave (Badertscher et al., 2011; Shumilovskikh et al., 2013b). For the Eemian part of core 22GC-8, a previous age model carried out on core 22GC- 3 is used (Shumilovskikh et al., 2013a), which bases on the correlation of pollen records from this core to well-dated pollen records from Ioannina 284 (Tzedakis et al., 2003) and Monticchio Lake (Allen and Huntley, 2009). Ages from core 22GC-3 were transferred to 22GC-8 by correlating XRF Ca records of both cores. Even though the Eemian sequence is incomplete due to a hiatus between ca. 65 and 120 ka BP in both cores (Fig. 11), the coverage of about 8,000 years of Eemian sapropel formation justifies a comparison to the Holocene.

5.2 Inorganic, organic, and ostracod geochemistry

A detailed overview about the multi-proxy approach applied in the present study is provided in Fig. 12. Additional parameters used in the present study are shown in Table 2.

Total carbon (TC) was determined by using elemental analyser EA 1110 (CE- instruments). Total inorganic carbon (TIC) was analysed after carbonate dissolution with 2M

HClO4 using elemental analyser multi EA 2000 (Analytik Jena). Total organic carbon (TOC) was calculated from the difference between TC and TIC.

The inorganic geochemical analyses involved total acid digestions of the freeze-dried and homogenised bulk sediment samples with a mixture of HNO3, HF, and HClO4

(Schnetger, 1997) and final dilution after acid evaporation with 2 vol.% HNO3. Major (e.g. Al, Ca, Fe, K, Ti) and trace (e.g. Sr, Zr) elements were measured by ICP-OES (iCAP 6300 Duo, Thermo Fisher Scientific). The trace elements Mo, Re, and W were measured by ICP- MS (iCAP Q, Thermo Fisher Scientific).

For organic geochemical analyses the bulk sediment samples were separated by column chromatography into three main fractions (n-alkanes, alkenones, and glycerol dialkyl glycerol tetra-ethers [GDGTs]) after accelerated solvent extraction (Thermo Scientific™ Dionex™ ASE™ 350). The fractions containing n-alkanes and alkenones were measured on a multichannel TraceUltra gas chromatograph (Thermo Scientific). The alkenone-based

24 INTRODUCTION

K’ palaeotemperatures were calculated according to a calibration of the U 37 unsaturation index (Prahl and Wakeham, 1987) of surface sediment (Conte et al., 2006). The fraction comprising GDGTs were measured by high-performance liquid chromatography/atmospheric pressure chemical ionization mass spectrometry (HPLC-APCI-MS; HPLCMS1100 Series, Hewlett Packard, USA) during two stays (one month each) in the Laboratory of Prof. Dr. E. Bard at CEREGE (Centre Européen de Recherche et d’Enseignement des Géosciences de l’Environnement, Géochimie et géochronologie, Aix-Marseille University). The TEX86-and BIT-indices were calculated according to Schouten et al. (2002) and Hopmans et al. (2004).

TEX86-based palaeotemperatures were calculated by using the global lake calibration by Powers et al. (2010) and the global marine subsurface (0–200 m) calibration by Kim et al. (2008).

After counting the ostracod valves of the species Candona spp. from the wet-sieved sediment samples, the cleaned valves of adult individuals (>500 µm in length) were dissolved in 2 mL 2vol.% HNO3 of suprapure quality for analysis of Ca, Mg, and Sr by ICP-OES (iCAP 6300Duo, Thermo) and ICP-MS (iCAP Q, Thermo Fisher Scientific). The concentration of U was determined on selected aliquots by HR-ICP-MS (Element II, Thermo) during a stay in the Microbiogeochemistry Group of Prof. Dr. H.-J. Brumsack at the Institute for Chemistry and Biology of the Marine Environment (ICBM, University of Oldenburg). For the determination of the Sr isotope composition, ostracods were cleaned with a few drops of H2O2 (5%), methanol, and deionised water. The 87Sr/86Sr ratios of the ostracods were measured by MC- ICP-MS (Neptune Plus, Thermo) at the University of Tübingen (Department of Earth Sciences, Isotope Geochemistry).

25 INTRODUCTION

Fig. 12: Schematic overview of the multi-proxy approach showing the main methods applied in the present work as well as possible interpretative information.

26 INTRODUCTION

Table 2: Overview about additional parameters from the investigated sediment cores used in the present study as well as interpretation and responsibilities.

Parameter/Methods Interpretation Institute and responsibility 18 δ Oostracods temperature and salinity GFZ Potsdam, B. Plessen and H. W. Arz

87 86 Sr/ Srostracods water input source University of Tübingen, I. C. Kleinhanns, M. Wille, A. Wegwerth

IRDC winter strength GFZ Potsdam, H. W. Arz TIC and TOC productivity, GFZ Potsdam, B. Plessen H. W. Arz, N.R. (25GC-1) preservation Nowaczyk Climate models temperature inland PIK Potsdam, A. Ganopolski propagation Isorenieratane photic zone euxinia ICBM Oldenburg, J. Köster, S. Eckert δ98Mo redox conditions Leibniz University Hannover, S. Weyer, S. Eckert, A. Brüske δ56Fe redoxcline migration Institute of Marine and Coastal Sciences, Rutgers University, USA, S. Severmann Inorganic redox conditions ICBM Oldenburg, S. Eckert, B. Schnetger, H.-J. geochemistry Brumsack (22GC-7, 22MUC-2, 22GC-3/Eemian) XRF (25GC-1) terrestrial input GFZ Potsdam, MARUM, Bremen University, H. W. Arz, N. R. Nowaczyk Palynology (pollen, vegetation, temperature, University of Göttingen, L. S. Shumilovskikh dinoflagellates) rainfall, salinity, productivity

6. Manuscript outline and personal contribution

The main part of this thesis comprises three manuscripts published/submitted to international peer-reviewed journals and one manuscript in preparation. The abstracts of three additional co-author publications related to the present study are presented together with the personal contribution in the appendix.

The first manuscript “Meltwater events and the Mediterranean reconnection at the Saalian-Eemian transition in the Black Sea” by A. Wegwerth, O. Dellwig, J. Kaiser, G. Ménot, E. Bard, L. Shumilovskikh, B. Schnetger, I. C. Kleinhanns, M. Wille, and H. W. Arz; published in Earth and Planetary Science Letters (2014) (404, 124-135) presents the climatic and environmental history of the Black Sea during the penultimate glacial-interglacial transition between 134 ka BP and 122 ka BP. The main foci are the deglacial meltwater events, the Termination II and Eemian warming, the Mediterranean Black Sea reconnection,

27 INTRODUCTION as well as a brief comparison with the Termination I. Personal contribution involves a multiproxy-approach including biomarker analysis, element ratios on ostracods, inorganic sediment geochemistry, as well as the data interpretation and concept and writing of the manuscript.

The second manuscript entitled “Black Sea temperature response to glacial millennial- scale climate variability” by A. Wegwerth, A. Ganopolski, G. Ménot, J. Kaiser, O. Dellwig, E. Bard, F. Lamy, and H. W. Arz; published in Geophysical Research Letters (2015) (42, 8147-8154) demonstrates for the first time pronounced continental temperature variability between 64,000 and 20,000 years ago in the Black Sea “Lake” coinciding with the Greenland Dansgaard-Oeschger cycles. A proxy-model comparison served, by including other published temperature records from the North Atlantic - Mediterranean corridor, for estimation of the inland propagation modes of Dansgaard-Oeschger cycles and Heinrich events. Personal contribution includes organic geochemistry and reconstruction of palaeotemperatures, analysis of Mg/Ca ratios on ostracods, as well as interpretation of data, concept and writing of the manuscript.

The third manuscript “Northern hemisphere climate control on the environmental dynamics in the glacial Black Sea “Lake” by A. Wegwerth, J. Kaiser, O. Dellwig, L. S. Shumilovskikh, N. R. Nowaczyk, and H. W. Arz, currently under review in Quaternary Science Reviews, provides a detailed reconstruction of the dynamics in riverine and aeolian input, salinity, and productivity in the Black Sea “Lake” in response to the long-term and millennial-scale northern hemispheric climate changes from late MIS 4 to early MIS 2. The analyses of n-alkane composition, counts and Sr/Ca ratios on ostracods, sediment inorganic geochemistry (except for XRF) as well as interpretation of data, concept and writing of the manuscript were performed by the author of this thesis.

The fourth manuscript “Redox evolution during Eemian and Holocene sapropel formation in the Black Sea” by A. Wegwerth, S. Eckert, O. Dellwig, J. Köster, B. Schnetger, S. Severmann, S. Weyer, A. Brüske, J. Kaiser, H. W. Arz, and H.-J. Brumsack; in preparation, revisits the Eemian and presents the first geochemical comparison of the redox evolution during the Holocene and Eemian sapropel formation in the Black Sea. Personal contribution includes the analyses of palaeotemperatures, ostracods geochemistry, major and redox-sensitive trace elements, TIC, and TOC in cores 22GC-8, 22MUC-1, and the Holocene part of 22GC-3, balance calculations, adaption of pre-existing age models to additional sediment cores as well as interpretation of data, concept and writing of the manuscript. S.

28 INTRODUCTION

Eckert contributed inorganic geochemical data of additional cores (Table 2). Analysis of isorenieratene derivatives were perfomed by J. Köster and S. Eckert. Stefan Weyer and A. Brüske measured Mo-Isotopes, and S. Servermann is responsible for Fe-Isotopes. All co- authors were involved in interpretation and discussion of the manuscript.

The data used in the manuscripts are stored online at the Data Publisher for Earth and Environmental Science PANGAEA (http://www.pangaea.de). Data already published in peer- reviewed journals are freely accessible and all other data will be unlocked after publication or are available upon request.

29 MANUSCRIPTS

MANUSCRIPTS

30 Meltwater events and the Mediterranean reconnection at the Saalian-Eemian transition in the Black Sea

Meltwater events and the Mediterranean reconnection at the Saalian- Eemian transition in the Black Sea

Antje Wegwertha*, Olaf Dellwiga, Jérôme Kaisera, Guillemette Ménotb, Edouard Bardb, Lyudmila Shumilovskikhc, Bernhard Schnetgerd, Ilka C. Kleinhannse, Martin Willee, and Helge W. Arza

a Leibniz-Institute for Baltic Sea Research Warnemünde (IOW), Marine Geology, Rostock, Germany b CEREGE, Aix-Marseille University, Collège de France, CNRS, IRD, Aix en Provence, France c Department of Palynology and Climate Dynamics, University of Göttingen, Germany d Institute for Chemistry and Biology of the Marine Environment (ICBM), University of Oldenburg, Germany e Isotope Geochemistry, Department of Earth Sciences, University of Tübingen, Germany

31 Meltwater events and the Mediterranean reconnection at the Saalian-Eemian transition in the Black Sea

Abstract

The last glacial-interglacial transition or Termination I (T I) is well documented in the Black Sea, whereas little is known about climate and environmental dynamics during the penultimate Termination (T II). Here we present a multi-proxy study based on a sediment core from the SE Black Sea covering the penultimate glacial and almost the entire Eemian interglacial (133.5 ±0.7-122.5 ±1.7 ka BP). Proxies comprise ice-rafted debris (IRD), O and Sr isotopes as well as Sr/Ca, Mg/Ca, and U/Ca ratios of benthic ostracods, organic and K’ inorganic sediment geochemistry, as well as TEX86 and U 37 derived water temperatures. The ending penultimate glacial (MIS 6, 133.5 to 129.9 ±0.7 ka BP) is characterised by mean annual lake surface temperatures of about 9°C as estimated from the TEX86 palaeothermometer. This period is impacted by two Black Sea melt water pulses (BSWP-II-1 and 2) as indicated by very low Sr/Caostracods but high sedimentary K/Al values. Anomalously 87 86 high radiogenic Sr/ Srostracod values (max. 0.70945) during BSWP-II-2 suggest a potential Himalayan source communicated via the Caspian Sea. The T II warming started at 129.9 ± 18 0.7 ka BP, witnessed by abrupt disappearance of IRD, increasing δ Oostracod values, and a first

TEX86 derived temperature rise of about 2.5°C. A second, abrupt warming step to ca. 15.5°C as the prelude of the Eemian warm period is documented at 128.3 ka BP. The Mediterranean- Black Sea reconnection most likely occurred at 128.1 ±0.7 ka BP as demonstrated by increasing Sr/Caostracods and U/Caostracods values. The disappearance of ostracods and TOC contents >2% document the onset of Eemian sapropel formation at 127.6 ka BP. During K’ sapropel formation, TEX86 temperatures dropped and stabilised at around 9°C, while U 37 temperatures remain on average 17°C. This difference is possibly caused by a habitat shift of Thaumarchaeota communities from surface towards nutrient-rich deeper and colder waters located above the gradually establishing halo-and redoxcline.

Keywords: Black Sea, penultimate glacial, Termination II, Eemian sapropel, meltwater pulses, Mediterranean reconnection

32 Meltwater events and the Mediterranean reconnection at the Saalian-Eemian transition in the Black Sea

1. Introduction

The Black Sea currently represents the Earth’s largest anoxic brackish water body (e.g. Demaison and Moore, 1980; Özsoy and Ünlüata, 1997). However, changes between limnic and brackish/marine conditions occurred during the past glacial cycles (Schrader, 1979). Anoxic conditions are favoured by a pycnocline separating high-salinity bottom waters derived from Mediterranean inflow via the Bosporus sill and low-salinity surface waters from excess river discharge and precipitation (Özsoy and Ünlüata, 1997; Kaminski et al., 2002). Initiated by global and regional climate change, sea-level fluctuations allowed a connection to the Mediterranean Sea for at least 12 times since 670 ka BP (Badertscher et al., 2011). While glacial periods are characterised by low global sea level and limnic conditions due to the disconnection from the Mediterranean Sea, increasing global temperature and sea level during glacial-interglacial transitions, the so-called Terminations, resulted in Mediterranean-Black Sea reconnections and the establishment of brackish/marine and anoxic conditions.

The last glacial-interglacial Termination I (T I), is one of the most investigated periods of the Black Sea, however dominated by studies in the western Black Sea (e.g. Ryan et al., 1997; Major et al., 2006; Bahr et al., 2008; Kwiecien et al., 2009; Soulet et al., 2011). Recent exceptions are Shumilovskikh et al. (2012) and Eckert et al. (2013) presenting palynological and geochemical data from the SE Black Sea, respectively. Although Zubakov (1988) provided an excellent basis for a climatostratigraphic correlation of the Black Sea record to the global oxygen-isotope scale for the last 1Ma, allowing determination of several saline and freshwater events coinciding with glacial-interglacial cycles, detailed insights into high- resolution dynamics of older Terminations, like TII, are sparse due to the lack of well constrained and complete sediment cores (e.g. Ross, 1978; Quan et al., 2013). By contrast, TII and Eemian studies worldwide established innovative concepts for a regional and global point of view during the last decades. For example, a possible Younger Dryas-like event shortly after the beginning of the Eemian interglacial (Sarnthein and Tiedemann, 1990; Sanchez Goñi et al., 1999; Bianchi and Gersonde, 2002; Cannariato and Kennett, 2005; Risebrobakken et al., 2006) and/or a Termination pause during TII (Lototskaya and Ganssen, 1999; Risebrobakken et al., 2006) are still under debate. Several records disagree with such a Younger-Dryas like event (Carlson, 2008). Pollen records from Lake Urmia (Iran; Stevens et al., 2012) and the Black Sea (Shumilovskikh et al., 2013b) even argue against a Termination break. Furthermore, it seems that warming and sea-level rise during T II occurred earlier and slightly more rapid than during T I as inferred mainly from foraminiferal assemblages and

33 Meltwater events and the Mediterranean reconnection at the Saalian-Eemian transition in the Black Sea oxygen isotopes (e.g. Seidenkrantz, 1993; Ruddiman, 2006; Siddall et al., 2006; Thomas et al., 2009).

During the R/V “Meteor” cruise M72/5 in 2007, a unique sediment core (22GC-8) was retrieved from the SE Black Sea covering the sedimentary record from 133.5 to 122.5 ka BP thus comprising the late MIS 6 (Saalian) and MIS 5e (Eemian). Based on a multi-proxy approach at high resolution, we aim to provide new insights into the penultimate glacial- interglacial dynamics in the Black Sea region by using sediment and ostracod (isotope) geochemistry as well as biomarkers.

2. Environmental setting, hydrology, and climate

The Black Sea is a semi-enclosed epicontinental basin located in SE Europe/SW Asia (Figs. 1A, 1B) with a maximum water depth of about 2200 m. Apart from precipitation, freshwater input occurs mainly via rivers Danube, Don, Dnieper, and Dniester in the north and Sakarya, Kizilirmak, Yesilirmak, and Rioni in the south (Fig. 1A). The present connection to the Mediterranean Sea through the Bosporus strait (-35 m deep) allows inflow of saline waters, while simultaneously an outflow of brackish waters occurs in the surface layer (Özsoy and Ünlüata, 1997; Kaminski et al., 2002).

The resulting halocline which separates salty bottom waters (∼22.4 g kg-1) from low salinity surface waters (∼18 g kg-1) (Spencer et al., 1972; Calvert et al., 1987) prevents ventilation and oxygenation of the deeper waters and favours the formation of a pelagic redoxcline at about 100–150m water depth. Initiated by intense microbial activity, about 87% of the entire water volume is currently characterised by euxinic conditions (Özsoy and Ünlüata, 1997) finally leading to sapropel formation (e.g. Brumsack, 2006 and references therein).

The southern Black Sea region is influenced by Mediterranean climate causing hot/dry summers and cold/wet winters with cyclonic Mediterranean moisture sources (Cullen and deMenocal, 2000; Lamy et al., 2006; Kwiecien et al., 2009; Göktürk et al., 2011). However, the Black Sea itself forms a considerable humidity source in the region (Fleitmann et al., 2009; Badertscher et al., 2011; Göktürk et al., 2011). The NE Anatolian and E Black Sea regions are the most rain-laden areas of Turkey in summer due to orographic precipitation at the North Anatolian Mountains (Türkeş et al., 2009; Deniz et al., 2011). Compared to the

34 Meltwater events and the Mediterranean reconnection at the Saalian-Eemian transition in the Black Sea western part of the southern Black Sea coast, less extreme temperature fluctuations with mild winters and warm summers characterise the eastern part (Deniz et al., 2011).

Fig. 1. (A) Detailed map showing the Black Sea with the core site 22GC-8 at the Archangelsky Ridge and the Black Sea sites discussed in the text. (B) Map showing the region between the Mediterranean Sea and the Himalayas and the sites discussed in the text: 1. 22GC-8 this study; 2. MD042790 Black Sea (Ménot and Bard, 2012); 3.BLKS9810-NW Black Sea (Major et al., 2006); 4.GeoB7608-1 Black Sea (Bahr et al., 2008); 5.So6-3-Sofular Cave speleothem (Fleitmann et al., 2009;Badertscher et al., 2011); 6.Lake Ohrid, Albania (Lézine et al., 2010;Sulpizio et al., 2010;Vogel et al., 2010); 7.Lago Grande di Monticchio, Italy (Brauer et al., 2007;Allen and Huntley, 2009); 8.Pantelleria, Italy (Paterne et al., 2008); and 9.ODP-977A-Alboran Sea, Iberian Margin (Martrat et al., 2004; 2007). The arrows denote the coupled meltwater pathway of Fennoscandian (yellow) and potential Himalayan (red) origin.

35 Meltwater events and the Mediterranean reconnection at the Saalian-Eemian transition in the Black Sea

3. Material and methods

This study bases on the lower section of gravity core 22GC-8 (945cm total length) retrieved from the tectonically uplifted fault block Archangelsky Ridge (Meredith and Egan, 2002) in the SE Black Sea (42°13.5’N, 36°29.6’E) during the R/V Meteor cruise M72/5 in 2007 in a water depth of 847 m (Fig. 1A). The sediment core contains a well-preserved sequence (785–945 cm core depth) from the Saalian glacial (MIS 6, limnic clay) to the Eemian interglacial (MIS 5e, sapropel).

3.1. Age-depth model

18 The strong correspondence between δ Ospeleothem from Sofular Cave located in NW 18 Anatolia and δ Oostracods from the Black Sea for T I suggests that the Sofular isotope signature predominantly reflects changes in δ18O of precipitation arriving from the Black Sea (water vapour source effect; Badertscher et al., 2011). Following the same line of argument our age- 18 depth model for TII is derived from the correlation (four tie points) between the δ Oostracods 230 18 and the Th dated δ OSofular record (Figs. 1, 2; Badertscher et al., 2011; Shumilovskikh et al., 2013b).

A 1 cm thick tephra layer centred at 855.5 cm provides a further time marker. This tephra, identified by XRD analysis, is characterised by a high sanidine content of about 45%, a mineral typically enriched in volcanic material. Based on the preliminary age model, it most likely originates from the Mediterranean Pantelleria eruption (P-11; southern Italy; Fig. 1B) at 131 ±5 ka BP (Paterne et al., 2008). The widespread P-11 tephra was also observed in central Greece (Karkanas et al., in press), SE Aegean Sea (Grant et al., 2012), Lake Ohrid (Fig. 1B, Albania; Lézine et al., 2010; Sulpizio et al., 2010) and in the Ionian Sea (Paterne et al., 2008).

Finally, for the Eemian part, the pollen record of core 22GC-3 (Shumilovskikh et al.2013a, 2013b) has been correlated (two tie points corresponding to arboreal pollen maxima) with pollen records from Ioannina 284 (Tzedakis et al., 2003) and Monticchio Lake (Allen and Huntley, 2009; Figs. 1B, 2). Ages from 22GC-3 were transferred to 22GC-8 based on a correlation (seven tie points) of XRF Ca. For more details about the age model the authors refer to Shumilovskikh et al. (2013a, 2013b).

Mean squared estimates (MSEi) of age uncertainty were determined after Grant et al. (2012) and are ± 0.7 ka for the glacial part and ± 1.7 ka for the interglacial part.

36 Meltwater events and the Mediterranean reconnection at the Saalian-Eemian transition in the Black Sea

Fig. 2. Age-depth relation by means of stratigraphic correlation of (1) pollen from the sediment core (Shumilovskikh et al., 2013b) with pollen records from Ioannina 284 (Tzedakis 18 et al., 2003) and Monticchio Lake (Allen and Huntley, 2009), and (2) of δ Oostracods with a 230Th dated δ18O speleothem record from North Anatolian Sofular Cave (Badertscher et al., 2011; Shumilovskikh et al., 2013b). A further time marker forms the tephra layer at about 130.9 ± 0.7 ka BP originating from the Mediterranean Pantelleria eruption (P-11; southern Italy) originally dated at 131 ± 5 ka BP (Paterne et al., 2008). Age uncertainties were determined by the mean squared estimate (MSEi) from Grant et al. (2012, supplement) by considering four sensitive variables (radiometric dating error of Sofular Cave speleothem, 18 18 sample spacing in the δ Ospeleothem record, sample spacing in the 22GC-8 δ Oostracods record, and extra uncertainty allowance for more ambiguous tie-points). Sedimentation rates are given for each interval. Modified from Shumilovskikh et al. (2013a, 2013b).

3.2. Sediment geochemistry

Sedimentary Al, Ca, K, and Sr contents were measured from HF/HClO4 acid digestions (Schnetger, 1997) by ICP-OES (iCAP 6300Duo, Thermo). Precision and accuracy were checked by using the international reference material SGR-1 (USGS) and were better than 3.7% and 6.9%, respectively. Total carbon (TC) was determined by using elemental analyser EA 1110 (CE-instruments). Total inorganic carbon (TIC) was analysed after carbonate dissolution with 2M HClO4 using elemental analyser multi EA®2000 (Analytik Jena).

37 Meltwater events and the Mediterranean reconnection at the Saalian-Eemian transition in the Black Sea

Precision and accuracy were better than 2.3% (TC) and 1.3% (TIC). Total organic carbon (TOC) was calculated as the difference between TC and TIC.

3.3. Ostracod geochemistry and isotopic composition

For geochemical composition, intact valves of adult Candona spp. were hand-picked from wet-sieved sediments (>150μm, 2 cm resolution) and cleaned under a binocular microscope with a thin brush and a few drops of deionised water. 4 to 12 valves were dissolved in 2 mL 2 vol.% HNO3 of suprapure quality for analysis of Ca, Mg, and Sr by ICP- OES (iCAP 6300Duo, Thermo) using external calibration. Mg and Sr concentrations in calibration solutions were corrected for contamination originating from the Ca standard solution. Signal depressions caused by Ca were compensated using In as an internal standard. Possible contamination by detrital material was monitored by Al. Precision and accuracy were checked by an external solution measured every five samples and ranged from 1.0-1.3% and 0.2-1.1%, respectively.

Selected aliquots were investigated for U by HR-ICP-MS (Element II, Thermo) using external calibration and Tl as internal standard. Sample intensities ranged between about 10,000 cps and 160,000 cps U corresponding to U concentrations in the sample solutions of 0.02 μg L-1and 0.32 μg L-1. Precision and accuracy were checked by an external solution and were better than 2.1% and 1.9%, respectively.

For Sr isotope analyses, ostracods were cleaned with a few drops of H2O2 (5%), methanol, and deionised water (Janz and Vennemann, 2005; Major et al., 2006; Vasiliev et al., 2010). The ostracods were dissolved in 300 μL of 2% HNO3 and dried on a hot plate at 80°C. Separation of Sr from the sample matrix was achieved in microcolumns filled with

Eichrom© Sr-spec resin using the HNO3-H2O technique. Dried Sr separates were re-dissolved in 0.3 N HNO3 and fully automated isotope ratio measurements were performed in static mode on a MC-ICP-MS (Neptune Plus, Thermo). Instrumental mass bias was corrected using a 88Sr/86Sr ratio of 8.375209 and exponential law. External reproducibility for NBS SRM 987 (n =14) is 0.710273 ± 14 (2 s.d.) for the 87Sr/86Sr ratio. Total procedural blank was <228 pg for Sr.

38 Meltwater events and the Mediterranean reconnection at the Saalian-Eemian transition in the Black Sea

3.4. Organic geochemistry

Total lipids of homogenised samples (1-6 g sediment) were extracted by accelerated solvent extraction (Thermo Scientific™ Dionex™ ASE™ 350) with a dichloromethane and methanol mixture (DCM/MeOH 9:1). The extracts were further desulphurized by activated copper pieces.

The extracts were separated over Pasteur pipettes plugged with activated Al2O3 into neutral (hexane/DCM 1:1) and polar (DCM/MeOH 1:1) fractions. After adding hexatriacontane as injection/internal standard the neutral (containing alkenones) fraction was measured on a multichannel TraceUltra GC (Thermo Fisher Scientific). Peak identification was based on comparing peak retention times with a sediment standard containing alkenones. K’ Alkenone-based palaeotemperatures were calculated according to a calibration of U 37 unsaturation index (Prahl and Wakeham, 1987) of surface sediment (Conte et al., 2006).

The polar fractions (containing glycerol dialkyl glycerol tetraethers, GDGTs) were filtered through a 0.45 μm PTFE filter after the addition of a C:46 GDGT standard. The measurements were carried out by high-performance liquid chromatography/atmospheric pressure chemical ionisation mass spectrometry (HPLC-APCI-MS; HPLCMS1100 Series, Hewlett Packard, USA) using a Prevail Cyano column. Peak identification was performed by single ion monitoring (see for details Ménot and Bard, 2012). The calculations of TEX86- and BIT-indices were done according to Schouten et al. (2002) and Hopmans et al. (2004). Water temperatures were calculated using the global lake calibration by Powers et al. (2010) (943.5- 801.5 cm) and the global marine subsurface (0-200 m) calibration by Kim et al. (2008) (800.5-792.5 cm). The mean standard deviation for duplicate runs is 0.003 for BIT and 0.004 for TEX86 with the latter corresponding to 0.2°C.

4. Results

4.1. Sediment geochemistry

The contents of K reflecting detrital clay minerals (e.g. illite) and K-feldspars were normalised to Al to eliminate dilution effects by carbonate and organic matter (Fig. 4D). The ending glacial shows two intervals (133.5–132.7 and 132.3–129.9 ka BP) with strongly elevated but also fluctuating K/Al values. Afterwards, K/Al remains on a uniform low level of about 0.26 corresponding to the typical lithogenic background of the Black Sea (Piper and

39 Meltwater events and the Mediterranean reconnection at the Saalian-Eemian transition in the Black Sea

Calvert, 2011). Except for the very end of the penultimate glacial at about 130.8 ka BP, total inorganic carbon (TIC, Fig. 3D) remains roughly constant on a low level (1.5%) representing detrital carbonate as determined via microscopy. Thereafter, TIC increases with values ranging between 1.8% and 6.5% due to authigenic carbonate precipitation as identified via microscopy. Sr/Cabulk values, which help to differentiate between the carbonate types, remain low (35 ·104) during the penultimate glacial and Termination II (Fig. 3E). Thereafter, Sr/Ca increases and shows at least two distinct maxima (>100 ·104) occurring parallel to elevated TIC values indicating inorganic aragonite precipitation. Low Sr/Ca values are associated with two coccolith-dominated intervals. Total organic carbon (TOC) indicating productivity and/or preservation is comparatively low until 127.8 ka BP (about 0.3%) but starts to increase at 127.7 ka BP and remains on a high level (about 5%) within the Eemian sapropel (Fig. 3F).

4.2. Ostracod geochemistry

The benthic ostracod Candona spp. is widely distributed in the limnic/brackish sediments of the Black Sea during the last glacial (Fig. 1, Bahr et al.2006, 2008; Major et al., 2006; Boomer et al., 2010), thus forming contemporary witness of former climatic and environmental conditions. The isotopic and geochemical composition of ostracod shells may indicate changes in e.g. temperature, precipitation, salinity, nutrients, dissolved oxygen, and the catchment area (e.g. Chivas et al., 1986; De Deckker et al., 1999; Decrouy et al., 2011a, 2011b). Since Candona is a benthic organism and the water depths of the present core is about 800 m, the geochemical signatures of their shells are expected to reflect long-term conditions at the sediment-water interface rather than seasonal changes.

4.2.1. Element ratios

Mg/Caostracods from the glacial Black Sea “Lake” mainly depends on temperature and only in the case of strong changes in water chemistry on Mg/Cawater (Bahr et al., 2008). Generally, increasing temperature leads to enhanced incorporation of Mg into the ostracod -1 shells. In this study Mg/Caostracods values are almost constant at about 1.4 mmol mol until 131 ka BP (Fig. 3C). Thereafter they fluctuate around 3 mmol mol-1 to finally reach 6.9mmol mol- 1 at 127.6 ka BP. The lower value of the last sample containing ostracods is most likely not representative due to the extremely low number of individuals and possible diagenetic alteration, which might be also true for Sr/Caostracods and U/Caostracods. Primarily controlled by

40 Meltwater events and the Mediterranean reconnection at the Saalian-Eemian transition in the Black Sea

salinity (Wansard et al., 1998), Sr/Caostracods reveals two periods (133.5-132.5 and 131.5-130.5 ka BP) with minimum values of about 0.8 mmol mol-1 (Fig. 4C). After smoothly increasing values between 130.5 and 128.1 ka BP, Sr/Caostracods abruptly reaches a maximum of 1.8 mmol mol-1.

The U/Ca ratio of benthic carbonate shells (e.g. foraminifera, ostracods) may provide information about changes in salinity (Russell et al., 1994) and/or redox conditions (Ricketts et al., 2001). A continuously low level is seen until 128.3 ka BP (Fig. 4F) but strongly increasing values are determined coevally with the appearance of Mediterranean-type larval shells from Mytilus galloprovincialis (Aral, 1999; Lamy et al., 2006). At 126.5 ka BP a thin layer (about 1 cm) containing well preserved benthic foraminifera (Ammonia tepida; G. Schmiedl, personal communication) is seen.

4.2.2. Strontium isotopes

Since 87Sr/86Sr of ostracods is not affected e.g. by temperature and/or biogenic fractionation this proxy can help to identify the impact of different water sources (e.g. Major 87 86 et al., 2006). Sr/ Srostracods values range between 0.708945 and 0.709452 (Fig. 4E). After a 87 86 certain increase between 133.5-132.5 ka BP, Sr/ Srostracods reaches a maximum level between 131.5-130.5 ka BP strongly exceeding the last Black Sea glacial (maximum 0.7091; 87 86 Major et al., 2006). Afterwards, the Sr/ Srostracods values decrease during TII and increase again at around 128.1 ka BP.

4.3. Organic geochemistry

4.3.1. GDGT

GDGTs are widely distributed in aquatic and terrestrial environments and are produced by archaea and bacteria (for a review Schouten et al., 2013). For reconstruction of past water temperatures the TEX86-index (TetraEther IndeX of lipids with 86 carbon atoms) is used

(Schouten et al., 2002). In core 22GC-8, the suitability of the TEX86-derived temperatures is assured by BIT-indices being generally below 0.3 (Table S1; Hopmans et al., 2004; Weijers et al., 2006; Castañeda et al., 2010) and a typical Thaumarcheaota GDGT-0/crenarchaeol ratio below 2 (Table S1) excluding a potential temperature bias related to the input of soil-derived isoprenoid GDGTs and/or shifts in the archaean populations (Schouten et al. 2002, 2013;

41 Meltwater events and the Mediterranean reconnection at the Saalian-Eemian transition in the Black Sea

Turich et al., 2007; Blaga et al., 2009 and references therein). While TEX86-derived water temperatures fluctuate between 7.7-11.4°C during the ending penultimate glacial, Termination II warming shows a more gradual increasing trend (Fig. 3B). The most prominent feature forms the pronounced increase to about 21.5°C between 128.5 and 127.9 ka BP, followed by a decrease to ca. 15°C at 127.5 ka BP. As present-day Thaumarchaeota abundances are highest close to the redoxcline (ca. 100 m water depth; Lam et al., 2007; Wakeham et al.,

2007), the TEX86 calibration for subsurface (0–200 m) temperature reconstruction from Kim et al. (2008) has been used between 127.3 and 122.5 ka BP (Fig. 3B). During this interval

TEX86 temperature estimates have a constant level of about 9°C.

4.3.2. Long-chain alkenones

K’ The long-chain di-and tri-unsaturated C37 alkenones used forU 37-based temperature reconstructions are specific compounds, which are mainly produced by the coccolithophorids Emiliania huxleyii and Gephyrocapsa oceanic (Volkman et al., 1980; Marlowe et al., 1984; Coolen et al., 2009) and whose concentrations relative to each other are a function of growth temperature (Prahl and Wakeham, 1987).

The appearance of C37-39 alkenones at 127.6 ka BP marks the intrusion of marine haptophytes in the Black Sea (Fig. 3H). The concentrations show two further intervals reaching maximum values at 125.3 and 123 ka BP. The alkenone-based sea-surface K’ temperature estimates (SST U 37) are around 17°C and increase to ca. 20°C in the earliest part of the record (Fig. 3B). As 17°C is close to the modern mean SST value during early summer coccolithophorid blooming (Cokacar et al., 2004) and knowing that maximum alkenone concentration currently occurs in the upper 10m of the Black Sea (Freeman and K’ Wakeham, 1992), the U 37 proxy most likely reflects early summer surface temperature during the Eemian.

42 Meltwater events and the Mediterranean reconnection at the Saalian-Eemian transition in the Black Sea

Fig. 3. Temporal variation of temperature and productivity related proxies from the penultimate glacial towards the Eemian. (A) ice-rafted debris (IRD; Shumilovskikh et al., K’ 2013b), (B) water temperatures based on TEX86 and U 37, (C) Mg/Caostracods (data point in brackets extremely low abundance and diagenetically altered), (D) total inorganic carbon (TIC), (E) Sr/Cabulk, (F) total organic carbon (TOC), (G) percentages of freshwater and brackish dinocysts (modified after Shumilovskikh et al., 2013b), (H) the sum ∑C37-39 alkenones/TOC. The ▲denote tie points for the age model, ♦ the first appearance of Mediterranean-type mollusc species Mytilus galloprovincialis, and the ● a layer of benthic foraminifera.

43 Meltwater events and the Mediterranean reconnection at the Saalian-Eemian transition in the Black Sea

5. Discussion

During the transition from the penultimate glacial (MIS 6) into the Eemian (MIS 5e) the Black Sea shows a complex environmental response to several important factors including the decaying northern ice sheet, the hemisphere-wide climatic amelioration, and the Mediterranean Sea reconnection related to the eustatic sea-level rise. In the following, we concentrate on the three major periods: (i) the end of the penultimate glacial (133.5-129.9 ka BP), (ii) the warming during Termination II (129.9-128.1 ka BP), and (iii) the Mediterranean- Black Sea reconnection during the Eemian (128.1-122.5 ka BP).

5.1. Ending penultimate glacial (133.5–129.9 ka BP)

The ending penultimate glacial shows several indications for cold but also relatively unstable climate conditions in the Black Sea region. Thus, TEX86-derived water temperatures reveal a pronounced variability (Fig. 3B). Although BIT and GDGT-0/crenarchaeol values assure the suitability of the TEX86 palaeothermometer, this proxy has to be interpreted carefully as it is still under debate whether TEX86 reflects summer, winter, or annual mean temperatures. This uncertainty is also valid for the water depth from where the signals may originate. As reviewed recently by Castaneda and Schouten (2011), several studies on lacustrine systems suggest that TEX86 temperatures represent surface winter conditions due to maximum archaea production at that season. On the other hand, according to Ménot and Bard

(2012) the TEX86 temperatures from the last glacial likely represent summer conditions at the lake surface (upper 30 m). The comparatively high temperature level of 9°C for the ending MIS 6 glacial of this study argues against a temperature signal produced during winter. Frequent formation of winter coastal ice as indicated by IRD (Fig. 3A, Nowaczyk et al., 2012;

Shumilovskikh et al., 2013b) is not compatible with the comparatively high TEX86 temperature level, whereas the comparison of IRD and TEX86 temperatures is certainly limited by the lower resolution of the latter parameter. We postulate that TEX86 temperature estimates during MIS 6 reflect annual mean surface conditions because: (i) the global lake calibration of TEX86 was generated for mean annual surface temperatures (Powers et al., 2010) and (ii) the mean annual SST estimates from the Mediterranean Sea of about 11-12°C during glacials MIS 6, 4, and 2 (Cacho et al., 2001; Martrat et al., 2004; Fig. 1B) show strong similarity to TEX86 temperature estimates considering the continental climate of the Black

Sea region. Following this suggestion, we conclude, that our low TEX86-derived mean annual

44 Meltwater events and the Mediterranean reconnection at the Saalian-Eemian transition in the Black Sea lake surface temperatures (LST) along with fluctuating IRD presence during the ending penultimate glacial reflect relatively cold but also unstable conditions as a prelude for the following warming phase.

The surface temperature changes were however not strong enough to influence the deeper water column at least until 131 ka BP. Low Mg/Ca values of benthic ostracods (Fig. 3C) indicate relatively stable and cold conditions in the deeper Black Sea (800m) that likely were not influenced by seasonal changes. The observed Mg/Caostracods fluctuations of maximum 0.5 mmol mol-1, imply relative changes in temperature of less than 1°C (Bahr et al., 2008). A higher variability is seen only during the last 1000 years of the ending penultimate glacial probably reflecting the initial TII warming.

In accordance with the generally cold lake surface temperatures we observe permanently low TOC values (Fig. 3F) that may indicate limited primary productivity. The use of bulk TOC, comprising freshly produced aquatic as well as altered terrestrial organic matter, suffers from decomposition in the water column and sediments (e.g. Burdige, 2007; Zonneveld et al., 2010) thus asking for further proxies reflecting productivity. The lack of authigenic carbonates as an indirect proxy for phytoplankton activity (Bahr et al., 2008; Nowaczyk et al., 2012) also argues for a generally reduced productivity. Furthermore, low dinocyst percentages (Shumilovskikh et al., 2013b) suggest a limited summer productivity (Fig. 3G; De Vernal and Marret, 2007).

5.2. Meltwater pulses

As a prominent feature of the ending penultimate glacial, Badertscher et al. (2011) suggested a melt water pulse from the Fennoscandian ice sheet possibly entering the Black Sea via the Caspian Sea. The authors determined very light δ18O values in the Sofular speleothem in NW Anatolia between 133.3-130.5 ka BP (Figs. 1A, 4A) which are assumed to reflect the isotopic composition of the Black Sea surface waters and thus the input of a large quantity of freshwater to the Black Sea “Lake”.

45 Meltwater events and the Mediterranean reconnection at the Saalian-Eemian transition in the Black Sea

Fig. 4. Temporal variation of terrigenous input and salinity related proxies from the penultimate glacial towards the Eemian. (A) δ18O record of Sofular speleothems So-6 18 (Fleitmann et al., 2009; Badertscher et al., 2011), (B) δ Oostracods (Shumilovskikh et al., 87 86 2013b), (C) Sr/Caostracods, (D) K/Albulk, (E) Sr/ Srostracods, (F) U/Caostracods, (G) percentages of marine dinocysts (modified after Shumilovskikh et al., 2013a), and (H) the smoothed relative sea level with lower and upper limits (shadow) from Grant et al., 2012. Sample in brackets extremely low concentrations and diagenetically altered. The ▲ mark tie points for the age model, the ♦ first appearance of Mediterranean-type mollusc species Mytilus galloprovincialis, and the ● a layer of benthic foraminifera.

46 Meltwater events and the Mediterranean reconnection at the Saalian-Eemian transition in the Black Sea

This meltwater pulse is most likely also documented in the SE Black Sea record 18 because δ Oostracods values show a pronounced negative excursion between 132.1-130.2 ka BP 18 (Fig. 4B). As the δ Oostracods signal originate from benthic organisms in depths of about 800m such strong depletion could not be caused by usual precipitation and riverine input alone but must have its origin in an extensive meltwater contribution. This assumption is supported by the Sr/Caostracods pattern indicating a lowering of the residual salinity of the Black Sea “Lake”

(Fig. 4C) for this interval. Correspondingly low Sr/Caostracods values occur also between 133.5- 132.5 ka BP and point to two separate meltwater pulses in contrast to the idea of a continuous meltwater phase (Badertscher et al., 2011). These meltwater pulses are here referred to as BSWP-II-1 (133.5-132.5 ka BP) and BSWP-II-2 (131.5-130.5 ka BP) following the nomenclature of Soulet et al. (2013) on T I.

While the presence of “red layers” during Termination I imply a Dnieper drainage basin source for BSWP in the NW Black Sea (e.g. Major et al., 2006; Bahr et al., 2008; Soulet et al., 2013, Fig. 1A), such deposits are not found at our SE study site for TII. Nevertheless, during both BSWP-II, enhanced sedimentary K/Al indicates an input of detrital suspended matter differing significantly from the local background level of about 0.26 (Fig. 4D). According to Piper and Calvert (2011) elevated K/Al values are caused by the deposition of illite- and K-feldspar-rich material from the northern rivers (Danube, Dniester, Dnieper) and Azov Sea, which generally points towards a northern meltwater source.

87 86 The most remarkable feature of both BSWP-II are increasing Sr/ Srostracods values 87 86 (Fig. 4E). While the Sr/ Srostracods level of BSWP-II-1 is roughly similar to the T I melt water event (Major et al., 2006; Bahr et al., 2008), BSWP-II-2 reaches much more radiogenic values of up to 0.70945 distinctly exceeding the T I level (about 0.7091; Major et al., 2006). Close-by rock and sediment sources with radiogenic Sr might be the Anatolian/Pontic Mountains in the south and/or the Caucasus in the east (Fig. 1A). Unfortunately, available 87Sr/86Sr data are limited to specific rock types in both orogens but an average 87Sr/86Sr value of about 0.7054 calculated from studies in the Eastern Pontides most likely excludes an Anatolian/Pontic Mountain source (Altherr et al., 2008; Aydin et al., 2008; Kaygusuz and Aydınçakır, 2009). This is also true for the Caucasus showing an average 87Sr/86Sr value of about 0.7041 (Mengel et al., 1987; Lebedev et al., 2007).

While the freshwater most likely originated directly from the Fennoscandian Ice Sheet and from the Dnieper drainage basin (e.g. Major et al., 2006; Soulet et al., 2013) during the last glacial melt water pulse, northern rivers hardly represent potential sources for the BSWP-

47 Meltwater events and the Mediterranean reconnection at the Saalian-Eemian transition in the Black Sea

II-2 Sr-isotope anomaly as their signatures are less radiogenic (av. 0.708792; Major et al., 2006). Even though K/Al suggests illite-rich (Müller and Stoffers, 1974) northern sediments as the source of suspended matter, the waters eroding this material from the shelf regions may originated from a more distant source carrying a different isotopic signature.

Page et al. (2003) reported increasing Sr-isotope values in mollusc shells from the Caspian Sea during the last 100 ka reaching a maximum level at ∼100 ka (0.70842-0.70864). Because the Sr-isotope values of the rivers Volga (0.70802) and Ural (0.7082; Vasiliev et al., 2010) north of the Caspian Sea are too low, the authors suggested that the Amu Darya river, which currently flows into the Aral Sea, may have contributed highly radiogenic Sr from the Himalayas (Henderson et al., 1994 and references therein) to the Caspian Sea during periods of elevated discharge and enhanced Himalayan runoff. Meltwater with extremely high radiogenic Sr from lesser Himalayan streams (0.7166-1.023; Jacobson et al., 2002) could have been sufficient to account for measureable changes in Sr-isotope signatures. Therefore, we propose that water from the Amu Darya drainage basin fed by radiogenic Sr from enhanced western Himalayan weathering during massive deglacial melt water pulses, entered the Caspian Sea and finally reaching the Black Sea through e.g. the Manych depression (Fig. 1B). This assumption is supported by studies suggesting that the Amu Darya River directly entered the Caspian Sea at the beginning of the Holocene (Boomer et al., 2000) and that it was strongly influenced by the western Himalayas (Leroy et al., 2013). Assuming comparable dynamics during the penultimate glacial, the crucial prerequisite remains the Caspian-Black Seas connection, which may have occurred between 135 and 130 ka BP, as suggested from speleothem and faunal studies (Svitoch et al., 2000; Badertscher et al., 2011, Fig. 1B). This connection was likely strengthened by Fennoscandian melt water discharge from the Volga River as assumed for the last glacial.

Unfortunately, to the authors’ best knowledge, there is a lack of studies on Sr-isotopes in the Caspian and Aral Seas and other Eurasian lakes (and between the Black Sea and Himalayas) during this period, which would help to understand and reconstruct ancient meltwater pathways. Future studies in these regions could also bring more light into prominent shoreline terraces surrounding the Caspian Sea, which are up to now not well-dated but may be also partly related to MIS 6 meltwater events (for review see Forte and Cowgill, 2013). Unless further studies will provide evidence for a different source of the radiogenic Sr, our record likely demonstrates a far-field component of Asian climate dynamics on the Black

48 Meltwater events and the Mediterranean reconnection at the Saalian-Eemian transition in the Black Sea

Sea/Northern Anatolian region during the BSWP-II-2 at the end of the penultimate glacial period.

5.3. Termination II and early Eemian warming (129.9-128.1 ka BP)

Termination II marks the transition period between the penultimate late glacial and the Eemian interglacial lasting for about 1.4 ka (Fig. 2). The onset of TII is here defined after the last pronounced appearance of IRD at ca. 129.9 ka BP (Fig. 3A), which is in accordance with sedimentological studies from Lake Ohrid (Fig. 1B, Albania; Vogel et al., 2010) and pollen records from Lago Grande di Monticchio (Italy, Fig. 1B; Allen and Huntley, 2009). The disappearance of IRD document milder winters and a gradual increase of TEX86-derived LST rising mean annual temperatures from 9°C to about 15.5°C at 128.4 ka BP in the early Eemian comparable to present annual mean values of 15°C (Fig. 3B; Locarnini et al., 2010). Besides further Eemian warming, the following abrupt increase to about 21.5°C within only 0.4 ka has to be attributed to a shift in Thaumarchaeota productivity towards summer because an amplitude of almost 13°C between glacial and interglacial mean temperatures appears rather unlikely. Irrespective of aforementioned limitations caused by the used calibration, this assumption is also in accordance with modern summer water temperatures at the coring site of about 23°C (Locarnini et al., 2010). Increasing summer productivity during this period is further reflected by rising phytoplankton-induced carbonate precipitation (TIC; Fig. 3D) and increasing percentages of freshwater and brackish dinocysts (Fig. 3G; Shumilovskikh et al., 2013a; De Vernal and Marret, 2007). TOC contents (Fig. 3F), however, are still low in this interval probably due to elevated decomposition of freshly produced organic material within the oxic water column.

The TII warming also affected the deeper water body indicated by Mg/Caostracods values behaving similar to LST at least until 129.2 ka BP thereby implying a well-mixed water column (Fig. 3C). Decreasing Mg/Caostracods values thereafter may point to a LST warming induced thermo(haline) stratification and consequently less mixing of water masses. The initial and a supra-regional warming in our record is further supported by sharply increasing 18 18 δ Oostracods values (Fig. 4B) due to enhanced temperature-driven contribution of O from atmospheric precipitation (Bahr et al., 2006) as also suggested from the Sofular speleothem record (Badertscher et al., 2011; Figs. 1B, 4A).

49 Meltwater events and the Mediterranean reconnection at the Saalian-Eemian transition in the Black Sea

The early TII warming is characterised by a reduction in sedimentation rates from an average glacial level of 30 cm ka-1 to 20 cm ka-1 (Fig. 2), which has to be attributed to lower inputs of terrestrial material due to decreasing ice sheet melting and river discharge. 87 86 Additionally, decreasing K/Al and Sr/ Srostracods values also document the disappearance of northern runoff (Figs. 4D, 4E). The lower terrigenous input may further attributed to the gradual spreading of the Anatolian vegetation, thus favouring soil stabilisation (Shumilovskikh et al., 2013b). Especially, the spreading of oaks, which is sensitive to initial warming in the Mediterranean region, confirms warmer and wetter conditions on land (Shumilovskikh et al., 2013b).

5.4. The Mediterranean-Black Sea reconnection during the Eemian (128.1-122.5 ka BP)

In the course of global warming, the rising sea level allowed the Mediterranean-Black Sea reconnection via the Bosporus sill not later than about 129.5 ka BP (Grant et al., 2012; Fig. 4H) documented by several salinity-related parameters in our record. Steeply increasing 87 86 values of Mg/Caostracods, Sr/Caostracods, Sr/ Srostracods, and U/Caostracods at about 128.1 ± 0.7 ka BP (Figs. 3C, 4C, 4E, 4F) clearly indicate the intrusion of salty bottom waters. The significance of the observed time lag between the connection as expected from the eustatic sea level and as indicated from our Black Sea records is certainly limited due to age model uncertainties, a potentially differing Bosporus sill depth, and transient hydrological/environmental changes.

Even though, Mg/Caostracods is generally used as a proxy for temperature, it also reflects changing salinity if water chemistry experiences fundamental changes (Bahr et al., 2008). 87 86 Sr/ Srostracods values reaching the level of the modern Aegean and Black Seas (0.709157 and 0.709133; Major et al., 2006) also support the impact of Mediterranean waters at least on the deeper parts of the Black Sea. The higher concentration of U in seawater when compared -1 -1 with freshwater (Useawater: 3.2 μg l ; Ufreshwater/river: 0.24 μg l ; Martin and Whitfield, 1983) may favour elevated incorporation in ostracod shells thereby representing a further salinity proxy. Nevertheless, this assumption needs further investigation and verification. While Russell et al. (1994) also suggested that U/Ca in benthic and planktonic foraminifera reflects changes in seawater U concentrations; Ricketts et al. (2001) used the U/Ca ratio of ostracods as an indicator for varying redox conditions.

50 Meltwater events and the Mediterranean reconnection at the Saalian-Eemian transition in the Black Sea

The abrupt disappearance of the ostracods after 1 ka at 127.1 ka BP suggests a salinity exceeding the tolerance of the freshwater species Candona spp. (0-8 g kg-1; Aladin and Potts, 1996; Meisch, 2000) and/or the establishment of anoxic bottom waters. Favoured by rising temperature, rapid enhancement of primary productivity and preservation, the formation of the TOC-rich Eemian sapropel started at about 127.6 ka BP (Fig. 3F). Euxinic bottom water conditions, however, most likely evolved not before the disappearance of benthic foraminifera at 126.5 ka BP (Ammonia tepida; G. Schmiedl, personal communication). Although Ammonia tepida is able to survive anoxic periods its growth is generally favoured by oxic conditions (Aksu et al., 2002). As these foraminifera also suggest a bottom water salinity >7 g kg-1 about 1.5 ka after the reconnection (Bradshaw, 1957; Çağatay et al., 2009), a salinity-dependent disappearance of the ostracods appears to be more likely. If not of allochtonous origin, the occurrence of Mediterranean-type larval Mytilus galloprovincialis at 128.1 ka BP may also imply more saline surface waters (at least 15 g kg-1; Aral, 1999). A more reliable indication for increasing surface water salinities is given by the high percentages of marine dinocysts and the appearance of haptophyte produced alkenones at least since 127.6 ka BP (Figs. 4G, 3H).

The change in the hydrochemical conditions and biogeochemical processes after the Mediterranean-Black Sea reconnection allowed at least two significant periods of authigenic aragonite precipitation documented in parallel maxima of TIC and Sr/Cabulk at ca. 126.9 and 124.4 ka BP (Figs. 3D, 3E). Higher temperatures and the inflow of Sr-rich Mediterranean water favour precipitation of aragonite (Müller et al., 1972; Brauer et al., 2007; Milner et al., 2012) and is typically observed after Mediterranean reconnections (e.g. Kwiecien et al., 2008). Two layers dominated by coccoliths occur in between the aragonite layers (Fig. 3E) being consistent with the increasing ∑C37-39 alkenone/TOC predominantly derived from E. Huxleyi (e.g. Volkman et al., 1980; Coolen et al., 2009; Fig. 3H).

Just after the reconnection, TEX86 temperatures decrease within 1 ka (Fig. 3B), which might be not related to a shift in temperature but to the associated major hydrographic and biogeochemical changes impacting the behaviour of Thaumarchaeota. Vertical profiles from the modern Black Sea water column show an elevated nutrient availability including ammonium close to the redoxcline (about 80-130 m; e.g. Yakushev et al., 2006; Wakeham et al., 2007). Since the cold intermediate layer positioned between 60-80 m marks the maximum depth of the seasonal subsurface water renewals, the deeper waters at the redoxcline currently reveal an almost constant temperature of about 8 to 9°C throughout the year (Yakushev et al.,

51 Meltwater events and the Mediterranean reconnection at the Saalian-Eemian transition in the Black Sea

2007). Indeed, elevated abundances of ammonium-oxidising Thaumarchaeota and crenarchaeol (a diagnostic biomarker for Thaumarchaeota) concentrations are reported from the redoxcline of the Black and Baltic Seas thus strongly suggesting a TEX86 temperature signal produced in greater water depths (e.g. Wakeham et al.2003, 2007; Menzel et al., 2006; Lam et al., 2007; Coolen et al., 2007; Labrenz et al., 2010). Therefore, we attribute this temperature decrease in our record to a habitat change of the Thaumarchaeota community from euphotic surface to deeper waters. When using the subsurface calibration from Kim et al. (2008), the resulting temperature estimates are stable around 8-9°C (Fig. 3B) and closely correspond to the present-day temperature at the redoxcline (Yakushev et al., 2007; Locarnini et al., 2010). First measurements of eight surface sediment samples (0-1cm) from the NW,

NE, and SE Black Sea (Table S1) revealed a mean modern TEX86 temperature of ca. 9°C when using the same calibration, which is also similar to the mean Eemian temperature. The K’ alkenone-derived SST (U 37) of on average 17°C (Fig. 3B) are close to the present-day SST during spring/early summer coccolithophores blooming (Fig. 3B; Cokacar et al., 2004; Locarnini et al., 2010), which also supports the idea that mean annual subsurface (upper redoxcline) temperatures are estimated by the TEX86 proxy in the stratified Black Sea. Although the postulated Thaumarcheota habitat shift appears plausible, a conclusive explanation requires further microbial studies in the water column as well as biomarker measurements including sediments and suspended matter to attain an appropriate calibration.

5.5. The last two glacial-interglacial transitions in the Black Sea

This chapter briefly compares the three major periods, i.e., glacial, Termination, and interglacial of the last two glacial-interglacial transitions in the Black Sea. Both ending glacials reveal certain fluctuations in TEX86 surface temperatures with the penultimate showing a generally higher temperature level of about 3°C (Fig. 5A; Ménot and Bard, 2012). Colder conditions during the penultimate glacial from other European records encompassing K’ the same time intervals, e.g. U 37 reconstructions from the Iberian Margin (Pailler and Bard, 2002; Martrat et al., 2007), suggest a difference in regional climate rather than a general difference between the last and penultimate glacial. Today, the climate pattern differs in the NW and SE Black Sea region (e.g. Capet et al., 2012) with higher precipitation, milder winters, and less seasonal temperature differences in the SE part (Deniz et al., 2011). Comparable regional gradients might have been active during the late glacial periods probably explaining the deviations between both time periods.

52 Meltwater events and the Mediterranean reconnection at the Saalian-Eemian transition in the Black Sea

A stronger meltwater impact on the Black Sea’s salinity during the penultimate glacial 18 is reflected by δ Oostracods and Sr/Caostracods patterns with two discrete periods of intense meltwater supply, whereas a less pronounced meltwater impact lasting about 2 ka is recorded during the last late glacial (Figs. 5C, 5D; Bahr et al., 2006, 2008). One reason for this difference might be the 56% larger Saalian ice sheet allowing a significantly higher freshwater discharge during its retreat (Colleoni et al., 2009). Furthermore, extreme 87 86 radiogenic Sr/ Srostracods signatures especially during the penultimate BSWP-II-2, suggest a freshwater surplus from a potential Himalayan/Caspian source (Fig. 5E).

Fig. 5. Comparison between the last two glacial-interglacial transitions in the Black Sea (penultimate transition: this study; last transition: literature data). Literature data from the last transition are from the NW Black Sea: (A) TEX86 temperature (Ménot and Bard, 2012), (B) 18 Mg/Caostracods (Bahr et al., 2008), (C) δ Oostracods (Bahr et al., 2006), (D) Sr/Caostracods (Bahr et 87 86 al., 2008), and (E) Sr/ Srshells (Major et al., 2006). Core names and water depths are given. BSWP indicates meltwater pulses.

Relative to the timing of the Saalian (140 ka BP) and Weichselian (21 ka BP) glacial maxima, the initial TII warming appears later than during TI. This difference can be also

53 Meltwater events and the Mediterranean reconnection at the Saalian-Eemian transition in the Black Sea attributed to the larger Saalian ice sheet, expanding further to the east and south into the Eurasian continent (Svendsen et al., 2004), having a direct impact on the regional Black Sea climate. This could have locked the Black Sea region more systematically in a more glacial state lagging behind the global climate change. While TEX86 and Mg/Caostracods patterns (Fig. 5A, 5B) clearly document the Bølling-Allerød warming and the Younger Dryas (YD) cooling during T I (Bahr et al., 2008; Ménot and Bard, 2012), a rather gradual but also less pronounced warming is indicated by TEX86 during TII. Although the drop in Mg/Caostracods may suggests a YD-like event at the end of TII, several proxies in the sediment record argue against such cooling (compare Section 5.3).

Concerning the Black Sea interglacials the TEX86 temperature increase was considerably stronger and faster during the Eemian (Fig. 5A) resulting in distinctly higher summer surface temperatures. This was probably favoured by the lack of a YD-like event facilitating the undisturbed warming. Since the Mediterranean-Black Sea reconnection is coupled to the global sea-level change at the glacial Terminations the time between the onset of the interglacials in the Black Sea and its reconnection to the Mediterranean appears shorter during TII.

6. Conclusions

This multi-proxy study provides new comprehensive insights into the evolution from the penultimate glacial Black Sea “Lake” towards the Eemian Black Sea between 133.5 ± 0.7and 122.5 ± 1.7 ka BP.

The ending penultimate glacial in the SE Black Sea is impacted by two melt water 18 pulses (BSWP-II-1 and BSWP-II-2). A remarkable negative δ Oostracods excursion and unusual high radiogenic Sr-isotope ratios of ostracods during BSWP-II-2 most likely indicate a Himalayan source and enhanced discharge of the Amu Darya River entering the Black Sea via the Caspian Sea. Termination II, identified by abrupt IRD disappearance since 129.9 ± 0.7 ka BP, is characterised by a gradual warming favouring productivity and coupled carbonate precipitation. First interglacial conditions at around 128.5 ± 0.7 ka BP are documented by a rapid increase of summer SST followed by the Mediterranean-Black Sea reconnection at 128.1 ± 0.7 ka BP. Elevated primary productivity due to rising temperature initiated Eemian sapropel formation at about 127.6 ± 1.7 ka BP with bottom water euxinia most likely established after the disappearance of benthic foraminifera at 126.5 ka BP. Surface

54 Meltwater events and the Mediterranean reconnection at the Saalian-Eemian transition in the Black Sea temperature estimates from the penultimate late glacial (9°C) to the Eemian (17°C) reveal a warming of about 8°C in the Black Sea region.

This study on the penultimate glacial-interglacial transition demonstrates the complex relationship between global and regional climate change and the paleoceanographic evolution of the largely isolated epicontinental Black Sea.

Acknowledgements

The authors wish to thank Birgit Plessen and Rudolf Naumann from Helmholtz Centre Potsdam-German Research Centre for Geosciences (GFZ-Potsdam) for performing the oxygen isotope and X-ray powder diffraction (XRD) measurements, respectively. We express our gratitude to Norbert Nowaczyk (GFZ) and Gerhard Schmiedl (University of Hamburg) for providing helpful comments. Olga Körting and Sibylle Fink are gratefully acknowledged for their assistance during ostracod preparation. A.W. is indebted to L.A. Reed for the inspiring atmosphere during manuscript preparation. We thank G. Henderson, two anonymous reviewers and W.B. Ryan for helpful comments improving an earlier version of the manuscript. Finally, we thank the captain and crew of RV METEOR and the cruise leader Christian Borowski for their support during the M72/5 Black Sea cruise in 2007. This work was funded by the Deutsche Forschungsgemeinschaft (BE2116/20-1, AR 367/9-1, AR 367/9- 2, and FL 710/1-1) within the priority program 1266 INTERDYNAMIC and by the Leibniz Institute for Baltic Sea Research (IOW).

Appendix A. Supplementary material

Supplementary material related to this article can be found online at http://dx.doi.org/10.1016/j.epsl.2014.07.030. The Excel file of the Supplementary material provides information about the age model and the original data set used.

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64 Meltwater events and the Mediterranean reconnection at the Saalian-Eemian transition in the Black Sea

Supplementary Material

Age-depth model

Pollen stratigraphic correlation from Shumilovskikh et al. (2013) Quaternary Research 80, 349-360

22GC-3 depth [cm] age [ka BP] 794 123.0 812 127.5

XRF Ca correlation between 22GC-3 and 22GC-8

22GC-3 depth [cm] 22GC-8 depth [cm] age [ka BP] 790.00 782 122.0 794.00 787 123.0 796.50 790 123.6 799.00 792 124.3 802.50 795 125.1 809.00 799 126.8 812.00 801 127.5

18 230 18 Correlation between δ Oostracods with Th dated δ Ospeleothem from Sofular Cave

22GC-8 depth [cm] age [ka BP] 812.5 128.8 832.5 130.4 859.5 131.0 916.5 132.7

Final age-depth model for 22GC-8

22GC-8 depth [cm] age [ka BP] 782.0 122.0 787.0 123.0 790.0 123.6 792.0 124.3 795.0 125.1 799.0 126.8 801.0 127.5 812.5 128.8 832.5 130.4 859.5 131.0 916.5 132.7

65 Meltwater events and the Mediterranean reconnection at the Saalian-Eemian transition in the Black Sea

22GC-8 Sediment Geochemistry depth [cm] age [ka BP] TOC [%] TIC [%] Al [%] Ca [%] K [%] Sr [mg/kg] 784.5 122.50 4.27 5.70 3.81 19.73 1.04 907 785.5 122.70 5.10 4.86 4.08 17.43 1.14 2153 786.5 122.90 4.08 5.97 3.73 19.33 1.04 961 787.5 123.11 4.38 6.25 3.15 21.63 0.89 1303 788.5 123.32 5.37 4.18 4.83 14.07 1.35 779 789.5 123.53 5.06 2.95 5.98 9.94 1.67 504 790.5 123.79 4.88 3.28 5.76 10.90 1.65 712 791.5 124.10 3.59 4.82 4.95 15.10 1.40 1977 792.5 124.40 3.91 6.56 2.80 25.10 0.76 4830 793.5 124.69 4.10 4.40 5.01 15.75 1.33 1903 794.5 124.98 4.86 3.79 5.31 13.65 1.39 667 795.5 125.33 4.17 4.44 4.83 16.25 1.26 773 796.5 125.74 4.44 3.17 5.94 11.50 1.55 550 797.5 126.14 4.34 2.92 6.21 10.75 1.56 803 798.5 126.55 4.60 5.67 3.61 21.20 0.92 3927 799.5 126.94 4.05 6.05 3.45 20.85 0.88 3917 800.5 127.31 3.23 6.02 4.09 18.45 1.05 3110 801.5 127.56 2.58 4.84 5.01 15.30 1.27 2003 802.5 127.67 0.93 4.12 5.86 12.80 1.50 898 803.5 127.78 0.21 4.60 5.83 14.00 1.50 395 804.5 127.90 0.27 5.03 5.52 15.80 1.41 402 805.5 128.01 0.21 5.64 5.08 17.30 1.30 412 806.5 128.12 0.06 6.23 4.77 19.25 1.24 416 807.5 128.23 0.19 5.82 4.99 18.40 1.29 390 808.5 128.35 0.23 4.94 5.04 15.50 1.32 332 809.5 128.46 0.21 3.11 6.85 10.10 1.79 250 810.5 128.57 0.26 4.41 5.84 14.00 1.53 308 811.5 128.69 0.21 4.19 5.99 13.25 1.56 294 812.5 128.80 0.29 3.96 6.17 12.75 1.57 286 813.5 128.88 0.27 3.87 6.04 12.90 1.70 282 815.5 129.04 0.20 3.34 6.70 10.90 1.81 261 817.5 129.20 0.27 2.90 6.83 9.22 1.81 236 819.5 129.36 0.26 2.70 7.40 9.14 1.91 244 821.5 129.52 0.32 1.82 8.20 6.37 2.13 229 823.5 129.68 0.34 1.43 8.27 4.95 2.26 206 825.5 129.84 0.30 1.51 8.27 5.24 2.33 206 827.5 130.00 0.26 1.18 8.12 4.21 2.60 169 829.5 130.16 0.33 1.60 8.44 5.52 2.40 214 831.5 130.32 0.30 1.45 8.40 5.04 2.45 198 833.5 130.42 0.31 0.81 7.75 2.89 2.64 138 835.5 130.47 0.26 0.83 8.24 3.04 2.71 149

66 Meltwater events and the Mediterranean reconnection at the Saalian-Eemian transition in the Black Sea

22GC-8 Sediment Geochemistry (continued) depth [cm] age [ka BP] TOC [%] TIC [%] Al [%] Ca [%] K [%] Sr [mg/kg] 837.5 130.51 0.29 0.89 8.11 3.29 2.86 153 839.5 130.56 0.24 1.29 8.08 4.46 2.49 176 841.5 130.60 0.19 1.28 7.77 4.53 2.47 170 843.5 130.64 0.19 1.38 7.05 4.75 2.38 162 845.5 130.69 0.23 1.40 7.12 4.84 2.37 171 847.5 130.73 0.22 1.53 8.49 5.45 2.62 219 849.5 130.78 0.36 1.76 8.66 6.31 2.70 223 851.5 130.82 0.15 2.11 8.36 7.29 2.63 244 853.5 130.87 0.17 1.46 7.36 5.18 2.50 179 855.5* 130.91 0.47 0.10 5.65 1.84 2.25 1840 857.5 130.96 0.25 1.37 8.86 4.93 2.74 201 859.5 131.00 0.21 1.33 8.27 4.86 2.59 187 861.5 131.06 0.18 1.27 8.11 4.43 2.64 180 863.5 131.12 0.29 1.19 7.93 4.29 2.57 173 865.5 131.18 0.24 1.19 7.30 4.16 2.53 160 867.5 131.24 0.23 1.17 6.95 4.14 2.49 156 869.5 131.30 0.29 1.18 7.45 4.14 2.60 157 871.5 131.36 0.25 1.15 7.74 4.09 2.69 162 873.5 131.42 0.31 1.16 8.04 4.45 2.73 179 875.5 131.48 0.28 1.32 7.99 4.66 2.65 174 877.5 131.54 0.31 1.32 7.40 5.16 2.62 175 879.5 131.60 0.23 1.44 6.87 4.45 2.47 152 881.5 131.66 0.25 1.35 6.87 4.70 2.24 168 883.5 131.72 0.16 1.44 7.38 4.83 2.39 178 885.5 131.78 0.28 1.41 7.30 4.58 2.31 169 887.5 131.84 0.28 1.36 6.80 3.72 2.45 142 889.5 131.89 0.26 1.06 6.66 4.34 2.32 150 891.5 131.95 0.24 1.30 7.24 4.87 2.32 175 893.5 132.01 0.23 1.40 6.15 4.01 2.00 145 895.5 132.07 0.27 1.35 5.40 3.55 1.80 125 897.5 132.13 0.23 1.28 7.07 4.32 2.41 158 899.5 132.19 0.27 1.20 5.83 3.68 1.97 133 901.5 132.25 0.23 1.23 7.01 4.40 2.17 163 903.5 132.31 0.24 1.23 7.07 4.44 2.09 168 905.5 132.37 0.32 1.30 7.62 4.84 1.99 185 907.5 132.43 0.33 1.39 7.28 4.53 1.88 172 909.5 132.49 0.32 1.25 6.87 4.35 1.90 162 911.5 132.55 0.36 1.30 6.56 4.56 1.88 161 913.5 132.61 0.31 1.41 6.94 4.60 2.23 162 915.5 132.67 0.44 1.16 6.96 4.14 2.31 151 917.5 132.73 0.36 1.27 7.09 4.89 2.54 165

67 Meltwater events and the Mediterranean reconnection at the Saalian-Eemian transition in the Black Sea

22GC-8 Sediment Geochemistry (continued) depth [cm] age [ka BP] TOC [%] TIC [%] Al [%] Ca [%] K [%] Sr [mg/kg] 919.5 132.79 0.33 1.34 7.30 4.91 2.48 177 921.5 132.85 0.32 1.48 7.44 5.41 2.38 189 923.5 132.91 0.26 1.61 5.74 4.20 1.84 149 925.5 132.97 0.27 1.50 7.26 5.03 2.37 175 927.5 133.03 0.27 1.35 7.25 4.90 2.31 171 929.5 133.09 0.27 1.43 7.42 5.34 2.26 185 930.5 133.12 0.27 1.56 7.38 5.40 2.30 184 931.5 133.15 0.27 1.50 7.69 5.54 2.38 197 932.5 133.18 0.18 1.16 7.25 5.35 2.20 197 933.5 133.21 0.18 1.47 7.13 5.32 2.19 189 934.5 133.24 0.28 1.45 7.05 5.21 2.32 169 935.5 133.27 / / 7.16 5.08 2.37 170 936.5 133.30 0.30 1.29 7.53 4.72 2.36 171 937.5 133.33 / / 7.58 4.35 2.39 165 938.5 133.36 0.32 1.17 7.56 4.32 2.41 166 939.5 133.39 / / 7.74 4.15 2.57 167 940.5 133.42 0.22 1.34 7.31 4.42 2.47 166 941.5 133.45 / / 7.42 4.91 2.47 174 942.5 133.48 0.21 1.51 7.44 4.89 2.52 173 943.5 133.51 0.27 1.41 7.34 4.79 2.36 170 944.5 133.54 0.23 1.44 7.43 4.76 2.31 171 *tephra

TIC data from 806.5-944.5cm were published in: Shumilovskikh, L. S., Arz, H. W., Wegwerth, A., Fleitmann, D., Marret, F., Nowaczyk, N., Tarasov, P., and Behling, H. (2013b). Vegetation and environmental changes in Northern Anatolia between 134 and 119 ka recorded in Black Sea sediments. Quaternary Research 80, 349-360

68 Meltwater events and the Mediterranean reconnection at the Saalian-Eemian transition in the Black Sea

22GC-8 Ostracod Geochemistry depth [cm] age [ka BP] Sr/Ca [mmol/mol] Mg/Ca [mmol/mol] U/Ca *10-6 87Sr/86Sr 800 127.1 1.56 4.63 1.62 0.70903 802 127.6 1.84 6.94 3.18 0.70907 804 127.8 1.80 5.60 1.81 0.70907 805 128.0 1.36 / 806 128.1 1.53 3.53 1.14 0.70895 808 128.3 1.42 3.21 0.63 0.70897 810 128.5 1.41 3.31 / / 812 128.7 1.32 3.83 0.70 0.70895 814 128.9 1.33 3.99 0.61 / 816 129.1 1.22 4.11 / 0.70899 818 129.2 1.20 4.22 0.61 0.70905 820 129.4 1.18 3.03 / / 822 129.6 1.18 3.49 / / 824 129.7 1.19 2.08 / / 826 129.9 1.18 3.18 / / 828 130.0 1.20 3.19 / / 830 130.2 1.19 2.85 0.99 / 832 130.4 1.21 3.07 / / 834 130.4 1.11 1.90 / 0.70934 836 130.5 1.09 1.56 / / 838 130.5 1.08 1.58 / / 840 130.6 1.03 1.96 / 0.70921 842 130.6 1.05 1.73 / / 844 130.7 0.91 2.00 / / 846 130.7 0.80 1.99 / / 848 130.7 0.81 2.26 0.49 0.70932 850 130.8 0.80 2.37 0.69 / 860 131.0 0.82 1.71 / / 862 131.1 0.81 1.43 / / 864 131.1 0.83 1.19 / / 866 131.2 0.80 1.27 / / 868 131.3 0.82 1.37 / 0.70945 870 131.3 0.83 1.14 / / 872 131.4 0.81 1.38 / / 874 131.4 0.80 1.50 / / 876 131.5 0.89 1.24 / / 878 131.6 0.86 1.26 0.41 0.70925 880 131.6 0.88 1.61 0.42 / 882 131.7 0.96 1.13 / / 884 131.7 0.98 1.21 / / 886 131.8 0.97 1.43 / /

69 Meltwater events and the Mediterranean reconnection at the Saalian-Eemian transition in the Black Sea

22GC-8 Ostracod Geochemistry (continued) depth [cm] age [ka BP] Sr/Ca [mmol/mol] Mg/Ca [mmol/mol] U/Ca *10-6 87Sr/86Sr 888 131.9 1.04 1.57 / / 890 131.9 1.03 1.40 / 0.70917 892 132.0 1.04 1.52 / / 894 132.0 1.01 1.49 0.46 / 896 132.1 1.05 1.40 / / 898 132.1 1.12 1.47 / / 900 132.2 1.13 1.31 / 0.70910 902 132.3 1.08 1.26 / / 904 132.3 1.10 1.19 / 0.70915 906 132.4 1.10 1.20 0.59 / 908 132.4 0.97 1.23 / / 910 132.5 0.91 1.34 / / 912 132.6 0.89 1.36 / / 914 132.6 0.85 1.26 / / 916 132.7 0.85 1.19 / / 918 132.7 0.87 1.22 / / 920 132.8 0.85 1.42 / 0.70921 922 132.9 0.83 1.49 / / 924 132.9 0.85 1.26 / / 926 133.0 0.83 1.67 0.41 / 928 133.0 0.87 1.40 / / 930 133.1 0.81 1.34 / / 932 133.2 0.83 1.52 / / 934 133.2 0.77 1.51 / / 936 133.3 0.80 1.44 0.38 / 938 133.3 0.80 1.14 0.42 0.70909 940 133.4 0.83 1.39 / / 942 133.5 0.89 1.42 / / 944 133.5 0.84 1.31 / 0.70913

70 Meltwater events and the Mediterranean reconnection at the Saalian-Eemian transition in the Black Sea

22GC-8 GDGT and TEX86

1 2 depth [cm] age [ka BP] BIT TEX86 GDGT-0/crenarchaeol LST [°C] SST 0-200m [°C] 784.5 122.50 0.07 0.41 1.20 / 9.0 785.5 122.70 0.06 0.41 1.23 / 8.8 787.5 123.11 0.05 0.41 1.22 / 8.9 788.5 123.32 0.06 0.42 1.26 / 9.5 789.5 123.53 0.06 0.41 1.25 / 8.8 790.5 123.79 0.04 0.41 1.15 / 8.8 791.5 124.10 0.04 0.40 1.15 / 8.3 792.5 124.40 0.05 0.40 0.97 / 8.7 793.5 124.69 0.02 0.40 0.86 / 8.4 794.5 124.98 0.04 0.41 1.01 / 9.0 795.5 125.33 0.04 0.40 0.99 / 8.5 796.5 125.74 0.05 0.40 1.03 / 8.7 797.5 126.14 0.06 0.43 0.99 / 9.8 798.5 126.55 0.06 0.41 0.87 / 8.9 799.5 126.94 0.04 0.41 0.85 / 9.0 800.5 127.31 0.05 0.44 0.76 / 10.1 801.5 127.56 0.04 0.49 0.59 14.39 / 802.5 127.67 0.04 0.52 0.41 15.80 / 803.5 127.78 0.12 0.59 0.36 19.64 / 804.5 127.90 0.10 0.63 0.30 21.63 / 805.5 128.01 0.10 0.62 0.30 21.10 / 806.5 128.12 0.13 0.63 0.33 21.73 / 807.5 128.23 0.11 0.61 0.34 20.52 / 808.5 128.35 0.09 0.51 0.41 15.56 / 809.5 128.46 0.13 0.41 0.49 10.27 / 810.5 128.57 0.08 0.42 0.44 11.16 / 811.5 128.69 0.08 0.42 0.45 10.90 / 812.5 128.80 0.10 0.42 0.44 10.74 / 813.5 128.88 0.10 0.41 0.44 10.39 / 814.5 128.96 0.11 0.42 0.49 10.84 / 815.5 129.04 0.13 0.42 0.46 10.79 / 817.5 129.20 0.15 0.41 0.46 10.20 / 818.5 129.28 0.14 0.41 0.48 10.27 / 819.5 129.36 0.13 0.40 0.44 9.73 / 821.5 129.52 0.13 0.40 0.45 9.92 / 823.5 129.68 0.13 0.39 0.44 9.48 / 824.5 129.76 0.14 0.39 0.46 9.54 / 825.5 129.84 0.14 0.39 0.44 9.16 / 827.5 130.00 0.12 0.38 0.48 8.81 / 828.5 130.08 0.14 0.39 0.44 9.19 / 829.5 130.16 0.14 0.39 0.41 9.57 / 830.5 130.24 0.13 0.39 0.43 9.47 / 831.5 130.32 0.13 0.39 0.40 9.44 /

71 Meltwater events and the Mediterranean reconnection at the Saalian-Eemian transition in the Black Sea

22GC-8 GDGT and TEX86 (continued)

1 2 depth [cm] age [ka BP] BIT TEX86 GDGT-0/crenarchaeol LST [°C] SST 0-200m [°C] 833.5 130.42 0.17 0.39 0.44 9.17 / 835.5 130.47 0.19 0.39 0.44 9.54 / 837.5 130.51 0.16 0.43 0.39 11.24 / 839.5 130.56 0.20 0.38 0.44 9.00 / 841.5 130.60 0.19 0.39 0.47 9.55 / 843.5 130.64 0.13 0.39 0.44 9.64 / 845.5 130.69 0.19 0.41 0.51 10.64 / 848.5 130.76 0.15 0.37 0.49 8.51 / 849.5 130.78 0.15 0.37 0.52 8.53 / 850.5 130.80 0.15 0.37 0.50 8.23 / 859.5 131.00 0.20 0.37 0.47 8.45 / 868.5 131.27 0.22 0.39 0.52 9.34 / 878.5 131.57 0.24 0.37 0.49 8.34 / 884.5 131.75 0.19 0.38 0.46 8.86 / 894.5 132.04 0.25 0.36 0.48 7.74 / 896.5 132.10 0.29 0.38 0.47 9.12 / 906.5 132.40 0.30 0.37 0.46 8.18 / 908.5 132.46 0.28 0.37 0.46 8.58 / 916.5 132.70 0.26 0.43 0.49 11.41 / 922.5 132.88 0.26 0.38 0.46 8.79 / 926.5 133.00 0.22 0.39 0.47 9.16 / 932.5 133.18 0.28 0.37 0.49 8.52 / 943.5 133.51 0.20 0.38 0.52 9.01 /

1 LST = lake surface temperature; calculated after Powers et al. (2010) 2 SST = sea surface-subsurface temperature (0-200 m water depth); calculated after Kim et al., (2008)

Core Top Samples (0-1 cm sediment depth) GDGT and TEX86

1 sample water Lat°N Long°E TOC BIT TEX86 GDGT-0/ SST depth[m] [deg/min] [deg/min] [%] crenarchaeol 24 MUC-1 207.7 41° 28.69'N 37° 11.74'E 2.2 0.03 0.40 1.05 8.3 14 MUC-7 274.9 44° 35.80'N 36° 21.28'E 1.8 0.03 0.41 0.97 8.9 25 MUC-2 418.1 42° 6.210'N 36° 37.44'E 5.4 0.03 0.40 1.08 8.5 17 MUC-4 523.2 44° 41.05'N 36° 2.01'E 1.9 0.04 0.42 0.90 9.2 22 MUC-1 844.5 43° 13.56'N 36° 29.60'E 3.8 0.05 0.40 1.06 8.6 18 MUC-4 978.9 44° 34.17'N 36° 0.78'E / 0.03 0.41 0.91 8.9 7 MUC-5 1531.3 43° 59.96'N 32° 1.76'E 3.4 0.04 0.44 0.96 10.2 20 MUC-3 2047.2 43° 57.26'N 35° 38.47'E 5.3 0.06 0.43 1.02 10.0 1 SST (0-200 m water depth), calculated after Kim et al. (2008)

72 Meltwater events and the Mediterranean reconnection at the Saalian-Eemian transition in the Black Sea

22GC-8 Alkenones

K' 1 depth [cm] age [ka BP] U 37 SST [°C] C37+38+39 /TOC [µg/g] 784.5 122.50 0.59 16.3 37.37 785.5 122.70 0.71 19.9 12.91 786.5 122.90 0.73 20.5 47.98 787.5 123.11 0.71 20 46.99 788.5 123.32 0.67 18.8 28.75 789.5 123.53 0.65 18.2 13.44 790.5 123.79 0.62 17.1 11.30 791.5 124.10 0.57 15.7 8.64 792.5 124.40 0.53 14.6 17.68 794.5 124.98 0.56 15.4 43.93 795.5 125.33 0.50 13.7 60.46 796.5 125.74 0.55 15.1 39.25 797.5 126.14 0.46 12.3 14.27 798.5 126.55 0.64 17.7 3.12 799.5 126.94 0.57 15.7 11.43 800.5 127.31 0.59 16.2 6.86 801.5 127.56 0.55 14.9 3.98 1 SST calculated after Conte et al. (2006)

73 Black Sea temperature response to glacial millennial-scale climate variability

Black Sea temperature response to glacial millennial-scale climate variability

Antje Wegwertha, Andrey Ganopolskib, Guillemette Ménotc, Jérôme Kaisera, Olaf Dellwiga, Edouard Bardc, Frank Lamyd, and Helge W. Arza

a Leibniz-Institute for Baltic Sea Research Warnemünde (IOW), Marine Geology, Rostock, Germany b Potsdam Institute for Climate Impact Research (PIK), Earth System Analysis, Potsdam, Germany c CEREGE, Aix-Marseille University, Collège de France, CNRS, IRD, Aix en Provence, France d Alfred-Wegener-Institut Helmholtz-Zentrum für Polar und Meeresforschung, Bremerhaven, Germany

74 Black Sea temperature response to glacial millennial-scale climate variability

Abstract

The Eurasian inland propagation of temperature anomalies during glacial millennial- scale climate variability is poorly understood, but this knowledge is crucial to understanding hemisphere-wide atmospheric teleconnection patterns and climate mechanisms. Based on biomarkers and geochemical paleothermometers, a pronounced continental temperature variability between 64,000 and 20,000 years ago, coinciding with the Greenland Dansgaard- Oeschger cycles, was determined in a well-dated sediment record from the formerly enclosed Black Sea. Cooling during Heinrich events was not stronger than during other stadials in the Black Sea. This is corroborated by modeling results showing that regular Dansgaard- Oeschger cycles penetrated deeper into the Eurasian continent than Heinrich events. The pattern of coastal ice-rafted detritus suggests a strong dependence on the climate background state, with significantly milder winters during periods of reduced Eurasian ice sheets and an intensified meridional atmospheric circulation.

Key Points

• Impact of Dansgaard-Oeschger cycles on the Black Sea surface temperature • Absence of extra cooling during Heinrich events • Different modes of DO cycles and HE propagation suggested by a proxy-model comparison

75 Black Sea temperature response to glacial millennial-scale climate variability

1. Introduction

The last glacial millennial-scale climate fluctuations involved major atmosphere-ocean reconfigurations that resulted in strong changes in temperature, precipitation, and wind fields in the Northern Hemisphere [Bond et al., 1993; Dansgaard et al., 1993; Bard et al., 2000; Wang et al., 2001]. Greenland ice cores and marine archives from the North Atlantic provide evidence of abrupt changes from cold stadials to warmer interstadials on a millennial timescale, so-called Dansgaard-Oeschger (DO) cycles [Bond et al., 1993; Dansgaard et al., 1993]. Stadials associated with massive iceberg discharges occurring on a multimillennial timescale, referred to as Heinrich events (HE) [Heinrich, 1988; Bond et al., 1993], were significantly colder in some areas of the midlatitude North Atlantic [Heinrich, 1988; Bond et al., 1993; Bard et al., 2000; Martrat et al., 2007]. These abrupt climate changes associated with DO cycles and HE were probably caused by major reorganizations in the Atlantic Meridional Overturning Circulation (AMOC) and strongly amplified by shifts in sea ice expansion and other forms of climate feedback [Bond et al., 1993; Ganopolski and Rahmstorf, 2001; Van Meerbeeck et al., 2011; Zhang et al., 2014]. Climate model simulations [Ganopolski and Rahmstorf, 2001; Zhang et al., 2014] suggest very different spatial patterns of oceanic and continental temperature anomalies during DO cycles and HE in the Northern Hemisphere. While stadials (interstadials) have been attributed to a weak (strong) AMOC associated with decreased (increased) ocean heat transport and most likely sea ice controlled, HE can be explained by an off mode of the AMOC [Bond et al., 1993; Ganopolski and Rahmstorf, 2001; Van Meerbeeck et al., 2011; Zhang et al., 2014].

Proxy-based temperature records including DO cycles and HE come mainly from the North Atlantic sector. Interstadial warming was shown to be strongest in Greenland [Kindler et al., 2014] and HE cooling most prominent in the midlatitude North Atlantic and western Mediterranean Sea [Bard et al., 2000; Martrat et al., 2004, 2007]. Climate models suggest that this spatial heterogeneity in the North Atlantic climate response was transmitted to the continental interior, with DO cycle temperature anomalies penetrating deeper into the continent than HE-associated cooling [Ganopolski and Rahmstorf, 2001; Zhang et al., 2014]. Pollen and stalagmite records from eastern and southern Europe reveal a pronounced variability in rainfall that coincided with DO cycles characterized by wetter conditions and lower wind stress during interstadials [Allen et al., 1999; Tzedakis et al., 2004; Fleitmann et al., 2009; Fletcher et al., 2010] (Figure S1 in the supporting information). While studies from the western Mediterranean Sea [Martrat et al., 2004; Allen et al., 1999] recognized distinct

76 Black Sea temperature response to glacial millennial-scale climate variability differences between HE and non-HE stadials, the detailed patterns of DO cycles and HE in the eastern Mediterranean region seem to be more ambiguous [Tzedakis et al., 2004; Fleitmann et al., 2009; Shumilovskikh et al., 2014]. Farther to the east, in the southeastern Asian climate domain, the significant influence of DO cycles and HE on the monsoonal system is evidenced through changes in precipitation and wind intensity, based on data reconstructed from Chinese stalagmites [Wang et al., 2001] and loess deposits [Sun et al., 2012]. Here stadials are characterized by the reduced precipitation of East Asian summer monsoons and HE stadials by a strengthening of the winter monsoon circulation linked to an enhanced zonality in the Northern Hemisphere atmospheric circulation [Wang et al., 2001; Sun et al., 2012]. In contrast to the hydroclimatic impacts over Eurasia, little is known about the thermal expression of DO cycles and HE. To test climate models and thus complete our mechanistic understanding of glacial climate oscillations requires quantitative temperature reconstructions.

Here we present three temperature-related proxy records from sediment core 25GC-1, recovered in the southeastern Black Sea, a key region bridging the Atlantic and continental Eurasian climate realms. The record spans the period from ~64 to 20 thousand years ago (ka) and includes Marine Isotope Stages (MIS) 4 and 2 (Figure 1b), when the global sea level was lower and the Black Sea was disconnected from the Mediterranean Sea [Nowaczyk et al., 2012; Shumilovskikh et al., 2014]. Based on these data, we provide ample evidence for the impact of the MIS 3 millennial-scale climate events on the Black Sea region.

2. Material and Methods

The gravity core 25GC-1 was retrieved during the R/V Meteor cruise M72/5 in 2007 from the southeastern Black Sea (Archangelsky Ridge; 42°06.2′N, 36°37.4′E) at a water depth of 418m. The sediment core contains an undisturbed sediment sequence (952–307 cm) spanning the period from 64 to 20 ka (MIS 4 – MIS 2). An initial age model was based on the absolutely dated Laschamp geomagnetic excursion (40.70 ± 0.95 ka), the Campanian Ignimbrite tephra (Y5; 39.28 ± 0.11 ka), and eight accelerator mass spectrometry radiocarbon dates [Nowaczyk et al., 2012]. Subsequently, fine tuning of the records of coastal ice-rafted detritus (IRDC), carbonate content, Ca XRF counts, and the high-resolution magnetic susceptibility of core 25GC-1 to the GICC05 chronology of the North Greenland Ice Core

77 Black Sea temperature response to glacial millennial-scale climate variability

Project (NGRIP) resulted in a final chronology with an average sampling resolution of about 230 years. The age model (Figure 1b) is discussed in greater detail in Nowaczyk et al. [2012].

The initial preparation for glycerol dialkyl glycerol tetraethers (GDGTs) analysis was carried out at the Leibniz Institute for Baltic Sea Research Warnemünde (IOW, Rostock, Germany). The 148 homogenized samples (4–6 g sediment; average resolution of 0.29 ka) were subjected to accelerated solvent extraction (Thermo Scientific™ Dionex™ ASE 350) with a dichloromethane/methanol mixture (DCM/MeOH 9:1), followed by desulfurization with activated copper pieces and GDGT separation using another dichloromethane/methanol mixture (DCM/MeOH 1:1) over Pasteur pipettes plugged with activated Al2O3. After the addition of a C46 GDGT standard and filtration through a 0.45 μm polytetrafluorethylene filter, GDGTs were measured at the CEREGE (Aix-en-Provence, France) using a high- performance liquid chromatography/atmospheric pressure chemical ionization mass spectrometer (HPLCMS1100 Series, Hewlett Packard, USA) equipped with a Prevail Cyano column [Ménot and Bard, 2012]. Single-ion monitoring was used for peak identification. The mean standard deviation for duplicate runs was 0.006 for TEX86 (TetraEther IndeX of lipids with 86 carbon atoms), corresponding to 0.3°C. The laboratory setting was tested through two interlaboratory comparison experiments carried out in 2009 and 2013 [Schouten et al.,

2013a]. The TEX86 values were converted to mean annual surface temperature estimates using the global lake calibration [Powers et al., 2010]. The calculated temperatures are within the calibration range of lake surface temperatures, which is 4–28°C [Powers et al., 2010]. The

TEX86 paleothermometer (TetraEther IndeX of lipids with 86 carbon atoms) records the ambient water temperature during the growth of Thaumarchaeota [Schouten et al., 2013b].

The TEX86 is assumed to reflect the mean near-surface annual temperatures during glacial stages of the Black Sea [Wegwerth et al., 2014].

Mg/Ca ratios of benthic ostracods were determined at the IOW (Rostock, Germany) as follows: Intact valves of adult Candona spp. were handpicked from wet-sieved sediments (>150 μm; 135 samples with an average resolution of 0.32 ka) and cleaned under a binocular microscope using a thin brush and a few drops of deionized water. Up to five valves were dissolved in 2.5 mL of 2 vol% HNO3 (subboiled) and then centrifuged, followed by the analysis of a 1 mL aliquot for Ca and Mg by inductively coupled plasma mass spectrometry (iCAP Q; Thermo Fisher Scientific). Separate calibration solutions were prepared for Mg (six) and Ca (five) (due to contamination). Matrix effects were compensated by using Be and Rh as internal standards. Possible contamination by detrital material was monitored by Al.

78 Black Sea temperature response to glacial millennial-scale climate variability

Precision (2.2%) and accuracy (1.4%) were checked every five samples by comparison with the international reference material ECRM 752-1 (BAS). The Mg/Ca ratios of benthic ostracod carbonate shells (Mg/Caostracods) document the mean annual changes in bottom water temperatures [Bahr et al., 2008; Wegwerth et al., 2014]. The methods for the total inorganic carbon (TIC) and IRDC analyses are described in Nowaczyk et al. [2012].

Model results were obtained using the CCSM3 fully coupled comprehensive general circulation model of the National Center for Atmospheric Research. Details regarding the experimental design are provided in Zhang et al. [2014].

3. Millennial-Scale Temperature Variability in the Black Sea

The mean annual Black Sea surface TEX86 temperature record closely mimics the DO oscillations seen in the NGRIP Greenland ice core [Dansgaard et al., 1993; Kindler et al., 2014] (Figure 1a) and ranges between ~5°C during stadials and ~9°C during interstadials

(today: 15°C). Our TEX86 and Mg/Caostracods records suggest that HE and non-HE stadials cooled to about 5.5°C, with no significant difference between them. Variations in

Mg/Caostracods, which closely resemble the TEX86 surface temperature record during MIS 3 (Figure 1a), correspond to temperature changes of roughly 2–4°C [Bahr et al., 2008] at this intermediate water depth and confirm the DO-related temperature changes, although we cannot provide absolute temperature values due to the absence of an appropriate calibration for the formerly limnic Black Sea. The magnitude of the temperature variability for the bottom water is therefore comparable to the amplitudes of the TEX86 surface temperature and suggested a generally well mixed water column at the coring site during the last glacial period and especially during MIS 3 (Figure 1a). The pronounced DO patterns in our temperature records closely parallel the humidity changes in Northern Anatolia (Figure S1), as inferred from pollen assemblages in the same sediment core [Shumilovskikh et al., 2014] and from the well-dated Sofular stalagmite record [Fleitmann et al., 2009]. Contrasting evidence comes from a 40 ka old sediment record in the northwestern Black Sea [Ménot and Bard, 2012], where cooling occurred only during HE stadials and distinct DO patterns are missing (Figure S1). This discrepancy remains largely unexplained, but it contrasts with the overregional paleoenvironmental evidence [Allen et al., 1999; Tzedakis et al., 2004; Fleitmann et al., 2009; Fletcher et al., 2010; Shumilovskikh et al., 2014] and may in part be related to both the large

79 Black Sea temperature response to glacial millennial-scale climate variability heterogeneity in the climatic conditions and the sediment sources of the northern Black Sea drainage basins.

Figure 1. Climate conditions in Greenland and the Black Sea “Lake” during the last glacial (65–20 ka). (a) Greenland temperature [Kindler et al., 2014] (NGRIP); TEX86-based mean annual lake surface temperature (LST). The variability in bottom water temperature was estimated from the Mg/Ca of benthic ostracods, with ΔT reflecting a minimum estimate of the temperature changes [Bahr et al., 2008]. Total inorganic carbon (TIC) served as a proxy for phytoplankton productivity. Winter severity was inferred from the accumulation rate of coastal ice-rafted detritus (IRDC, modified [Nowaczyk et al., 2012]). The Eurasian ice volume is shown relative to present [Bintanja and van deWal, 2008]. Red numbers and yellow bars denote the warm interstadials of DO cycles; HE denotes Heinrich events. (b) Age model of core 25GC-1 with absolute age control points and fine tuning using sedimentary parameters [Nowaczyk et al., 2012].

80 Black Sea temperature response to glacial millennial-scale climate variability

Together with enhanced concentrations of dinoflagellates [Shumilovskikh et al., 2014] and Thaumarchaeota (Figure S1), other proxies from our core, such as the total inorganic carbon (TIC) record, suggest a strong environmental response with enhanced biological productivity during interstadials.

TIC mainly reflects elevated calcite precipitation induced by augmented phytoplankton blooms during relatively warm interstadials [Shumilovskikh et al., 2014] (Figure 1a) and therefore qualitatively supports the temperature reconstructions. IRDC accumulation rates

(Figure 1a) also show a clear DO pattern, with IRDC virtually missing during interstadials and high values suggesting the occurrence of extremely cold winter temperatures that favored coastal ice formation during stadials [Nowaczyk et al., 2012]. The pronounced millennial- scale environmental changes are further supported by the enhanced content of arboreal pollen during interstadials and a reduction thereof during the stadials in our record [Shumilovskikh et al., 2014] (Figure S1), thus reflecting the pronounced response of Northern Anatolian vegetation to the variability in temperature and especially in precipitation during DO cycles. The enhanced arboreal pollen content is indicative of milder and wetter conditions during interstadials [Shumilovskikh et al., 2014]. Wetter conditions may have contributed to enhanced planktonic activity through high inputs of macronutrients from land.

4. Long-Term Temperature Changes in the Black Sea

On a multimillennial timescale, the relatively low accumulation rates of IRDC during MIS 3 indicate much milder winters compared to MIS 4 and MIS 2. Furthermore, the low TIC content during MIS 2 suggests strongly reduced planktonic productivity even during DO interstadials 3 and 4 (Figures 1a and S1). The milder winters during the first part of MIS 3 together with the increased humidity in Northern Anatolia [Shumilovskikh et al., 2014] (Figure S1) point to a decreased continentality during periods of reduced continental ice sheet volume [Sánchez Goñi et al., 2008; Bintanja and van de Wal, 2008; Helmens, 2014] (Figure 1a) and therefore both a stronger zonality in the atmospheric circulation and a far-field impact toward Eurasia of the North Atlantic climate. A stronger oceanic influence on continental climate during a period of low ice sheet volume was also inferred from palynological and microfossil records from the Iberian continental margin [Sanchez Goñi et al., 2013]. Similarly, the reduced latitudinal δ18O gradients across Greenland (NGRIP versus GRIP and GISP2 sites) during MIS 3 were related to a reduction in ice sheet volume and sea ice extent and to changes in the atmospheric circulation [Seierstad et al., 2014].

81 Black Sea temperature response to glacial millennial-scale climate variability

The long-term trend in our Black Sea IRDC record mimics the Eurasian ice sheet volume (Figure 1a) and suggests a high impact of the Eurasian ice sheet on the severity of winter in the Black Sea region. Shortly before the MIS 3/MIS 2 transition (since ~34 ka),

IRDC accumulation rates are relatively high not only during HE but also during non-HE stadials, indicating long-lasting cold winters and presumably stronger seasonal temperature contrasts than during MIS 3, since the mean annual surface temperatures (TEX86) remain at about the same level during stadials (Figure 1a). The progressively colder winters in the Black Sea region after 34 ka presumably reflect insolation-driven long-term changes, i.e., the expansion of Eurasian ice sheets [Bintanja and van de Wal, 2008] (Figure 1a) toward the Last Glacial Maximum (LGM) and the southward migration of the atmospheric polar front. Likewise, Chinese Loess Plateau records [Sun et al., 2012] that document distinct monsoonal changes in phase with Greenland DO cycles between 60 and 34 ka suggest a reduced sensitivity to the North Atlantic climate variability between 34 ka and the LGM, when ice sheets expanded [Bintanja and van de Wal, 2008]. A reduced inland propagation of the North Atlantic climate at times coinciding with the presence of large ice sheets is further corroborated by paleoclimate studies in central and eastern Europe. These have documented an increase in ice masses that during MIS 2 [Pollard and Barron, 2003; Feurdean et al., 2014] formed a barrier against Atlantic air toward the east, with strong anticyclonal circulation resulting in a continental climate over Europe. Therefore, the Black Sea records provide additional evidence for the role of the Northern Hemispheric ice sheet in modulating the North Atlantic and Eurasian continental climate systems.

5. Proxy- and Model-Based Assessments of Spatial Temperature Patterns in the North Atlantic-Eurasian Region

To compare the sensitivity of DO- and HE-related temperature changes in the continental Black Sea region with those of the North Atlantic sector, we compiled sea surface temperature (SST) records from the midlatitude North Atlantic [Martrat et al., 2007] and the western Mediterranean Sea [Martrat et al., 2004]. These marine records are consistent with DO cycles but, unlike the Black Sea record, they show considerably colder conditions during some HE compared to non-HE stadials [Martrat et al., 2004, 2007] (Figure 2).

Based on the temperature records, we were able to estimate the inland propagation of temperature anomalies related to DO cycles, by calculating for each considered proxy-based

82 Black Sea temperature response to glacial millennial-scale climate variability temperature record [Martrat et al., 2004, 2007; Kindler et al., 2014] the temperature amplitudes for DO cycles.

Figure 2. Temperature variability in the North Atlantic and Europe during the last glacial (65–20 ka). Greenland temperature [Kindler et al., 2014] (NGRIP), alkenone-based sea surface temperature (SST) from the North Atlantic/Iberian Margin [Martrat et al., 2007] (core MD01-2444) and from the western Mediterranean Sea/Alboran Sea [Martrat et al., 2004] (Ocean Drilling Program (ODP) Site 977A) are shown together with the TEX86-based lake surface temperature (LST) from the southeastern Black Sea (core 25GC-1, this study). Red numbers denote DO interstadials, blue numbers HE.

First, the temperature amplitudes for DO cycles (ΔTDO; DO 3–16) were calculated based on the difference between the averaged maximum interstadial and averaged minimum non-HE stadial temperatures (Figure 3a). The extra cooling related to the HE (ΔTHE) was defined as the difference between the averaged minimum temperatures of HE (HE 2–6) and non-HE stadials (Greenland stadials (GS) 4–17). Finally, ΔTDO and ΔTHE were compared with the model-based ΔT air temperature fields calculated in the same way as for the proxy-based

83 Black Sea temperature response to glacial millennial-scale climate variability

ΔT. Although the proxy data reflect SSTs (except for Greenland [Kindler et al., 2014]) and the climate model simulations reveal surface air temperatures, we assume that for perennially sea ice free conditions (at least in the pelagic zones), the temperature changes occurring on these timescales were similar.

The proxy-based ΔTDO (compare Figures 2 and 3) decreases from ~12°C over

Greenland to ~2°C toward the midlatitudes (ΔTDO 1.4°C at the Iberian margin and ΔTDO

2.2°C in the western Mediterranean Sea). A ΔTDO of ~1.8°C for the isolated Black Sea “Lake” is comparable with the temperature amplitudes from the midlatitude North Atlantic and Mediterranean Sea, suggesting a close atmospheric connection between these regions. By contrast, the proxy-based ΔTHE in the midlatitude eastern North Atlantic and western Mediterranean Sea seem to be larger as in the continental Greenland and the SE Black Sea, which would suggest that HE cooling did not exert similar effects on continental regions far from the eastern North Atlantic (Figures 2 and 3). Supporting evidence for the different patterns comes from temperature anomalies calculated for the individual HE and the non-HE stadials (Figure S2 in the supporting information).

A compilation of climate model simulations [Ganopolski and Rahmstorf, 2001; Zhang et al., 2014] for DO cycles and HE yields qualitatively similar temperature patterns (Figures 3b and 3c). Maximum air temperature anomalies > 8°C associated with DO cycle warming are located over the Nordic Seas (Figure 3b) and can be attributed to an intensification of the AMOC, the northward shift of deep water formation areas, and a significant retreat of sea ice. Such DO-type anomalies are subsequently propagating eastward from the North Atlantic over the entire extratropics of the Northern Hemisphere, through atmospheric circulation. Simulated positive temperature anomalies of approximately 4–6°C over both the Mediterranean Sea and the Black Sea point to similar temperature changes in the midlatitude North Atlantic and on the continent, as determined by the proxy records (Figure 3; for individual warmings, see Figure S3 in the supporting information). HE-type cooling (i.e., additional cooling compared to non-HE stadials) can be attributed to a significant weakening or even complete shutdown of the AMOC and a strongly reduced northward ocean heat transport [Ganopolski and Rahmstorf, 2001; Van Meerbeeck et al., 2011; Zhang et al., 2014]. It is therefore restricted mostly to the eastern subtropical Atlantic realm, with much less penetration into the continental interiors (Figures 2, 3c, and S2).

84 Black Sea temperature response to glacial millennial-scale climate variability

Figure 3. Comparison of the proxy-based sea surface and model-simulated surface air temperature amplitudes (ΔT) during DO cycles and HE, determined between DO cycles 2–17. (a) Proxy-based ΔT during DO cycles and HE for Greenland [Kindler et al., 2014] (NGRIP), the North Atlantic [Martrat et al., 2007] (MD01-2444), the western Mediterranean Sea [Martrat et al., 2004] (ODP Site 977A), and the southeastern Black Sea (25GC-1, this study). ΔT values for the DO cycles were calculated as the difference between the averaged maximum Greenland Interstadial (GIS) and minimum non-HE Greenland Stadial (GS), and for HE between the averaged minimum HE and Greenland Stadial temperatures (HE-GS). Error bars denote the cumulative mean standard deviation. (b) Model-based and proxy-based (circles) ΔT between GIS and GS. (c) Model-based and proxy-based (circles) ΔT for additional cooling during HE versus non-HE stadials. Temperature patterns from warming during DO cycles and cooling during HE were determined by compiling the modeling results obtained with the CLIMBER-2 model by Ganopolski and Rahmstorf [2001] and with the CCSM3 fully coupled comprehensive general circulation model by Zhang et al. [2014]. Details regarding the experimental design are provided in the supporting information.

The model-based temperature patterns largely resemble the differences in DO and HE patterns observed in the paleodata and likely explain why neither Greenland nor the Black Sea region was affected by significant extra-HE cooling. Paleodata to further test these contrasting atmospheric teleconnections are sparse for regions east of the Black Sea. While the hydroclimatic impact, even though not well captured in models [Kageyama et al., 2013], is more obvious in, e.g., East Asian records [Wang et al., 2001; Sun et al., 2012], adequate paleotemperature data are not yet available.

85 Black Sea temperature response to glacial millennial-scale climate variability

6. Conclusion

The TEX86-based temperature record from the SW Black Sea provides evidence of the impact of Dansgaard-Oeschger cycles on the Eurasian continent. Temperature amplitudes between stadials and interstadials were as high as 4°C, comparable to the values in other midlatitude marine records. In contrast to the North Atlantic and western Mediterranean Sea, cooling was not stronger during HE-equivalent stadials than during regular stadials, which is in line with climate models. The present record confirms a deep Eurasian inland propagation of DO cycle-related temperature changes driven by a close atmospheric teleconnection between the North Atlantic and Eurasia during MIS 3. Over the long term, winters were much milder during the first part of MIS 3. The distinct cooling that started at ~34 ka underlines the importance of both insolation-driven changes in the Eurasian ice volume and the associated atmospheric circulation patterns.

Acknowledgments

The authors thank the captain and crew of the R/V Meteor and the cruise leader, C. Borowski, for their support during the M72/5 Black Sea cruise in 2007. A.W. acknowledges Y. Fagault for assistance during GDGT measurements and O. Körting, K. Häusler, and S. Fink for supporting ostracod sampling. A.W. is indebted to F. Black for constructive contributions during the writing of the manuscript. The authors thank X. Zhang for providing unpublished material for climate modeling. We thank two anonymous reviewers for helpful comments that improved the manuscript and Kim Cobb for editorial handling. This work was funded by the German Science Foundation (AR 367/9-1, 2) within the priority program 1266 INTERDYNAMIC, by the Gary Comer Science and Education Foundation, by the CEREGE institute (Aix-en-Provence), by the Helmholtz Centre Potsdam-German Research Centre for Geosciences (GFZ), and by the Leibniz Institute for Baltic Sea Research Warnemünde (IOW). The data used in the present study are available online at the Data Publisher for Earth and Environmental Science PANGAEA (http://www.pangaea.de). The Editor thanks Eleanor Ivanova and an anonymous reviewer for their assistance in evaluating this paper.

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89 Black Sea temperature response to glacial millennial-scale climate variability

Supporting Information for

Black Sea temperature response to glacial millennial-scale climate variability

Antje Wegwerth1, Andrey Ganopolski2, Guillemette Ménot3, Jérôme Kaiser1,

Olaf Dellwig1, Edouard Bard3, Frank Lamy4, and Helge W. Arz1

1Leibniz-Institute for Baltic Sea Research Warnemünde (IOW), Marine Geology, Rostock, Germany.

2Potsdam Institute for Climate Impact Research (PIK), Earth System Analysis, Potsdam, Germany.

3CEREGE, Aix-Marseille University, Collège de France, CNRS, IRD, Aix en Provence, France.

4Alfred-Wegener-Institut Helmholtz-Zentrum für Polar und Meeresforschung, Bremerhaven, Germany.

Contents of this file

Text S1 to S3 Figures S1 to S3

90 Black Sea temperature response to glacial millennial-scale climate variability

Introduction

The supplementary information provides additional information regarding the contrasting temperature patterns during Dansgaard-Oeschger (DO) cycles and Heinrich events (HE) in the southeastern (SE) and northwestern (NW) Black Sea. We also present additional records confirming DO cycles in the Black Sea region and point out temperature anomalies during individual cooling and warming events in Greenland, the North Atlantic, the Mediterranean, and the SE Black Sea. Finally, the spatial patterns of DO cycles and HE deduced from climate modeling are described in greater detail.

Text S1. Contrasting TEX86 temperature patterns of DO cycles and HE in the NW and SE Black Sea

As seen in Fig. S1, the TEX86-temperature record (from 40 ka BP to 9 ka BP) from the NW Black Sea [Ménot and Bard, 2012] shows cooling during HE but no clear pattern of DO cycles, in contrast to the record from the SE Black Sea (this study). These differing temperature records initially suggest different impacts of the climate variability of the Greenland DO cycles on the NW and SE Black Sea that were probably due to differences in the regional atmospheric circulation. The DO cycle signatures from other records in these regions [Allen et al., 1999; Tzedakis et al., 2004; Fleitmann et al., 2009; Fletcher et al., 2010; Shumilovskikh et al., 2014], however, suggest that these differences in the Black Sea temperatures were not climate-driven.

One reason might be that the NW coring site was subjected to higher terrestrial inputs from the large and heterogeneous NW Black Sea drainage area. Thus, relatively large amounts of isoprenoid GDGTs derived from Thaumarchaeota thriving in soils [Schouten et al., 2013] could have biased the reconstructed temperatures. The Branched and Isoprenoid Tetraether (BIT) index, defined as the normalized (0–1) ratio between terrestrial-derived branched GDGTs and crenarchaeol [Hopmans et al., 2004], is a useful tool for estimating terrestrial organic matter input. A BIT index close to 0 is typical for open marine systems, while a BIT index close to 1 characterizes terrestrial systems [Hopmans et al., 2004]. BIT values from the SE Black Sea (Fig. S1) are <0.3 during the entire investigated period, indicative of generally low inputs of soil-derived isoprenoid GDGTs and providing evidence of the suitability of the TEX86 paleothermometer [Weijers et al., 2006]. BIT indices from the NW region [Ménot and Bard, 2012] are in many cases, >0.3 (Fig. S1), indicating a higher

91 Black Sea temperature response to glacial millennial-scale climate variability

terrigenous contribution in that region. Even though the TEX86 temperatures were BIT- corrected, the temperature record lacks a distinct DO cycle pattern but shows HE-related cooling. A recent study [Liu et al., 2014] suggests that the branched GDGTs were derived from the water column and/or the sediment of the brackish/marine Holocene Black Sea, thus calling into question the application of the BIT as proxy for the input of soil-derived organic matter. Whether this is also the case for the limnic/brackish phases of the Black Sea is unclear.

Further TEX86 temperature reconstructions from other regions of the Black Sea may clarify the different spatial patterns of temperature. Although certain changes during HE are indicated in other regions around the Black Sea, for example, by some but not all palynological data in Greece [Tzedakis et al., 2004] and in Anatolian Lake Van [Çağatay et al., 2014], clear evidence of temperature changes (i.e. HE cooling) is not contained in these records.

92 Black Sea temperature response to glacial millennial-scale climate variability

Figure S1. Climate and environmental conditions in the Black Sea and northern Anatolia during the last glacial. TEX86 mean annual lake surface temperature (LST) for the NW Black Sea (MD042790; gray, BIT-corrected [Ménot and Bard, 2012]) and the SE Black Sea, 25GC- 1 (red, this study). The BIT index for the NW Black Sea (gray, [Ménot and Bard, 2012]) and the SE Black Sea (red, this study) is shown. Crenarchaeol was normalized to TOC and is an indicator of Thaumarchaeota productivity (this study). The sum of all dinoflagellate cysts is a proxy for productivity [Shumilovskikh et al., 2014]. The percentage of arboreal pollen reflects vegetation changes during DO cycles in Northern Anatolia [Shumilovskikh et al., 2014]. δ18O serves as a record of Sofular stalagmites So-1 and So-2 [Fleitmann et al., 2009]. The chronology for 25GC-1 was based primarily on NGRIP (GICC05) and that of MD042790 on Hulu Cave speleothem record [Wang et al., 2001; Ménot and Bard, 2012].

93 Black Sea temperature response to glacial millennial-scale climate variability

Text S2. Temperature anomalies during individual cooling and warming events

The pattern of the temperature anomalies during cooling events (Fig. S2) shows that cooling was generally stronger during HE than during most non-HE stadials at the Iberian Margin and in the Alboran Sea. Note that HE 5a was apparently milder than the other HE, but this was most likely due to an undersampling during this interval.

Figure S2. Temperature anomalies for cooling during the individual HE (HE 2–6) and non- HE stadials (Greenland Stadials GS 4–17) for Greenland ([Kindler et al., 2014], NGRIP); the North Atlantic/Iberian Margin ([Martrat et al., 2007], core MD01-2444), the western Mediterranean Sea/Alboran Sea ([Martrat et al., 2004], ODP Site 977A), and the SE Black Sea (this study). Each temperature anomaly describes the deviation during the strongest cooling from the average temperature throughout the records between 20 and 65 ka BP. The average temperature during the investigated period is -44.9°C for Greenland, 13.9°C for the Iberian Margin, 13.2°C for the Alboran Sea, and 7.0°C for the Black Sea record.

94 Black Sea temperature response to glacial millennial-scale climate variability

By contrast, the temperature anomalies during the HE are generally within the range of those determined for non-HE stadials in Greenland and the Black Sea. The individual temperature departures from the long-term mean (average between 20 and 65 ka BP; Fig. S2) supports the spatial pattern of the proxy-based HE extra cooling, as shown in Fig. 3a, c in the main text. The GIS temperature anomalies (Fig. S3) show that warming was slightly stronger during the GIS succeeding the HE. However, in contrast to the GS temperature anomalies, this difference compared to non-HE associated GIS warming is minor.

Figure S3. Temperature anomalies for warming during the individual Greenland Interstadials (GIS) following HE (diamonds) and during GIS following non-HE stadials (circles) for Greenland ([Kindler et al., 2014], NGRIP); the North Atlantic/Iberian Margin ([Martrat et al., 2007], core MD01-2444), the western Mediterranean Sea/Alboran Sea ([Martrat et al., 2004], ODP Site 977A), and the SE Black Sea (this study). Each temperature anomaly describes the deviation during the strongest warming from the average temperature throughout the records (between 20 and 65 ka BP). The average temperature during the investigated period is - 44.9°C for Greenland, 13.9°C for the Iberian Margin, 13.2°C for the Alboran Sea, and 7.0°C for the Black Sea record.

95 Black Sea temperature response to glacial millennial-scale climate variability

Text S3. Spatial patterns of DO cycles and HE deduced from climate modeling

The difference between the spatial patterns of the temperature anomalies during DO cycles and HE was previously explained by the existence of three distinct modes of the glacial Atlantic Meridional Overturning Circulation (AMOC) [Ganopolski and Rahmstorf, 2001]. According to Ganopolski and Rahmstorf [2001], the DO cycle pattern corresponds to the difference between the interstadial (similar to the present) and non-HE stadial modes of the AMOC, which is weaker than the interstadial mode but is also characterized by a more southward location of the deep-water formation area in the North Atlantic and a more widely expanded sea-ice cover in Nordic Seas. The significant reduction of the sea-ice cover during the transition from a non-HE stadial to an interstadial mode of the AMOC strongly amplifies the direct temperature response to the increased meridional ocean heat transport, especially during the winter. This concept is additionally supported by the results of Li et al. [2005], which showed (Fig. 1B) that the temperature effect of sea-ice removal from the Nordic Seas is similar to the DO pattern. The HE temperature anomaly pattern represents additional cooling from the non-HE stadial to the HE stadial. This cooling can be primarily attributed to a significant reduction of the AMOC during HE, a conclusion supported by the larger magnitude of the bipolar seesaw pattern during HE stadials [Margari et al., 2010]. The results reported by Ganopolski and Rahmstorf [2001], who used a coarse-resolution model of intermediate complexity (CLIMBER-2), were recently corroborated by those of modeling experiments obtained with the high-resolution fully coupled general circulation model CCSM3 of Zhang et al. [2014]. Figure 1b in that paper shows annual mean temperature differences between experiments of “+0.02 Sv” and “−0.02 Sv”, which can be interpreted as the difference between non-HE stadial and interstadial modes. The results are in qualitatively good agreement with the warming pattern of the DO shown in Fig. 3a of Ganopolski and Rahmstorf [2001]. In particular, simulations with the CCSM3 model showed maximum temperature anomalies (>10°C) over the Nordic Seas as well as a large change in the maximum sea-ice extent in the North Atlantic. The temperature differences between experiments “+0.02 Sv” and “+0.2 Sv” (Xiao Zhang, personal communication) may reflect the difference between HE stadials (“off” mode of the AMOC) and non-HE stadials. Again, the pattern of additional HE cooling is in agreement with the CLIMBER-2 model simulations shown in Fig. 3c [Ganopolski and Rahmstorf, 2001]. In spite of the strong additional reduction in AMOC strength determined in this experiment, the magnitude of additional cooling is significantly smaller than the differences between experiments “+0.02 Sv” and

96 Black Sea temperature response to glacial millennial-scale climate variability

“−0.02 Sv”. Moreover, the maximum of the temperature anomalies associated with the AMOC shutdown is located south of Iceland. As described by Ganopolski and Rahmstorf [2001], a complete shutdown of the AMOC in the CCSM3 model causes only modest additional cooling in Greenland and has a very small effect on the maximum sea-ice extent. It is, however, important to note that, although the annual mean DO signal is stronger than the additional HE cooling, it is strongest during the cold season whereas additional HE cooling is less seasonally dependent. In addition, HE cooling occurred closer to the Equator and therefore should have had a stronger impact on the position of the Intertropical Convergence Zone, which is crucial for the strength of the Southeast Asian summer monsoon. This can explain why in the monsoon proxies HE stadials differed from non-Heinrich stadials [Wang et al., 2001].

Considering the slight differences between model-based and proxy-based temperature changes during HE cooling (Fig. 3c), the temperature reconstructions (apart from those in Greenland) may be biased towards summer, whilst the response to AMOC changes is stronger during winter. This can explain why annual mean model temperature changes are greater than those induced from data.

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99 Northern hemisphere climate control on the environmental dynamics in the glacial Black Sea “Lake”

Northern hemisphere climate control on the environmental dynamics in the glacial Black Sea “Lake”

Antje Wegwertha*, Jérôme Kaisera, Olaf Dellwiga, Lyudmila S. Shumilovskikhb,c,

Norbert R. Nowaczykd, and Helge W. Arza

aLeibniz-Institute for Baltic Sea Research Warnemünde (IOW), Marine Geology, Rostock,

Germany bDepartment of Palynology and Climate Dynamics, University of Göttingen, Göttingen,

Germany cLaboratory of biogeochemical and remote techniques of environmental monitoring, National

Research Tomsk State University, Tomsk, Russia dHelmholtz Centre Potsdam, GFZ German Research Centre for Geosciences, Section Climate

Dynamics and Landscape evolution, Potsdam, Germany

under review for publication in “Quaternary Science Reviews”

100 Northern hemisphere climate control on the environmental dynamics in the glacial Black Sea “Lake”

Abstract

The climate variations during the last glacial occurring on multi-millennial and millennial timescales propagated over large parts of the Northern Hemisphere through atmospheric teleconnections. Over the long-term, the Marine Isotope Stages (MIS) 2 and 4 were colder, drier, and windier than the MIS 3. The MIS 3 stands out due to its abrupt changes from cold and dry stadials to warm and humid interstadials, the so-called Dansgaard- Oeschger cycles that also affected temperature and rainfall in the Black Sea region. This study is based on a gravity core from the southeastern (SE) Black Sea that covers the last glacial lake stage from 64 to 20 ka BP. By using the composition of major and trace elements in the sediments, terrestrial plant-derived n-alkane flux, and Sr/Ca from benthic ostracods, we reconstruct the variability of riverine and aeolian input, salinity, and productivity in the SE Black Sea region in response to the Northern Hemisphere climate oscillations. During colder and drier stadials, the aeolian input increased relative to the riverine discharge, potentially due to southward shifted and/or stronger westerly winds and due to changes in the vegetation cover. An evaporation exceeding freshwater supply by rainfall and rivers possibly caused higher salinity and a lower lake level. The environmental status during MIS 4 and 2 is very much comparable with the stadial conditions during MIS 3. During warmer and more humid interstadials, lower salinity and presumably positive lake level changes most likely resulted from increased precipitation and river discharge. This likely increased primary productivity through an augmented nutrient supply. Lowest average salinities are suggested for the middle part of MIS 3 in response to enhanced meltwater from the disintegrating Fennoscandian Ice Sheet and/or by generally more humid conditions.

Keywords: Black Sea “Lake”, Dansgaard-Oeschger stadials and interstadials, riverine and aeolian input, salinity, productivity

Highlights • Black Sea record between 64 and 20 ka BP is presented.

• Main parameters are K/Ti, Ti/Al, Zr/Al, n-C27 - n-C31 n-alkanes, and Sr/Caostracods. • Higher aeolian input and salinity and low productivity during stadials. • Higher riverine input and productivity and low salinity during interstadials. • Long-term pattern shows freshening during MIS 3.

101 Northern hemisphere climate control on the environmental dynamics in the glacial Black Sea “Lake”

1. Introduction

The Marine Isotope Stage 3 (MIS 3) was characterised by pronounced millennial-scale climate fluctuations characterised by abrupt shifts from cold stadials to warmer interstadials in the northern high latitudes (Dansgaard et al., 1993). These so-called Dansgaard-Oeschger cycles (DO) were originally detected in Greenland ice cores by pronounced oxygen isotope variations occurring on a centennial to millennial timescale (Dansgaard et al., 1993). Meanwhile, proxy data and climate models provide evidence for large temperature variations with amplitudes of 6-16.5°C during DO in Greenland and northern latitudes of continental Europe (Ganopolski and Rahmstorf, 2001; Johnsen et al., 2001; Huber et al., 2006; Kindler et al., 2014). In the mid-latitude North Atlantic and Mediterranean regions temperature changes during DO were smaller but amplitudes amounted on average 2°C (e.g. Martrat et al., 2004, 2007). A similar level of DO associated temperature changes was observed also for the continental interior from the formerly enclosed Black Sea “Lake” (Wegwerth et al., 2015). Interstadial warming and stadial cooling most likely resulted from a stronger and weaker Atlantic Meridional Overturning circulation (AMOC) triggered by major changes in freshwater supply from ice sheets (e.g. Ganopolski and Rahmstorf, 2001). Several stadials, referred to as Heinrich events (HE) mark the coldest periods in certain regions across the northern hemisphere (Ruddiman, 1977; Heinrich, 1988; Bond et al., 1993; Bard et al., 2000). The exceptional cooling during HE was most likely a result from an off-mode of the AMOC and an associated reduced ocean heat transport (Ganopolski and Rahmstorf, 2001). The impact of HE seems most pronounced in the mid-latitude North Atlantic (Bard et al., 2000; Ganopolski and Rahmstorf, 2001; Martrat et al., 2007).

Studies currently available investigating the climate evolution during MIS 3 in continental Eastern Europe, apart from the Black Sea, are from the Carpathians (Stevens et al., 2011; Sümegi et al., 2013), Ukraine (Rousseau et al., 2011; Sima et al., 2013), Anatolia (Fleitmann et al., 2009; Rowe et al., 2012), Greece (Tzedakis et al., 2004; Margari et al., 2009), Albania/Macedonia (Vogel et al., 2010), and the Aegean Sea (Geraga et al., 2005; Ehrmann et al., 2013). These studies are in general agreement that DO interstadials were characterised by mild and humid climate conditions, favouring growth of temperate forests and soil development. In contrast, DO stadials were much colder and more arid. Together with less dense vegetation cover, pronounced westerly winds allowed the formation of large European loess deposits during stadials (e.g. Rousseau et al., 2011; Stevens et al., 2011; Sima et al., 2013; Sümegi et al., 2013). The periods of enhanced loess deposition in Europe are

102 Northern hemisphere climate control on the environmental dynamics in the glacial Black Sea “Lake” about the same as the well-known Chinese loess deposits (e.g. Sun et al., 2012) and suggest a common atmospheric forcing.

The Black Sea with its sedimentary archives represents a key location for palaeoclimate reconstructions bridging the North Atlantic and Eurasian climate realms. Over the last 670 ka, the Black Sea was mostly a (freshwater) lake isolated from the Mediterranean Sea during glacials (e.g. Schrader, 1979; Badertscher et al., 2011). Previous studies focussed extensively on the last two glacial terminations as well as on the Holocene and Eemian climatic and environmental changes in the Black Sea (e.g. Ryan et al., 2003; Lamy et al., 2006; Major et al., 2006; Bahr et al., 2008; Kwiecien et al., 2009; Soulet et al., 2011; Shumilovskikh et al., 2012, 2013a, 2013b; Wegwerth et al., 2014). These studies revealed a generally strong response of the Black Sea to global and hemisphere-wide climate. Nevertheless, not much is known about the impact of the pronounced climate fluctuations during Marine Isotope Stage 3 (MIS 3, the period between 57 and 29 ka BP; (Heinrich, 1988; Bond et al., 1993; Dansgaard et al., 1993; Bard et al., 2000). In the northwestern part of the Black Sea paleoenvironmental data were published by Ménot and Bard (2012), Rostek and Bard (2013), and Sanchi et al. (2014, 2015) covering the last glacial back to about 40 ka BP. Nowaczyk et al. (2012, 2013), Shumilovskikh et al. (2014), and Wegwerth et al. (2015) presented first palaeomagnetic, palynologic, and temperature reconstructions from the south-eastern part of the basin extending back to about 64 ka BP.

Here, we document the environmental response of the SE Black Sea to last glacial millennial-scale climate variability. By using a combination of the detrital element composition (Al, K, Ti, Zr etc.), the terrestrial plant derived n-alkane flux, and ostracod geochemistry (Sr/Ca) data with previously published data sets including total inorganic carbon, coastal ice-rafted detritus, pollen and dinoflagellate records from the same sediment core 25GC-1 (Nowaczyk et al., 2012; Shumilovskikh et al., 2014; Wegwerth et al., 2015), we reconstruct changes in riverine and aeolian input, salinity, and productivity. The results are then set into an over-regional context and linked to major atmospheric changes across Eurasia during the last glacial.

2. Climate and environmental setting of the Black Sea region

Presently, the Black Sea region is under Mediterranean influence resulting in dry summers and humid winters (Cullen and deMenocal, 2000; Göktürk et al., 2011).

103 Northern hemisphere climate control on the environmental dynamics in the glacial Black Sea “Lake”

The winter climate in the southern Black Sea and northern Anatolia is influenced by the polar front, the Siberian anticyclone, cold air masses from the Balkan Peninsula, Mediterranean cyclones, and northwesterly winds (Fig. 1; Wigley and Farmer, 1982; Tolmazin, 1985, Akçar and Schlüchter, 2005; Kwiecien et al., 2009; Deniz et al., 2011), when wind intensities are strongest (Staneva and Stanev, 1998).

Figure 1: Idealised scheme of the modern atmospheric circulation patterns with broad position of the polar front during January and the Subtropical Jet during July influencing regional climate in the Black Sea region (modified after Wigley and Farmer, 1982; Akçar and Schlüchter, 2005). Ice sheet limits during the MIS 3 and LGM are from Larsen et al. (2006). The red circle denotes the location of the sediment core 25GC-1 and numbers denote the outlets of major rivers draining into the Black Sea. The rivers are listed in the lower right inlay. Yellow circles denote sites discussed in the text: MD95-2043 (western Mediterranean Sea; Moreno et al., 2005), Co-1202 (Lake Ohrid; Vogel et al., 2010), So-1/2 (Sofular Cave; Fleitmann et al., 2009), and Başyayla Valley (NE Anatolia; Reber et al., 2014). The map was created with the generic mapping tool provided by woodshole.er.usgs.gov/mapit/index.html.

During summer the northward shifted polar front gives place to the Azores high- pressure system and the subtropical Jet causing maritime conditions with predominantly southwestern and southern winds (Fig. 1; Tolmazin, 1985, Deniz et al., 2011). NE Anatolia and the eastern Black Sea are the most summer rain-laden regions due to orographic precipitation in the Pontic and Caucasian Mountains (Staneva and Stanev, 1998; Türkeş et al., 2009; Deniz et al., 2011), whereas the Black Sea itself forms an additional moisture source (Fleitmann et al., 2009; Badertscher et al., 2011; Göktürk et al., 2011). Mean annual air temperatures are about +10°C in the northern and +14 to +15°C in the southern Black Sea

104 Northern hemisphere climate control on the environmental dynamics in the glacial Black Sea “Lake” region (Kosarev et al., 2008). The spatial variability is most pronounced during winter, when mean February temperatures are -2°C in the north and +7.5°C in the south (Kosarev et al., 2008). In August, mean temperatures are +21.5°C in the northern and +24°C in the southern basin (Kosarev et al., 2008).

The modern climate (i.e. air and sea surface temperatures, precipitation patterns) in that region is strongly related to the North Atlantic Oscillation (NAO). The NAO is expressed as the difference between the subtropical Azores high-pressure system and the subpolar Icelandic low-pressure system, which influence is most important for the North Atlantic Ocean and its marginal seas (Oguz et al., 2006). Positive winter NAO causes strong westerly winds carrying cold and dry air to the Black Sea, while negative NAO phases result in milder winters and more humid conditions (Oguz et al., 2006).

The shelf and catchment areas as well as the riverine input are much larger in the NW Black Sea when compared to the SE Black Sea. The Danube River in the NW contributes to about 50-70% of the total riverine input (Fig. 1; Özsoy and Ünlüata, 1997; Nezlin, 2008). Further major rivers in this region are the Dnieper and the Dniester, whose discharge is about one third of that of the Danube (Özsoy and Ünlüata, 1997). The freshwater supply from all remaining rivers is < 20% (Özsoy and Ünlüata, 1997).

3. Material and methods

This study is based on a 952 cm long gravity core 25GC-1 recovered from the Archangelsky Ridge (42°06.2’N, 36°37.4’ E; 418 m water depth) in the SE Black Sea during R/V Meteor cruise M72/5 in 2007 (Fig. 1; Nowaczyk et al., 2012; Shumilovskikh et al., 2014; Wegwerth et al., 2015). Core 25GC-1 contains a well-preserved sequence (952-307 cm core depth) from the late MIS 4 to MIS 2 and consists of detrital mud and intercalated carbonate- rich intervals (Nowaczyk et al., 2012; Nowaczyk et al., 2013; Shumilovskikh et al., 2014; Wegwerth et al., 2015).

3.1 Age-depth model

Compared to the western Black Sea, sedimentation on the Archangelsky Ridge is comparatively low but favours the recovery of older glacial and previous interglacial periods from shallower sub-bottom depths (e.g. Shumilovskikh et al., 2013a, 2013b, 2014; Wegwerth et al., 2014, 2015).

105 Northern hemisphere climate control on the environmental dynamics in the glacial Black Sea “Lake”

The age-depth model of core 25GC-1 (Fig. 2) has been determined by Nowaczyk et al. (2012). For the investigated period, it is based on eight AMS (accelerator mass spectrometry) radiocarbon datings on juvenile shells of the bivalve Dreissena rostriformis from Black Sea core 24GC-3 (208 m water depth), reaching back to 39 ka BP, the Campanian Ignimbrite tephra (Y5; 39.28 ± 0.11 ka; De Vivo et al., 2001), and the Laschamp geomagnetic excursion (40.70 ± 0.95 ka BP; Singer et al., 2009). The chronology of the core was further improved with the related palaeomagnetic records, including ChRM (characteristic remanent magnetisation) inclination and declination as well as relative palaeointensity (Nowaczyk et al. 2012). For the final chronology between 63.4 ka BP and 20 ka BP several sedimentary parameters comprising coastal ice-rafted detritus (IRDC), carbonate content, Ca-XRF (X-ray fluorescence) counts, and high-resolution magnetic susceptibility of the core were tuned to the oxygen isotope record from the North Greenland Ice core Project on the GICC05 timescale (NGRIP; Svensson et al., 2006, 2008).

Figure 2: Age-depth model of sediment core 25GC-1 between 20 and 65 ka BP (after Nowaczyk et al., 2012).

Tuning was done with the xtc-program (extended tool for correlation) allowing an interactive wiggle-matching routine (Nowaczyk et al.; 2012). Cullen et al. (2014) recently published the detailed tephrostratigraphic framework from core 25GC-1 where they identified, apart from the Campanian Ignimbrite tephra, 21 additional non-visible cryptotephra horizons largely confirming our NGRIP-tuned age model.

3.2 Sediment geochemistry

Geochemical analyses were performed on aliquots of freeze-dried and homogenised sediment samples. Total carbon (TC) and total organic carbon (TOC) were determined with a

106 Northern hemisphere climate control on the environmental dynamics in the glacial Black Sea “Lake”

EuroVector CHNS Elementar analyser (Nowaczyk et al., 2012) and TIC was then calculated by the difference between TC and TOC. For the determination of major and trace element contents, 50 mg of sediment material were digested with a HF/HClO4 mixture (Schnetger, 1997) in closed Teflon vessels (PDS-6, Heinrichs et al., 1986) at 180°C. After evaporation of the acids on a metal plate at 180°C, the digestions were fumed off 3-times with half- concentrated HCl and finally diluted with 2 Vol.% HNO3 to 25 mL. Al, K, Ti, and Zr were measured with ICP-OES (iCAP 6300 Duo, Thermo Fisher Scientific) and Re contents were obtained by ICP-MS (iCAP Q, Thermo Fisher Scientific) in CCT mode. 187Re data were corrected for interference by 171Yb16O+ by analysing a single solution containing only Yb. Precision and accuracy of ICP-OES (<4.2% and <6.6%) were determined by the international reference material SGR-1 (USGS) while PACS-1 (NRC Canada) was used for ICP-MS (6.3% and 2.8%). High-resolution core scanning (av. 1 cm) of K and Ti was carried out by X-ray fluorescence (XRF) with the Avaatech XRF core scanner configured with an Amptek detector (Nowaczyk et al., 2012). In the following, the different methods for elemental compositions

(X) are symbolised by XOES, XICP-MS, and XXRF. K/TiXRF ratios are shown as ln-ratios (natural logarithm) as recommended by Weltje and Tjallingii (2008).

3.3 Ostracod abundance and geochemistry

The ostracod valves of adult Candona spp. were handpicked and counted under a binocular from the dried, wet-sieved sediment samples (>150 µm, with an original wet volume of 10 cm3). The valves were cleaned with a thin brush and few drops of deionised water. Up to five valves were measured for Sr/Ca ratios by ICP-MS (iCAP Q; ThermoFisher

Scientific) after dissolution in 2 vol.% HNO3 (sub-boiled) and centrifugation. Compensation of matrix effects was achieved by using Be and Rh as internal standards. Al was measured for monitoring potential contamination by detrital material. An international reference material (ECRM 752-1, BAS) was measured every five samples and assured a precision of 2.2% and an accuracy of 1.4%.

3.4 Organic geochemistry

The preparation for n-alkane analysis involved accelerated solvent extraction (Thermo Scientific™ Dionex™ ASE™ 350) of freeze-dried and homogenised samples (4-6 g of sediment) with a dichloromethane/methanol mixture (DCM/MeOH 9:1), followed by

107 Northern hemisphere climate control on the environmental dynamics in the glacial Black Sea “Lake” desulphurisation with activated copper pieces and addition of an internal standard containing a mixture of squalane, 2-nonadecanone and androstanol. The separation of the apolar fraction containing n-alkanes has been done with hexane/DCM (9:1) over a Pasteur pipette plugged with activated Al2O3. The fractions were measured on a multichannel TraceUltra gas chromatograph (ThermoScientific) equipped with a DB-5 MS 30 m x 0.32 mm x 0.25 µm capillary column. Nitrogen was used as carrier gas and the temperature ramp was programmed from 40 °C to 290 °C at 4° C/min followed by a 20 min plateau. Peak identification was based on the comparison of peak retention times with an external standard containing n-C8 to n-C40 alkanes. The reproducibility of the n-alkane quantification was better than 15%.

4. Results

4.1 Sediment geochemistry

The contents of TOC and TIC are high during interstadials and low during stadials and follow the DO-pattern of the δ18O from the Greenland ice core NGRIP (Svennson et al., 2008) during MIS 3 (Figs. 3A-C). While TOC and TIC increase during interstadials in the late MIS 4, the contents remain low during interstadials 3 and 4 in the early MIS 2. One exception is the stadial DO 13 and 14, where TOC remains relatively high.

To eliminate dilution effects by carbonate, Ti and Zr contents are normalised to Al. 2 Ti/AlOES and Zr/AlOES ratios reveal a similar pattern (correlation coefficient R =0.69) with higher values during stadials (Figs. 3D, E). The period between DO 14 and 12 reveals most enhanced Ti/AlOES and Zr/AlOES ratios. The ln(K/TiXRF) ratios is inversely correlated to 2 Ti/AlOES and Zr/AlOES (correlation coefficient R =-0.65) with elevated values especially during DO interstadials 17-5. The pattern of the redox-sensitive Re normalised to Al

(Re/AlICP-MS) reveals maximum values during interstadials 14, 12, and 11.

108 Northern hemisphere climate control on the environmental dynamics in the glacial Black Sea “Lake”

Figure 3: Temporal variation of A) δ18O NGRIP (Greenland; Svennson et al., 2008) B) total organic carbon (TOC), C) total inorganic carbon (TIC), D) Ti/AlOES, E) Zr/AlOES, F) ln (K/TiXRF), G) Re/AlICP-MS, H) abundance from ostracod valves, I) Sr/Ca from ostracod carbonate shells, and J) sum of n-C27 to n-C31 odd-numbered n-alkanes. The red numbers and grey bars denote the interstadials during Dansgaard-Oeschger (DO) cycles and the blue numbers denote Heinrich event (HE) equivalent periods.

109 Northern hemisphere climate control on the environmental dynamics in the glacial Black Sea “Lake”

4.2 Ostracod abundance and geochemistry

Ostracods, mainly dominated by Candona spp., were widely distributed during the glacial stages of the limnic/brackish Black Sea (e.g. Bahr et al., 2008; Boomer et al., 2010; Wegwerth et al., 2014). In our record, the abundance of ostracod valves tends to increase during stadials and it reaches a maximum of up to 150 specimens/10 cm-3 during MIS 2 (Fig.

3H). Sr/Caostracods values show a long-term trend towards a minimum level during mid-MIS 3

(Fig. 3I). MIS 4 and 2 in turn have overall higher Sr/Caostracods values. Although there is no clear DO pattern, Sr/Caostracods values tend to be lower (higher) during interstadials (stadials).

4.3 Organic geochemistry

The sum of n-C27 to n-C31 odd-numbered n-alkanes represents the amount of leaf wax lipids from higher vascular plants (e.g. Eglinton and Eglinton, 2008). The sum of odd- -1 numbered long chain alkanes (n-C27+n-C29+n-C31) ranges between 200 and 2200 ng g (Fig. 3J) showing elevated values during stadials. The sum of n-alkanes is higher after ca. 43 ka

BP. An estimation of the maturity of n-alkanes can be obtained with the CPI index (n-C27 to n-C33 odd-over-even carbon-number preference index) (Bray and Evans, 1961; Marzi et al.,

1993). The relatively high mean CPI27-33 value of 6.3 (n = 163; average standard deviation = 1.2), are very similar to CPI values from the NW Black Sea (CPI = 6.6; Rostek and Bard, 2013) and suggest that the n-alkanes are largely unaffected by degradation and diagenesis and applicable for environmental reconstructions.

5. Discussion

5.1 Variability in the detrital sedimentation

18 The ln(K/TiXRF) ratio varies on a millennial timescale in concert with the δ ONGRIP record (Fig. 4 A, D). This implies an increasing contribution of K-rich minerals during the milder DO interstadials at least until ~32 ka BP. K is mainly originating from weathered K- feldspars and mica and incorporated into clay minerals (illite) that are transported under low energy conditions (Müller and Stoffers, 1974, Piper and Calvert, 2011). In the modern Black Sea, K mainly originates from detrital suspended matter of the northern rivers Danube, Dniester, and Dnieper (Fig. 1; Piper and Calvert, 2011). A δ13C record of the Sofular stalagmites originating from northwestern Anatolia (Fig. 4 C) documents moister and milder

110 Northern hemisphere climate control on the environmental dynamics in the glacial Black Sea “Lake” climate during interstadials, because the considerably decreasing values point to changes from

C4 plant (grasses) to C3 plant (trees, shrubs) dominance above the cave (Fleitmann et al., 2009). Similarly, a pollen record from our sediment core 25GC-1 (Fig. 5 C) shows increased temperate forest biomes during interstadials which suggests more humid conditions in northeastern Anatolia (Shumilovskikh et al., 2014). The high ln(K/TiXRF) values in our record (Fig. 4 D) may reflect, therefore, enhanced local river supply in response to more humid interstadial conditions. Arslanov et al. (2007) documented also pronounced interstadial freshwater input in the eastern Black Sea from a pollen record at the Dziguta River (Caucasian region) since 50 ka BP that is in agreement with the SE Black Sea record. Such strong MIS 3 variability in the fluvial successions was also determined in southern Europe by a comparison of fluvial successions and global and regional climate modelling (van Huissteden and Pollard, 2003).

The freshwater supply to the Black Sea was possibly amplified by melting of glaciers in the northern Anatolian mountains but also of the Fennoscandian Ice Sheet. During warming of Termination I and II, relatively large amounts of K-bearing material entered the Black Sea “Lake” due to melting of the Fennoscandian Ice Sheet with Black Sea meltwater events lasting about 220-1000 years (Soulet et al., 2013; Wegwerth et al., 2014). In the present study high ln(K/TiXRF) values suggest largest fluvial input between 60 and 45 ka BP (Fig. 4 D) that correspond to the maximum in summer insolation at 65°N (Fig. 4 B; Laskar et al., 2004) and the minimum volume of the Fennoscandian Ice Sheet (Fig. 5 D; Bintanja and van de Wal, 2008). We assume that a considerable amount of freshwater reached at that time the Black Sea from the north. Moreover, a climate and insolation forced ice sheet response could have driven the fluvial meltwater input also on a millennial time scale. Arz et al. (2007) indeed proposed by means of model simulations and proxy-based sea level reconstructions pronounced variations in the mass balance of northern hemisphere ice sheets between stadials and interstadials, especially at the southern continent-based margins of the Fennoscandian and Laurentide ice sheets, because of the pronounced magnitude of the temperature changes during DO cycles. It is therefore likely that the increased freshwater input during DO interstadials was a combination of higher rainfall and decaying northern ice sheets.

Periods with decreased ln(K/TiXRF) ratios indicate less riverine input during stadials synchronous to increased Ti/AlOES and Zr/AlOES (Figs. 3 D, 4 E). Ti and Zr are associated with heavy minerals as e.g. rutile, ilmenite, and zircon, generally enriched in sediments

111 Northern hemisphere climate control on the environmental dynamics in the glacial Black Sea “Lake” formed under high depositional energy and aeolian transport (Schnetger, 1992; Dellwig et al., 2000; Calvert and Pedersen, 2007; Piper and Calvert, 2011; Bout-Roumazeilles et al., 2013).

Figure 4: Temporal variability of terrestrial input contributed by riverine and aeolian sources. 18 A) δ ONGRIP (Greenland; Svennson et al., 2008), B) summer (June-July-August) insolation at 13 65°N (Laskar et al., 2004), C) δ CSofular Cave (stalagmites So-1 and So-2; Fleitmann et al., 2009), D) ln(K/TiXRF) (core 25GC-1), E) Zr/AlOES (core 25GC-1), F) sum of n-C27 to n-C31 odd-numbered n-alkanes (core 25GC-1), G) xerophytic steppe biome (core 25GC-1; Shumilovskikh et al., 2014), H) ln(Zr/TiXRF) from Lake Ohrid (core Co-1202; Vogel et al., 2010; triangles denote age control points), I) Si/(Si+K) from the western Mediterranean Sea (core MD95-2043; Moreno et al., 2005; triangles denote age control points), and J) northeastern Anatolian glacier advance (Başyayla Valley; Reber et al., 2014). The red numbers and grey bars denote the interstadials during Dansgaard-Oeschger (DO) cycles and the blue numbers denote Heinrich event (HE) equivalent periods.

112 Northern hemisphere climate control on the environmental dynamics in the glacial Black Sea “Lake”

Due to the location of the core 25GC-1 on the deeper part of the Archangelsky Ridge and the fact that the Kizilirmak plume is rather diverted westward by a local anticyclonic gyre (Oguz et al., 1993), the records potentially reflect the aeolian contribution rather than a sorting by bottom currents.

There are several studies documenting increased aeolian deposition in eastern and southeastern Europe during cold and dry periods throughout the Quaternary. Buggle et al. (2008) presented a geochemical characterisation of eastern and southeastern loess deposits formed during glacial and stadial stages of the Quaternary, and determined high amounts of Ti and Zr particularly within the Dnieper loess area. Loess accumulation during dry and windy stadials alternated with soil formation during moister and less windy interstadials with a denser vegetation cover in Eastern Europe (Rousseau et al., 2011; Sima et al., 2013). Up to two-fold higher stadial dust emissions were simulated for the Ukrainian/East European loess deposits (Sima et al., 2013). A study from the eastern Mediterranean Sea has further shown that Ti/Al and Zr/Al strongly increased during the arid and windy glacials on glacial- interglacial timescales (Wehausen and Brumsack, 2000). A Zr/TiXRF record from the nearby Lake Ohrid (Figs. 1, 4 H; Vogel et al., 2010) and a Si/(Si+K) record from the western Mediterranean Sea (Figs., 1; 4 I; Moreno et al., 2005) indicate increased erosion, aeolian input, wind activity, and aridity during some stadials, which is in broad accordance with the Black Sea “Lake” record.

A significant aeolian contribution to the central Black Sea has been deduced from n- alkanes in recent to sub-recent sediments of the central Black Sea way beyond the shelf areas (Wakeham, 1996). This evidence and the fact that the n-alkane record from core 25GC-1 calculated as the sum of n-C27 to n-C31 alkanes (Fig. 4 F) presents a similar pattern as the

Zr/AlOES record (Fig. 4 E) suggests the predominant aeolian transport of plant-derived n- alkanes. Thus, a less dense stadial vegetation cover dominated by xerophytic steppe biome which was favoured by dry conditions (Fig. 4 G; Shumilovskikh et al., 2014), and stronger winds (Sima et al., 2013) likely triggered deflation and transport of leaf waxes (Simoneit et al., 1977; Kusch et al., 2010). Less dense vegetation during stadials was also determined by palynological investigation from Greece (Tzedakis et al., 2004; Tzedakis, 2005; Müller et al., 2011). Multiple glacier advances since 57 ± 3.5 ka BP in NE Anatolia (Figs. 1; 5 I; Reber et al., 2014), could have additionally favoured deflation in that region. Their temporal pattern, however, does not allow linking glacier advance with millennial-scale climate variability but may suggest that main glacier advances took place during most HE.

113 Northern hemisphere climate control on the environmental dynamics in the glacial Black Sea “Lake”

There is evidence for enhanced westerly wind strength during stadials (e.g. Sima et al. 2013), which was possibly the main carrier for aeolian material into the Black Sea. Future provenance studies are required to understand its pathways and source regions.

5.2 Changes in the basins hydrology

The Sr/Ca ratio of benthic ostracods is a proxy that has been successfully used for the reconstruction of relative salinity changes during glacial stages in the Black Sea “Lake” (Bahr et al., 2008; Wegwerth et al., 2014). Due to moulting (ecdysis), ostracods build up carbonate shells up to nine times within their life cycle (e.g. Chivas, 1986). The higher the Sr concentrations in the host water are, the higher the uptake of Sr into the carbonate shell is (e.g. Chivas et al., 1986; Engstrom and Nelson, 1991; Wansard et. al., 1998). As ocean water generally contains higher concentrations of dissolved Sr when compared with average river -1 -1 water (8,000 µg l vs. 60 µg l ; Martin and Whitfield, 1983), Sr/Caostracods values can reflect past salinity changes (e.g. Chivas et al., 1986; Wansard et. al., 1998).

-1 The Sr/Caostracods long-term pattern in the Black Sea shows 0.2-0.3 mmol mol lower values during mid-MIS 3 when compared to MIS 4 and 2 (Fig. 5A). Although an appropriate calibration for a salinity reconstruction is not available for the limnic Black Sea, the

Sr/Caostracods pattern suggests lower salinity during MIS 3 when compared to MIS 4 and 2.

Since MIS 3 Sr/Caostracods values are close to the values measured for the meltwater events of the penultimate glacial in the SE Black Sea (Wegwerth et al., 2014) they suggest considerable meltwater contribution from the disintegrating Fennoscandian Ice Sheet (Fig. 5 D; Bintanja and van de Wal, 2008). This might have reached the Black Sea via the rivers Dnieper and Danube during that period. The decreasing abundance of marine indicators and increasing concentrations of freshwater algae (Shumilovskikh et al., 2014) in the SE Black Sea during early MIS 3 agrees well with the freshening suggested from the Sr/Caostracods record.

Higher fluvial input (Fig. 4 D) during MIS 3 also suggests an increased rainfall at least in the hinterland of the investigated sediment core and presumably, reduced evaporation over the whole Black Sea. This may explain the observed low Sr/Caostracods values. Support for this assumption comes from the carbon isotope records from north-western Anatolian stalagmites (Figs. 1, 5 B; Fleitmann et al., 2009), the record of temperate forest biome from the SE Black Sea (Fig. 5C; Shumilovskikh et al., 2014), and loess sequences from the Danube basin (Fitzsimmons et al., 2012), which imply wetter and milder conditions during MIS 3 when compared to MIS 4 and 2.

114 Northern hemisphere climate control on the environmental dynamics in the glacial Black Sea “Lake”

Figure 5: Temporal variations of A) Sr/Caostracods (core 25GC-1; raw data and simple 21-point 13 moving average), B) δ CSofular Cave (Fleitmann et al., 2009), C) temperate forest biome (core 25GC-1; raw data and 21-point moving average; Shumilovskikh et al., 2014), D) Eurasian ice volume (Bintanja and van de Wal, 2008), and E) accumulation rate of ice-rafted detritus (IRDC; core 25GC-1; modified from Nowaczyk et al., 2012; Wegwerth et al., 2015). The red numbers and grey bars denote the interstadials during Dansgaard-Oeschger (DO) cycles and the blue numbers denote Heinrich event (HE) equivalent periods.

Accumulation rates of coastal ice-rafted detritus (IRDC) from the SE Black Sea (Fig. 5 E), which most likely originates from winter coastal ice formation (Nowazcyk et al., 2012), are much lower during the early MIS 3 when compared with MIS 4, late MIS 3 and early MIS 2. This pattern reflects milder winter conditions in the Black Sea related to the retreat of the Eurasian ice sheet (Fig. 5D; Bintanja and van de Wal, 2008) and northward shifted atmospheric polar fronts during a phase of maximum summer insolation (Laskar et al., 2004).

115 Northern hemisphere climate control on the environmental dynamics in the glacial Black Sea “Lake”

Winter conditions similar to the modern ones might have prevailed during early MIS 3 with frequent cyclogenesis over the eastern Mediterranean Sea that increased rainfall along the southern margin of the Black Sea (Fig. 1; Wigley and Farmer, 1982). Accordingly, westerly winds may have prevailed during early MIS 3 and during the cold MIS 4 and 2 northwesterlies dominated the atmospheric circulation over Europe (van Huissteden and Pollard, 2003; Renssen et al., 2007; Bokhorst et al., 2011). Since ca. 42 ka BP (Fig. 5 A) salinity is again gradually increasing. The beginning expansion of the Eurasian Ice Sheet at that time (Figs. 1, 5 D; Bintanja and van de Wal, 2008) pushed the atmospheric polar front southwards and the influence of the Siberian High pressure system over Northern Anatolia (e.g. Çağatay et al., 2014) with its dry polar air masses increased relative to the tropical air masses and Mediterranean cyclone tracks (Fig. 1).

Aloisi et al. (2015) suggest that a connection between the Marmara Sea and the Black Sea existed since at least 50 ka BP. The authors assume that large amounts of freshwater reached the Sea of Marmara via spill-out of the Black Sea due to a pronounced freshwater supply from the large drainage area (Çağatay et al., 2009; Aloisi et al., 2015). According to the Sr/Caostracods record this is in agreement with the Black Sea freshening only until ca. 42 ka BP. We hypothesise, that the increased salinity after 42 ka is accompanied by a decreasing water level of the Black Sea and a gradual separation from the Marmara Sea interrupted the Black Sea overflow. Although we cannot completely rule out a continuous connection to the Sea Marmara since 50 ka BP, Constantinescu et al. (2015) proposed that the Black Sea “Lake” level was ≤ -110 m during the LGM. Similarly, Eris et al. (2011) proposed a Black Sea-Marmara Sea connection not before 15 ka BP.

Low Sr/Caostracods values during interstadials are much less pronounced (Fig. 5A) but suggest a negative response of salinity to increased fluvial discharge and melting of the continental ice sheets during the early stage of the interstadials (Arz et al., 2007).

5.3 Productivity in the Black Sea “Lake” during Dansgaard-Oeschger cycles

The Black Sea “Lake” productivity increased considerably during the warm and more humid interstadials, especially during the longer interstadials 12 and 14, as reflected by increased TOC, TIC, and total dinocysts (Figs. 6 A-C; Nowaczyk et al., 2012; Shumilovskikh et al., 2014; Wegwerth et al., 2015). While TOC represent a mixture of both aquatic and terrestrial sources, TIC indicates authigenic carbonate precipitation due to enhanced biological activity in the lake (Bahr et al., 2008; Nowaczyk et al., 2012). Interstadial

116 Northern hemisphere climate control on the environmental dynamics in the glacial Black Sea “Lake” productivity was most likely triggered by higher summer temperatures and elevated nutrient supply due to the higher riverine discharge. Interestingly, TIC values of 3-4% during most interstadials are similar to values obtained for Terminations I (Major et al., 2002; Kwiecien et al., 2008) and II (Wegwerth et al., 2014), which suggests comparable productivity and temperatures. Temperatures during both these terminations ranged between 8 and 11°C (Ménot and Bard, 2012; Wegwerth et al., 2014), interstadial temperatures increased up to 9°C (Wegwerth et al., 2015).

The ostracod abundance reproduce the DO cycles as well (Fig. 6 D), but with an opposite pattern, i.e. high abundances occur during stadials and low abundances, or even a complete absence, during interstadials. The possible factors controlling the abundance and productivity of ostracods are temperature, salinity, food supply, and oxygen availability (Kiss, 2007). Preservation effects and variations of the water depth may affect the sedimentary ostracod abundance as well. Temperature and salinity were most likely not the limiting factors for ostracods in the Black Sea because similar conditions prevailed during Terminations I and II, when ostracods were generally more abundant (e.g. Bahr et al., 2008; Wegwerth et al., 2014). The role of food supply can also be ruled out as high TOC, TIC, and dinocysts concentrations attest a pronounced productivity during interstadials. However, a high surface productivity and the export of high amounts of organic matter may have resulted in oxygen deficiency at the sediment water interface critical for the survival of ostracods. We indeed observe higher values of the redox-sensitive Re/AlICP-MS ratio during interstadials (Fig. 6E) which may indicate suboxic conditions at the sediment/water interface because oxygen deficiency favours sedimentary Re enrichments (Crusius et al., 1996). While studies on the tolerance of minimum oxygen demands for most ostracod species (including Candona) are not available (Frenzel et al., 2010), we suggest that oxygen depleted bottom waters disturbed the ostracods’ ontogeny, finally resulting in the interstadial reduction of their abundance. Whether the reduced ostracod abundance may also relate to calcification and /or preservation effects is still a matter of debate (Frenzel and Boomer, 2005 and references therein). Some authors relate stronger calcification to increased organic matter input and others prefer the opposite argument that calcification suffers from decomposition of organic matter leading to lower pH values. The ostracod valves during interstadials are often thin-shelled and more fragile possibly arguing for a weaker preservation.

117 Northern hemisphere climate control on the environmental dynamics in the glacial Black Sea “Lake”

Figure 6: Temporal variations in productivity in surface and bottom water as reconstructed from core 25GC-1, SE Black Sea: A) content of total organic carbon (TOC), B) content of total inorganic carbon (TIC; Nowaczyk et al., 2012; Wegwerth et al., 2015), C) sum of dinoflagellate cysts (dinocysts; Shumilovskikh et al., 2014), D) abundance of benthic ostracod valves (Candona spp.), E) Re/Al possibly representing oxygen deficiency during some interstadials in bottom waters or at least at the sediment-water interface. The red numbers and grey bars denote the interstadials during Dansgaard-Oeschger (DO) cycles and the blue numbers denote Heinrich event (HE) equivalent periods.

Finally, the ostracod decline during interstadials could be related to a higher Black Sea “Lake” level. Boomer et al. (2005) found that decreasing abundances of ostracods correspond to increasing water depth in the Caspian Sea, with strongest gradient of specimens’ reduction in the upper 400 m of the water column. Therefore, a relatively wet climate and an enhanced freshwater input during interstadials potentially induced rises of the Black Sea “Lake” level and decreases of ostracod abundances. Boomer et al. (2005) additionally concluded that

118 Northern hemisphere climate control on the environmental dynamics in the glacial Black Sea “Lake” ostracods were most abundant and productive during cooler climatic conditions, suggesting a strong resemblance to the SE Black Sea “Lake” record. Thus, we cannot rule out the effect of potential lake level fluctuations on ostracod abundances, at least for some but not all interstadials and for the long-term trend since at least the MIS 3.

6. Summary and concluding remarks

In combination with previous works (Nowaczyk et al., 2012; Shumilovskikh et al., 2014; Wegwerth et al., 2015), the present study demonstrates a strong response of the climate and environmental conditions in the SE Black Sea region to northern hemisphere climate variability at a (multi-)millennial time-scale between 64 and 20 ka BP as summarised in Figure 7:

Stadials (Fig. 7A)

The southward migration of the atmospheric polar front during stadials increased the influence of polar air masses and stronger westerly winds in the SE Black Sea leading to lower temperatures and reduced rainfall in this region. As one consequence, the stadials are marked by an increase in salinity and probably a decreased lake level. Riverine nutrient supply and lake surface productivity was relatively low during these intervals and bottom waters were well-oxygenated favouring the occurrence of benthic ostracods. In the hinterland, the vegetation was dominated by a xerophytic steppe and most stadials experienced extreme winters leading to coastal ice formation. Aridity, less dense vegetation cover, and stronger westerly winds likely favoured an increased aeolian input. Conditions prevailing during stadials are comparable to the glacial conditions during MIS 4 and 2.

Interstadials (Fig. 7B)

The northward migration of the atmospheric polar front and the stronger influence of Mediterranean cyclone tracks caused extra warming and higher rainfall in the SE Black Sea region during interstadials. Milder winters, increased humidity, and associated higher riverine input of nutrients favoured enhanced productivity, lower salinity, and probably increasing lake levels. Enhanced export production likely triggered microbial decomposition of settling organic matter and the consumption of oxygen at the sediment-water interface. Latter might

119 Northern hemisphere climate control on the environmental dynamics in the glacial Black Sea “Lake” explain low ostracod abundances during these periods. In the hinterland, more humid conditions increased the contribution of the temperate forest biome. The dense vegetation cover and less intense westerly winds diminished the aeolian input.

Figure 7: Compilation scheme of the atmospheric circulation (left) and environmental (right) changes in the Black Sea “Lake” during the last glacial for A) stadials and B) interstadials. E: evaporation P: precipitation, R: riverine runoff, red circle (left) denote the location of core 25GC-1.

In a long-term perspective, the SE Black Sea record suggests that milder winters and wetter conditions during MIS 3 favoured an increased productivity, enhanced riverine discharge, freshening, and a lake level rise with a potential overspill through the Bosporus sill. These changes were most likely associated with an increased northern hemisphere summer insolation, the concomitant shrinking of the continental ice sheets and a northward displacement of the atmospheric polar front.

Acknowledgements The authors thank Anne Köhler and Nadine Hollmann for assistance during sample preparation and measurements of inorganic and organic geochemical parameters, respectively. Olga Körting, Sibylle Fink, and Katharina Häusler are acknowledged for their assistance

120 Northern hemisphere climate control on the environmental dynamics in the glacial Black Sea “Lake” during ostracod sampling. A.W. thanks J. White for inspiring accompaniment during manuscript writing. Finally, we thank the captain and crew of RV “METEOR” and the cruise leader Christian Borowski for their support during the M72/5 Black Sea cruise in 2007. This work was funded by the Deutsche Forschungsgemeinschaft (BE2116/20-1, AR 367/9-1, AR 367/9-2, and FL 710/1-1) within the priority program 1266 INTERDYNAMIC, by the Gary Comer Science and Education Foundation, by the Helmholtz Centre Potsdam - German Research Centre for Geosciences (GFZ-Potsdam), by the Alfred-Wegener-Institut Helmholtz- Zentrum für Polar und Meeresforschung (Bremen), and by the Leibniz Institute for Baltic Sea Research (IOW).

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128 Redox evolution during Eemian and Holocene sapropel formation in the Black Sea

Redox evolution during Eemian and Holocene sapropel formation in the Black Sea

Antje Wegwertha*, Sebastian Eckertb, Olaf Dellwiga, Jürgen Kösterb, Bernhard Schnetgerb,

Silke Severmannc, Stefan Weyerd, Annika Brüsked, Jérôme Kaisera,

Helge W. Arza, and Hans-Jürgen Brumsackb

a Leibniz-Institute for Baltic Sea Research Warnemünde (IOW), Marine Geology, Rostock, Germany b Institute for Chemistry and Biology of the Marine Environment (ICBM), University of Oldenburg, Germany c Institute of Marine and Coastal Sciences, Rutgers University, New Brunswick, NJ 08901, USA d Leibniz Universität Hannover, Institut für Mineralogie, Hannover, Germany

in preparation

129 Redox evolution during Eemian and Holocene sapropel formation in the Black Sea

Abstract

The Black Sea is the type locality for modern and ancient sapropel-forming sulfidic (euxinic) basins and attracted attention since decades. During past glacial-interglacial transitions however, the basin experienced repeatedly fundamental changes from limnic/oxic to brackish/euxinic conditions. Although the Holocene brackish period that started at about 9 ka BP has been intensively studied, the redox evolution of previous euxinic stages is almost unexplored. Here we investigate for the first time a Holocene and Eemian sapropel record from the same site and water depth in the SE Black Sea at high-resolution in order to compare both redox scenarios. Trace metal contents and metal isotope proxies comprising Mo, Re, W, δ98Mo, and δ56Fe highlight a straightforward and textbook-like redox development from oxic, suboxic, and finally strong euxinic conditions after the Eemian reconnection to the Mediterranean Sea. In contrast to previously reported up- and downward migrations of the redoxcline during the Holocene, our Eemian sapropel section spanning about 6.5 ka reveals an almost continuously rising redoxcline. Favored by the higher Eemian sea level and salinity, the degree of euxinia and trace metal enrichments distinctly exceed the Holocene. This study identifies the different climate background as most important for the specific redox conditions in the Eemian and Holocene Black Sea.

Introduction

The Black Sea represents the world’s largest brackish basin due to the inflow of salty Mediterranean water through the shallow ca. 35 m deep Bosporus strait and freshwater input via rivers and rainfall. The resulting pycnocline, separating denser bottom from less saline surface waters, limits vertical exchange and deep-water ventilation finally promoting the formation of a pelagic redoxcline and euxinic conditions below ca. 150 m (e.g. Özsoy and Ünlüata, 1997). The redox evolution during the present brackish stage, which started at about 9 ka BP (e.g. Soulet et al., 2011), attracted attention since more than a century (e.g. Murray, 1900; Degens and Ross, 1972). After Degens and Ross (1972), the Holocene sapropel is divided into Unit I and Unit II separated by the massive invasion of the coccolithophoride Emiliana huxleyi at ca. 2.76 ka BP (Lamy et al., 2006). The analysis of redox-sensitive and sulfide-forming trace metals in sapropel samples by Brumsack (1989) displayed the ocean water portion of the Black Sea water column as the ultimate source for sedimentary metal enrichments. Strong enrichments of Mo, Re, and U just after the reconnection to the

130 Redox evolution during Eemian and Holocene sapropel formation in the Black Sea

Mediterranean reflect enhanced metal supply and reducing water column conditions (e.g. Piper and Calvert, 2011). Even though euxinic bottom waters favor trace metal fixation, the Holocene Black Sea sapropel is less enriched when compared to Mediterranean sapropels or Black Shales because the restricted inflow of ocean water is not able to compensate metal sequestration finally leading to severe bottom water depletion of several trace metals (e.g. Brumsack, 2006; Algeo and Maynard, 2008). First Mo-Isotope data (δ98Mo) revealed larger fractionation in the Unit II sapropel when compared with Unit I due to the different degree of euxinic conditions (Arnold et al., 2004; Nägler et al. 2005). Besides pronounced perturbations of the Black Sea redoxcline during the past centuries (Lyons et al., 1993), Eckert et al. (2013) could also show by Fe contents and isotopes (δ56Fe) that the gradual rise of the redoxcline after the reconnection to the Mediterranean was interrupted by a major redoxcline descent at ca. 5.3 ka BP before reaching the present situation. Because the Black Sea was at least 12 times connected to the Mediterranean Sea since about the last 670 ka BP (Badertscher et al., 2011), such fundamental changes from limnic/oxic to brackish /euxinic conditions most likely form a typical feature. However, the knowledge about the redox conditions during the last interglacial (Eemian, MIS 5e, 130-115 ka BP) is relatively poor (e.g. Quan et al., 2013) because of the lack of appropriate sediment material.

Based on well-preserved gravity cores from the SE Black Sea covering the period from 133.5 ± 0.7 to 119.7 ± 1.7 ka BP (Shumilovskikh et al., 2013a, b; Wegwerth et al., 2014), we unravel for the first time at high resolution the Eemian redox evolution and compare it with the Holocene history at the same site and water depth by using new and previously published datasets (Table S 1). Because the warmer Eemian climate caused an up to 10 m higher global sea level (Grant et al., 2012; Martrat et al., 2014) and consequently a higher salinity in the Black Sea (Shumilovskikh et al., 2013a), the Eemian redox history is expected to differ considerably from that of the Holocene. While the TEX86 paleothermometer, Sr/Ca of benthic ostracod carbonate shells, and dinoflagellate assemblages provide environmental background information, trace metals including e.g. Mo and Re as well as the δ98Mo and δ56Fe isotope proxies serve for the reconstruction of both redox scenarios.

Material and methods

Three gravity cores (22GC-3/-7/-8) and two short cores (22MUC-1/-2) were retrieved from the Archangelsky Ridge in the SE Black Sea in 850 m water depth covering large parts of the Eemian and the entire Holocene (Fig. S1). For details on the Eemian and Holocene age

131 Redox evolution during Eemian and Holocene sapropel formation in the Black Sea models, which are based on previous works (Shumilovskikh et al.; 2012, 2013a; Wegwerth et al., 2014) see Fig. S2. Although strongly increasing Ti/Al and Zr/Al values (Fig. S3) indicate a potentially earthquake-related turbidite, which caused hiatus at ca. 119.7 ka BP, at least 8.5 ka of the Eemian are recovered.

Calculation of TEX86 and corresponding temperatures were based on measurements of glycerol dialkyl glycerol tetraethers (GDGT) by HPLC-APCI-MS. Isorenieratene derivatives of polar and apolar fractions were measured by GC/GC-MS. Sr/Ca of ostracods was determined by ICP-MS. Contents of Mo, Re and W were determined by ICP-MS and major elements including Sr and Zr were measured by ICP-OES and XRF. Mo and Fe isotopic compositions, reported here as δ98Mo and δ56Fe notations, were determined by MC-ICP-MS. Total organic carbon (TOC) was calculated from the difference between total carbon (TC) and total inorganic carbon (TIC) measured by elemental analyzers and coulometry. For details on methodology see supplementary information.

Results and discussion

Environmental background during the Holocene and Eemian

The onset of Holocene warming is documented in mean annual surface TEX86-based temperatures (max. 18°C at 8 ka BP) after the Younger Dryas cold period (Fig. 1A). A stronger and especially more rapid glacial-interglacial temperature increase is seen for the Eemian reaching up to 22°C at ca. 128 ka BP (Wegwerth et al., 2014), which is in agreement with studies across the northern hemisphere revealing an about 3°C warmer early Eemian compared to the early Holocene (e.g. Seidenkrantz, 1993; Martrat et al., 2014). As previously postulated for the Eemian (Wegwerth et al., 2014), abruptly falling

TEX86 temperatures after the Holocene maximum warming most likely reflect a downward migration of isoGDGT producing Thaumarchaeota towards colder and ammonia-rich subsurface waters (e.g. Wakeham et al. 2003; Menzel et al., 2006). In relation to the Holocene, the higher sum of thermophilic dinocysts taxa clearly point to considerably higher temperatures during the Eemian (Fig. 1B; Shumilovskikh et al., 2013a). The presence of the (sub-)tropical dinocyst Tuberculodinium vancampoae, which is absent during the entire Holocene, even implies maximum Eemian summer temperatures of up to 27-29°C between 127.5 and 125 ka BP (Shumilovskikh et al., 2013a).

132 Redox evolution during Eemian and Holocene sapropel formation in the Black Sea

Figure 1: Temporal variation of temperature and salinity related proxies from the last two glacials towards the interglacials in the SE Black Sea at site M72/5 22. A) TEX86 temperatures (lake subsurface temperatures (circles) after Powers et al. (2010); sea surface temperatures (diamonds) after Kim et al. (2008); Eemian data from Wegwerth et al. (2014)), B) sum of thermophilic dinocyst taxa indicating surface temperature changes (Shumilovskikh et al. (2013a), C) Sr/Caostracods indicating changes in bottom water salinity (~850 m water depth at the study site 22; Eemian data from Wegwerth et al. (2014)). D) sum of fully marine dinocyst taxa indicating changes in surface salinity (Shumilovskikh et al. (2013a), and E) concentration of the fully marine dinoflagellate species Spiniferites pachydermus indicating salinity of 28-30‰ (Shumilovskikh et al. (2013a). The differentiation between the Holocene Units I and II bases on the CaCO3 content and Sr/Ca ratios shown in Fig. S3. The blue bars denote sediments enriched in coccoliths, G: Glacial, IG: Interglacial, blue dotted line indicates the Mediterranean-Black Sea reconnection.

Along with global warming and coupled sea-level rise, increasing Sr/Ca ratios of benthic ostracods (Sr/Caostracods, Fig. 1C) date the Mediterranean-Black Sea reconnection to ca. 9 ka BP in the Holocene and to 128.1 ka BP in the Eemian (Wegwerth et al., 2014). First appearance of fully marine dinocysts after both reconnections imply increasing surface water salinity at ca. 8.1 ka BP and ca. 127.6 ka BP, respectively (Fig. 1D; Shumilovskikh et al., 2013a). Sea surface salinity reconstructions using dinocyst process lengths suggest a gradual increase from 13-14‰ at ca. 8 ka BP towards the present level of 17-18‰ at ca. 4.1 ka BP (Mertens et al., 2012). The early Eemian dinocyst assemblages, in turn, indicate a surface

133 Redox evolution during Eemian and Holocene sapropel formation in the Black Sea salinity distinctly above 20‰ (Shumilovskikh et al., 2013a ) due to the ca. 10 m higher sea level (Grant et al, 2012). After 126.5 ka BP several fully marine dinocysts (e.g. Spiniferites pachydermus, Fig. 1E), which are absent during the Holocene, even suggest an Eemian surface salinity of up to 28-30‰ (Shumilovskikh et al., 2013a).

Holocene and Eemian redox evolution

In the course of increasing productivity and better preservation of organic matter within a progressively more oxygen-deficient water column, Holocene sapropel formation documented by TOC contents exceeding 2% (Calvert, 1983) started shortly after the reconnection at ca. 8.5 ka BP (Fig. 2A). A TOC level comparable to that of the Holocene Unit II is also seen in the Eemian since ca. 127.7 ka BP (Fig. 2A).

Redox-sensitive trace metals are widely used as diagnostic tools for the reconstruction of past redox states and productivity (e.g. Brumsack, 2006; Tribovillard et al., 2006). Due to their differing response to changing redox conditions, trace metals like Mo and Re potentially allow differentiation between suboxic/anoxic and euxinic conditions. While enhanced burial of Mo affords elevated H2Saq levels exceeding the “action point of switch” (APS) of about 11 2- µM at which highly soluble molybdate (MoO4 ) is transferred into particle-reactive - thiomolybdates (Helz et al., 1996; Erickson and Helz, 2000), perrhenate (ReO4 ) is already trapped under suboxic or intermediate reducing conditions (e.g. Crusius et al., 1996; Morford et al., 2005).

About 3 ka after the Holocene reconnection, maximum Mo/Al values were reached at ca. 6 ka BP and remain comparatively stable on a distinctly lower level thereafter (Fig. 2C). At a first glance, this pattern may be compatible with increasing sulfide concentrations in the water column followed by bottom water Mo depletion and/or less sulfidic conditions afterwards. By contrast, Eemian Mo/Al values increase almost gradually for ca. 6.5 ka until reaching a two-fold higher level just before the occurrence of the most pronounced coccolith ooze (Fig. 2C, compare also Figs. S 3C,D). Besides stronger sequestration from a progressively more sulfidic water column, the higher Mo level during the Eemian may also reflect a larger water column trace metal reservoir favored by the higher sea level (Grant et al., 2012) and elevated salinity (Figs. 1D,E), respectively. According to Algeo and Lyons (2006) the latter assumption can be tested by the covariation between Mo and TOC. While a slope m <15 indicates a strongly restricted and sulfidic basin like the modern Black Sea and

134 Redox evolution during Eemian and Holocene sapropel formation in the Black Sea

Framvaren Fjord, values ranging between 15 and 35 point towards moderately restricted but also strong sulfidic conditions prevailing in the Cariaco Basin. Indeed, the correlation between Mo and TOC reveal a distinctly steeper slope (m ≈22) for the Eemian compared to the Holocene (m ≈6) and thus a larger Mo water column reservoir due to better Eemian deep water renewal (Fig. S4).

Because seawater forms the ultimate source for sedimentary enrichments of trace metals like Mo, Re, or U, normalization to Al, which is contributed by detrital input, may be compromised by different sedimentation rates (Brumsack, 1989). Therefore, trace metal ratios like Re/Mo or Mo/U represent alternative proxies due to the different behavior of the individual elements under changing redox conditions (Crusius et al., 1996; Algeo and Tribovillard, 2009). As the Eemian record is possibly biased by U co-precipitation with carbonate phases like MnCO3 overgrowths on coccolithophorides (Fig. S5; Brumsack, 2006), we put our focus on the Re/Mo proxy in the present study.

Increasing Re/Mo values clearly imply preferable Re burial until ca. 7. ka BP and 126.5 ka BP and the establishment of suboxic conditions without substantial presence of H2S, respectively (Fig. 2D). The stronger increase of Re/Mo during the Eemian possibly further indicates both the faster development of oxygen-deficiency and higher salinity-related water column Re availability. Although, the use of Re/Mo may suffer from deepwater depletion due to a stronger loss of Mo relative to Re in the modern Black Sea water column (Algeo and Maynard, 2008), such circumstances appear unlikely for these periods shortly after the Mediterranean-Black Sea reconnections.

The presence of suboxic or at least less sulfidic conditions in the early Holocene and Eemian finds further support in Mo isotopes. While nearly complete stripping of Mo from a strongly euxinic water column (H2Saq>>APS; Helz et al., 1996) results in sedimentary Mo isotope signatures comparable to the oxygenated ocean (mean ocean molybdenum, MOMo; 98 δ Mo = +2.3‰, Siebert et al., 2003), Mo is prone to isotopic fractionation under lower H2Saq concentrations due to partial removal of the lighter Mo-isotopes (Neubert et al., 2008; Nägler et al., 2011). Following this concept, comparatively light δ98Mo values of about +1‰ imply fractionation during thiomolybdate formation in an early Eemian water column (127-125 ka 98 BP; Fig. 2E) still having relatively low H2Saq concentrations. Thereafter, δ Mo values increase and nearly approach MOMo around ca. 123 ka BP. Along with low Re/Mo and gradually rising Mo/Al values, the increase in Eemian δ98Mo suggests a progressive change towards stronger euxinic bottom waters and almost complete Mo-fixation. Moreover, δ98Mo-

135 Redox evolution during Eemian and Holocene sapropel formation in the Black Sea

based estimations of past H2Saq levels after Arnold et al. (2012) reveals concentrations exceeding the APS at least since 124.7 ka BP (Fig. S6). This potential onset of pronounced euxinia corresponds with enhanced presence of isorenieratene derivatives (Fig. 2B) indicating the occurrence of the photosynthetic green sulfur bacteria Chlorobiaceae and thus photic zone euxinia (e.g. Sinninghe-Damsté and Köster, 1998).

Figure 2: Temporal variation of the redox-related proxies from the last two glacials towards the interglacials in the SE Black Sea at site M72/5 22. A) Total organic carbon (TOC; dashed line indicates the limit >2% TOC for sapropels; Calvert (1983)), B) The abundance of isorenieratene derivatives in the polar and apolar fractions indicating photic zone euxinia, C) Mo/Al, D) Re/Mo (molar). E) δ98Mo, F) W/Al, G) Fe/Al, and H) δ56Fe (note reversed scale). The blue bars denote sediments enriched in coccoliths, G: Glacial, IG: Interglacial, blue dotted line indicates the Mediterranean-Black Sea reconnection, the redox conditions in the water column are color-coded for differentiation between oxic, suboxic, and euxinic conditions. Eemian TOC partly from Wegwerth et al. (2014), Holocene TOC, Mo/Al, Fe/Al, and δ56Fe partly from Eckert et al. (2013).

136 Redox evolution during Eemian and Holocene sapropel formation in the Black Sea

A further potential indicator for strong sulfidic conditions in the Eemian may represent

W because the formation of thiotungstate requires essentially higher H2Saq concentrations when compared with thiomolybdate (Mohajerin et al., 2014). Although the geogenic background of W and Mo is on a comparable level (Hu and Gao, 2008), 100-fold lower seawater concentrations of W (Mendel, 2005) may limit the detection of diagenetic W enrichments. Nevertheless, gradually increasing W/Al values are possibly compatible with distinctly higher water column H2S concentrations in the Eemian record, whereas the Holocene sapropel remains almost constant and only slight enrichments appear during the Mo/Al maximum (Figs. 2C,F).

98 Distinctly lighter δ Mo values and the calculated H2Saq level indicate only weakly euxinic bottom waters below the APS throughout the entire Holocene Unit II (Figs. 2E, S6). As scavenging of isotopically light Mo by redoxcline-derived Mn-oxide particles causes strong isotope fractionation, sedimentary δ98Mo values were possibly also influenced by the deposition of these particles under sulfide-poor conditions (Barling and Anbar, 2004; Dellwig et al., 2010; Scholz et al., 2013). A similar scenario causing a light Mo-isotope signature was also proposed for sediments deposited in the occasionally euxinic Landsort Deep of the Baltic Sea (Noordmann et al., 2015).

Sedimentary Fe/Al and δ56Fe signatures are able to track changes in the so-called shelf- to-basin Fe-shuttle associated with vertical migrations of the redoxcline because the impingement of suboxic waters on shelf areas allows release and transport of isotopically light reactive Fe (Lyons and Severmann, 2006; Severmann et al., 2008). Thus, Eckert et al. (2013) recently showed distinct Fe enrichments along with light δ56Fe values at our site, which indicate a short-term redoxcline highstand during the mid-Holocene (Figs. 2G,H). A low level of isorenieratene derivatives during this period (Fig. 2B), however argues against pronounced photic zone euxinia. Lower Fe/Al values and heavier Fe isotopes thereafter most likely reflect a weakening of the Fe-shuttle due to a descending redoxcline and less sulfidic conditions. This assumption would be compatible with the second increase in Re/Mo (Fig. 2D), which may also benefitted from an increasing salinity during that time (Mertens et al., 2012). By contrast, a prominent Fe peak within the Eemian sapropel section is missing but Fe/Al values generally reveal a distinctly higher when compared with the Holocene. After the early Eemian suboxic stage, δ56Fe values become almost gradually lighter until 120.7 ka BP and therefore imply a continuously shoaling redoxcline. Along with decreasing Fe/Al, the

137 Redox evolution during Eemian and Holocene sapropel formation in the Black Sea heavier δ56Fe value thereafter in the coccolith ooze is possibly related to a redoxcline descent, although a biased signal caused by the turbidite cannot be ruled out.

Because Mo/Al values decrease rapidly before Fe contents and isotopes imply a redoxcline highstand especially in the Eemian, it is worthwhile to address the possible reasons for this behavior of Mo. In addition to the exhaustion of the modern water column Mo reservoir due to insufficient replenishment via the Bosporus (Algeo and Maynard, 2008), decreasing Mo/Al ratios may also augmented by intruding seawater, which was possibly depleted in Mo during the expansion of Mediterranean euxinia. When considering the almost coevally forming Mediterranean S5 sapropel reaching even higher Mo/Al values of up to 60*10-4 as well as photic zone euxinia, this process might be of relevance at least for our Eemian sapropel (Gallego-Torres et al., 2010; Rohling et al., 2006). On the other hand, less pronounced trace metal enrichments (max. Mo/Al: 15*10-4, Gallego-Torres et al., 2010) and the lack of photic zone euxinia during S1 formation (Rohling et al., 2015) argue against a significance for the Holocene Black Sea. To test the importance of reservoir exhaustion for Mo, simple balance calculations have been performed (Table S2), which imply that the exhaustion of the initial Mo water column reservoir would affect sedimentary Mo patterns at about 5.8 ka BP and 122.3 ka BP. Since this simple calculation is in fair agreement with both sedimentary records (Fig. 2C), water column reservoir exhaustion appears to be the most important mechanism explaining the decreases in Mo enrichments.

Conclusions

The redox evolution during Eemian and Holocene sapropel formation in the Black Sea differed remarkably. While an ideal sequence from suboxic to finally strong euxinia is documented in the Eemian sapropel, the Holocene Unit II is subjected to redoxcline migrations and less pronounced euxinic conditions. The stronger enrichment of redox- sensitive trace metals in the Eemian sapropel is in all probability related to the up to 10 m higher sea level and the associated larger trace metal reservoir. In contrast to the Holocene, the Eemian record reveals a heavier δ98Mo signature and presence of isorenieratene derivatives pointing to considerably more pronounced sulfidic conditions and even photic zone euxinia. Furthermore, warmer conditions and elevated nutrient supply during the Eemian may also have favored the development of stronger euxinia. Considering present global warming and associated sea-level rise, the Eemian sapropel record may provide a window into the future evolution of the Black Sea.

138 Redox evolution during Eemian and Holocene sapropel formation in the Black Sea

Acknowledgments

The authors thank the captain and crew of R/V Meteor and the cruise leader, C. Borowski, for their support during the M72/5 Black Sea cruise. A.W. thanks L. S. Shumilovskikh for providing dinoflagellate datasets. The authors thank E. Gründken, A. Köhler, C. Lehners, and I. Scherff for assistance during sample preparation and bulk parameter measurements. A.W. acknowledges E. Bard and G. Ménot enabling laboratory work at CEREGE and Y. Fagault for kind assistance during GDGT measurements. O. Körting and S. Fink are thanked for supporting ostracod sampling. A.W. is indebted to R. Erickson for inspiring contributions during the writing of the manuscript. This work was funded through grants by the German Science Foundation to H. W. Arz (AR 367/9-1, 2) and to H.-J. Brumsack (Br 775/15 and 775/26).

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143 Redox evolution during Eemian and Holocene sapropel formation in the Black Sea

Supplementary Information

Redox evolution during Eemian and Holocene sapropel formation

in the Black Sea

Antje Wegwerth1*, Sebastian Eckert2, Olaf Dellwig1, Jürgen Köster2, Bernhard Schnetger2,

Silke Severmann3, Stefan Weyer4, Annika Brüske4, Jérôme Kaiser1,

Helge W. Arz1, and Hans-Jürgen Brumsack2

1Leibniz-Institute for Baltic Sea Research Warnemünde (IOW), Marine Geology, Rostock,

Germany

2Institute for Chemistry and Biology of the Marine Environment (ICBM), University of

Oldenburg, Germany

3Institute of Marine and Coastal Sciences, Rutgers University, New Brunswick, NJ 08901,

USA

4Leibniz Universität Hannover, Institut für Mineralogie, Hannover, Germany

144 Redox evolution during Eemian and Holocene sapropel formation in the Black Sea

Details on Material and Methods

Methods

The preparation for GDGT (glycerol dialkyl glycerol tetraethers) analysis involved accelerated solvent extraction (Thermo Scientific™ Dionex™ ASE™ 350) of freeze-dried and homogenized samples (4-6 g of sediment) with a dichloromethane/methanol mixture (DCM/MeOH 9:1), followed by desulfurization with activated copper pieces. The separation of the polar fraction containing GDGTs has been done with DCM/MeOH (1:1) over a Pasteur pipette plugged with activated Al2O3. After filtration through a 0.45μm PTFE filter and addition of a C:46 GDGT standard, GDGTs were determined by HPLC-APCI-MS

(HPLCMS1100 Series, Hewlett Packard, USA). TEX86 was calculated according to Schouten et al. (2002) and corresponding temperatures by using the global lake calibration for the glacial (Powers et al., 2010) and the global marine subsurface (0–200m) calibration for the interglacial (Kim et al., 2008). Concentrations of isorenieratene derivatives of polar and apolar fractions were measured by gas chromatography-mass spectrometry by a Thermo Scientific TSQ Quantum GC-MS coupled to a Thermo Scientific Trace GC Ultra gas chromatograph (Koopmans et al., 1996).

The ostracod valves of adult Candona spp. were handpicked under a binocular from the dried, wet-sieved sediment samples (>150 µm). The valves were cleaned with a thin brush and few drops of deionized water. Up to five valves were measured for Sr/Ca ratios by ICP-MS (iCAP

Q; ThermoFisher Scientific) after dissolution in 2 vol.% HNO3 (sub-boiled) and centrifugation.

For major and trace elements, sediment samples were digested by a HNO3/HClO4/HF mixture (Schnetger, 1997) in closed Teflon bottles at 180°C (PDS-6, Heinrichs et al, 1986). After evaporation of the acids on a metal plate at 180°C, the digestions were fumed off 3-times with half-concentrated HCl and finally diluted with 2 Vol.% HNO3 to 25 mL. From these acid digestions, Mo, Re, and W were measured by ICP-MS (iCAP Q; Thermo Fisher Scientific) and Al, Ca, Fe, Ti, Sr, and Zr by ICP-OES (iCAP 6300Duo, Thermo). For high-resolution records of Al, Ca, Fe, Ti, Sr, Zr and Mo, selected sections of cores 22GC-3/7 and 22MUC-2 were also analyzed by X-ray fluorescence (XRF) with a Philips PW-2400 WD-XRF spectrometer (Eckert et al., 2013).

145 Redox evolution during Eemian and Holocene sapropel formation in the Black Sea

Mo and Fe isotopic compositions, reported here as δ98Mo and δ56Fe notations were determined by MC-ICP-MS (Thermo Fisher Scientific Neptune) following the methods by Noordmann et al. (2015) for δ98Mo and by Arnold et al. (2004) for δ56Fe.

Total organic carbon (TOC) was calculated from the difference between total carbon (TC) and total inorganic carbon (TIC) measured by elemental analyzers (EA 1110, CE-instruments and EA®2000, Analytik Jena; Wegwerth et al., 2014) as well as by an ELTRA CS-500 IR- spectrometer and UIC CM-5012 (Eckert et al., 2013).

Study area

Figure S1: Study area and locations of gravity cores (GC) and multi-cores (MUC) retrieved at site 22 during RV “Meteor” cruise M72/5 in 2007.

146 Redox evolution during Eemian and Holocene sapropel formation in the Black Sea

Table S1: Overview about parameters applied on the sediment cores in this study. core period parameter source

22GC-3 Holocene TOC, TEX86, Sr/Caostracods, CaCO3, Sr/Ca, this study 98 δ Mo, W/Al, Ti/Al, Zr/Al, Mn/Al, U/Al dinoflagellates Shumilovskikh et al. (2013a) 22GC-7 Holocene TOC, Mo/Al, Fe/Al, δ56Fe Eckert et al. (2013) isorenieratene derivatives, Ti/Al, Zr/Al, this study Mn/Al, U/Al

22MUC-1 Holocene TOC, TEX86 Wegwerth et al. (2014) Mo/Al, Fe/Al, Re/Mo, δ98Mo, W/Al this study 22MUC-2 Holocene TOC, Mo/Al, Fe/Al this study 22GC-3 Eemian dinoflagellates Shumilovskikh et al. Fe/Al, Mo/Al, Re/Mo, δ98Mo, δ56Fe, (2013a) isorenieratene derivatives, W/Al, Ti/Al, this study Zr/Al, Mn/Al, U/Al 22GC-7 Eemian δ56Fe, isorenieratene derivatives this study

22GC-8 Eemian TOC, TEX86, Sr/Caostracods, CaCO3, Sr/Ca Wegwerth et al. (2014) Fe/Al, Mo/Al, Re/Mo, W/Al, Ti/Al, Zr/Al, this study Mn/Al, U/Al

147 Redox evolution during Eemian and Holocene sapropel formation in the Black Sea

Age model

Figure S2: Age models for the Holocene (left) and Eemian (right). a) Stratigraphic correlation for the Holocene between gravity cores M72/5 22GC-3 and 22GC-7 by means of TOC contents. Ca contents approve core correlation. Open circles denote surface sediments of 22MUC-1 (black; 0-1 cm) and 22MUC-2 (blue; 0-5 cm). Triangles denote tie points for age model of 22GC-3 obtained by Shumilovskikh et al. (2012). b) Age-depth relation of cores M72/5 22GC-3, 22GC-7, 22MUC-1, and 22MUC-2. c) Stratigraphic correlation for the Eemian between gravity cores M72/5 22GC-3, 22GC-7, and 22GC-8 by means of Ca contents. Triangles with dots denotes pollen-stratigraphic correlation on core 22GC-3 (Shumilovskikh et al., 2013a). All triangles denotes tie points of XRF-Ca between 22GC-3 and 22GC-8 (Wegwerth et al., 2014). 22GC-3 encompasses the largest part of the Eemian but the Eemian is probably not fully preserved due to a sharp transition from the Eemian to a turbidite (cp. Fig. S3). d) Age-depth relation of cores M72/5 22GC-3, 22GC-7, and 22GC-8.

148 Redox evolution during Eemian and Holocene sapropel formation in the Black Sea

Additional Geochemical Data

Figure S3: Ti/Al (A), Zr/Al values (B), CaCO3 contents (C), and Sr/Ca ratios (D) of the Holocene and Eemian Black Sea records. The sharp increases of Ti/Al and Zr/Al in the Eemian record indicate a turbidite base at ca. 119.7 ka BB causing a hiatus. While increasing CaCO3 contents after the Holocene reconnection are still caused by aragonite precipitation as seen in elevated Sr/Ca ratios, the carbonate peak at ca. 2.5 ka BP marks the transition from Unit II to Unit I (blue bar) due to the pronounced invasion of the coccolithophoride Emiliana huxleyi (Ross and Degens, 1974; Lamy et al., 2006). The transition to an Unit I analogue in the Eemian is possibly reflected in the massive deposition of coccolith ooze since 121 ka BP (blue bar). Relative to the Mediterranean reconnections, Eemian coccolithophoride invasions however, appeared distinctly earlier (light blue bars) than during the Holocene. Indeed, long- chain alkenones (Wegwerth et al., 2014) document a first occurrence of coccolithophorides already at 127.6 ka BP most likely favored by stronger intrusion of Mediterranean waters. Because Emiliana huxleyi dominance appeared not before 73-85 ka BP, Eemian coccolithophorides probably comprised of Gephyrocapsa oceanica (Thierstein et al., 1977). G: Glacial, IG: Interglacial, blue dotted line indicates the Mediterranean-Black Sea reconnection.

149 Redox evolution during Eemian and Holocene sapropel formation in the Black Sea

Figure S4: Scatter plot of TOC and Mo contents of the Holocene (black dots) and Eemian (red diamonds) sapropel. According to Algeo and Lyons (2006), steeper slopes of regressions indicate higher availability of dissolved Mo in the water column and better deep water renewal, respectively.

Figure S5: CaCO3 contents (A), Mn/Al (B), and U/Al (C) during the Holocene and Eemian in the Black Sea. Carbonate-rich intervals during the Eemian show elevated Mn/Al and U/Al values possibly implying incorporation in the carbonate phases like diagenetic Mn carbonate overgrowth on e.g. coccolithophorides. The blue bars denote sediments enriched in coccoliths, G: Glacial, IG: Interglacial, blue dotted line indicates the Mediterranean-Black Sea reconnection.

150 Redox evolution during Eemian and Holocene sapropel formation in the Black Sea

Figure S6: Mo contents (A) and estimated water column H2Saq concentrations (B) calculated from δ98Mo values (Fig. 2E) following the approach of Arnold et al. (2012). The “action point of switch” (red line; H2Saq> 11µM; Helz et al., 1996; Erickson and Helz, 2000) allowing nearly complete transfer of dissolved Mo into particle-reactive thiomolybdates is reached at least since 124.7 ka BP in the Eemian and since ca. 1 ka BP in the Holocene. The blue bars denote sediments enriched in coccoliths, G: Glacial, IG: Interglacial, blue dotted line indicates the Mediterranean-Black Sea reconnection, the redox conditions in the water column are color-coded for differentiation between oxic, suboxic, and euxinic conditions.

Balance Calculations

Whilst taking into account the modern inflow into the Black Sea (300 km3 a-1) with a salinity of 35.5 psu (Özsoy and Ünlüata, 1997), a salinity-based conservative mixing model after Nägler et al. (2011) would imply a yearly Mo input of 3,063 t. By considering a deep water salinity of 22.4 mg kg-1 for the Holocene, the initial inventory of the water column (total Black Sea volume: 530,000 km3; Özsoy and Ünlüata, 1997) amounts to about 3.5 Mt Mo. Due to the 10 m higher sea level during the Eemian (Grant et al., 2012) the inflow likely increased to about 386 km3 a-1 thereby enhancing the input of Mo to about 3,941 t a-1. Concerning the higher Eemian surface salinity of about 30 psu (Shumilovskikh et al., 2013a) the inventory further increases to about 4.7 Mt Mo. Along with the calculated annual Mo contribution via the Bosporus, salinity-based Mo inventories of the Black Sea basin would be reached after about 1.1 ka in the Holocene and 1.5 ka in the Eemian, which is for the Holocene in line with previous studies (e.g. Eckert et al., 2013). With respect to the Mediterranean reconnections, we thus set the onsets for complete Mo reservoirs at 7.9 ka BP and 126.5 ka BP. When considering these onsets, an average dry bulk density of 1.5 g cm-3 (Eckert et al., 2013), and estimated accumulation rates, exhaustion of the initial water column Mo reservoir would impact sedimentary Mo patterns at about 5.8 ka BP and 122.3 ka BP Mo.

151 Redox evolution during Eemian and Holocene sapropel formation in the Black Sea

Table S2: Parameters used for balance calculations of the Holocene and Eemian Mo inventories and of the possible Mo reservoir depletion.

parameters Holocene Reference

inflow [km3 a-1] 300 Özsoy and Ünlüata, 1997 inflow salinity [mg kg-1] 35.5 Özsoy and Ünlüata, 1997 Mo conc. inflow [µg l-1] 10.2 salinity-based conservative mixing model Nägler et al. (2011) yearly Mo input [t] 3,063 salinity-based conservative mixing model Nägler et al. (2011)

-1 salinity bottom water Black Sea [mg kg ] 22.4 Özsoy and Ünlüata, 1997 Black Sea area [km2] 420,000 Özsoy and Ünlüata, 1997 Black Sea volume [km3] 530,000 Özsoy and Ünlüata, 1997

-1 Mo at salinity bottom water [µg l ] 6.6 salinity-based conservative mixing model Nägler et al. (2011) initial Mo inventory [Mt] 3.5 calculated duration for complete Mo inventories [ka ] 1.1 calculated onset complete Mo reservoir [ka BP] 7.9 calculated av. Dry bulk density [g cm-3] 1.5 Eckert et al., 2013 sediment accumulation rate [g cm-2 ka-1] 4-27 (av. 11) calculated exhausted initial Mo reservoir [ka BP] 5.8 calculated by Mo accumulation rates (cumulative)

parameters Eemian Reference

inflow [km3 a-1] 386 calculated with assumed 10m higher sea level (Grant et al., 2012) inflow salinity [mg kg-1] 35.5 same as for Holocene is assumed Mo conc. inflow [µg l-1] 10.2 salinity-based conservative mixing model Nägler et al. (2011) yearly Mo input [t] 3,941 salinity-based conservative mixing model Nägler et al. (2011)

-1 salinity bottom water Black Sea [mg kg ] 30 Shumilovskikh et al., 2013a Black Sea area [km2] 420,000 same as for Holocene is assumed Black Sea volume [km3] 534,200 calculated with assumed 10m higher sea level (Grant et al., 2012)

-1 Mo at salinity bottom water [µg l ] 8.7 salinity-based conservative mixing model Nägler et al. (2011) initial Mo inventory [Mt] 4.7 calculated duration for complete Mo inventories [ka ] 1.5 calculated onset complete Mo reservoir [ka BP] 126.5 calculated av. Dry bulk density [g cm-3] 1.5 same as for Holocene is assumed

sediment accumulation rate [g cm-2 ka-1] 2.2-13.5 (av. 6) calculated exhausted initial Mo reservoir [ka BP] 122.3 calculated by Mo accumulation rates (cumulative)

152 Redox evolution during Eemian and Holocene sapropel formation in the Black Sea

References

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154 CONCLUSIONS AND OUTLOOK

CONCLUSIONS AND OUTLOOK

By using a multiproxy approach, the present work focusses on the climatic and environmental changes of the Black Sea over the last two glacial-interglacial cycles.

Cold annual mean temperatures ranged between 8 and 11°C during the ending penultimate glacial (Termination II), whose instability prelude the following Eemian warming phase. Indeed, two extensive meltwater pulses originating from the disintegrating Fennoscandian Ice sheet but also from a potential Himalayan source demonstrate pronounced climate instability (Fig. 13). The glacial-interglacial transition towards the Eemian involved rapid warming of up to 22°C during summer, whereas temperatures towards the Holocene reached only 18°C.

The increasing global sea level during the interglacials resulted in the Mediterranean- Black Sea reconnection at ca. 128.1 ka BP during the Eemian and at ca. 9 ka BP during the Holocene (Fig. 13). Favoured by the at least 10 m higher early Eemian sea level (e.g., Grant et al., 2012) and associated enhanced inflow of saline Mediterranean waters, the Black Sea’s salinity exceeded Holocene values by up to 10 (Shumilovskikh et al., 2013b). The higher Eemian salinity most likely is the reason for a different redox evolution compared to the Holocene. While the Eemian sapropel documents an almost ideal sequence from suboxic to strong euxinia, the Holocene sapropel reflects redoxcline migrations and weaker euxinia.

Prominent climate fluctuations seen in the Northern Hemisphere during the last glacial and especially during Marine Isotope Stage 3 (57-29 ka BP) could be documented in the Black Sea “Lake”. Temperature reconstructions revealed amplitudes of up to 4°C, which clearly coincided with Dansgaard-Oeschger stadial-interstadial cycles (Fig. 13). The record, therefore, demonstrates for the first time a deep inland propagation of DO-related temperature changes driven by a close atmospheric teleconnection between the North Atlantic and Eurasia during MIS 3 and coevally confirms climate models proposing that the influence of Heinrich event extra cooling might have been weaker in the Black Sea region.

The DO cycles affected not only temperatures in the Black Sea region but also rainfall and wind patterns, which all influenced the environmental conditions in the isolated basin. The warmer and moister interstadials, favoured by a northward-migrated atmospheric polar front and by an increasing role of Mediterranean cyclones, were characterised by increased riverine freshwater supply, enhanced productivity, decreased salinity, and probably by an increasing lake level. In contrast, the colder and more arid stadials were affected by an

155 CONCLUSIONS AND OUTLOOK increasing role of polar air masses and strong westerly winds. The reduced rainfall, higher evaporation, and lower riverine freshwater discharge caused an increasing salinity and probably a decreasing lake level. Aridity, less dense vegetation cover, and stronger westerly winds additionally favoured aeolian input.

In a long-term perspective, winters were much milder during the early MIS 3 when compared to MIS 4 and MIS 2 suggesting a strong influence of insolation-driven changes in Eurasian ice volume and associated atmospheric circulation patterns. Thus, an intensified meridional atmospheric circulation is inferred during MIS 3, when Eurasian ice sheets retreated.

Fig. 13: Overview about some major findings obtained in the present study over the last two glacial-interglacial cycles. M-BS denotes the Mediterranean-Black Sea reconnections.

Although this thesis contributed new insights into the glacial and interglacial history of the Black Sea (“Lake”) during the last 134 ka, there are still several temporal gaps and scientific questions, which are of crucial importance:

First, there is obviously a need for sediment cores covering periods older than 134 ka BP to learn more about the climatic and environmental conditions during the penultimate glacial maximum (140 ka BP) and Black Sea sapropels formed during older interglacials.

156 CONCLUSIONS AND OUTLOOK

Such cores may possibly help to detect further meltwater pulses and their pathways as well as earlier Mediterranean-Black Sea connection phases around 170-190 ka BP and 220-230 ka BP as it has been postulated by Badertscher et al. (2011), and thus potential episodes of sapropel formation. New gravity cores recovered during R/V “Maria S. Merian” cruise MSM33 in 2013 may help to answer such questions.

Second, the sediment cores recovered in the SE Black Sea reveal a hiatus between ca. 120 and 65 ka BP (Fig. 13). Since this was the period where the last isolation from the Mediterranean Sea took place, this interval is of fundamental importance to understand the still unknown process of the Black Sea freshening and oxygenation during a transition from interglacial to glacial conditions. Unfortunately, this has not been recovered in gravity cores up to now and will require future careful sampling strategies in the Black Sea. A combination of carbonate shell geochemistry of marine bivalves (Mytilus galloprovincialis) and limnic ostracods (Candona spp.) as well as organic biomarkers (e.g. alkenones) could contribute to our understanding of the process of freshening but also of salinification. For this purpose, further analytical experiments on the behaviour of trace element geochemistry of bivalves and ostracod valves during salinity changes are essential for the development of suitable calibrations.

Third, since the Sr-isotope signature of ostracods from the penultimate glacial meltwater events in the SE Black Sea “Lake” point to a potential Himalayan source, additional Sr-isotope records from the last glacial meltwater event should be collected from the same region and the NE Black Sea to exclude other local high radiogenic Sr-sources. Such approach could be improved by clay mineralogy and heavy mineral assemblages, which may serve for provenance studies and thus for tracing meltwater pathways. Additionally, samples of suspended particulate matter and waters from the Black Sea, Caspian Sea, Aral Sea, and rivers like Amu Darya and Dnieper should be analysed for its Sr-isotope signatures.

The discrepancy of the Dansgaard-Oeschger related temperature changes, which are well documented in the SE Black Sea records but to date not detected in the NW Black Sea should attract thorough examination. In this regard, further analysis of TEX86-based palaeotemperatures from sediment cores throughout the Black Sea may help to understand in which extent the amount of terrestrial (organic matter) input and/or the heterogenic climate conditions may control the TEX86 palaeothermometer during the limnic stages of the Black Sea. This purpose would require parallel investigation of alternative temperature-related proxies (e.g. Mg/Ca ratios of ostracods for changes in bottom water temperatures, amount of

157 CONCLUSIONS AND OUTLOOK coastal ice-rafted detritus for estimating winter severity) but also of inorganic detrital geochemical parameters (e.g. K, Ti, Zr) for assessment of terrestrial input.

Finally, similar to records from Greenland and the North Atlantic millennial-scale climate variability could have also occurred during periods older than 64 ka, which should get attention during future studies.

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178 APPENDIX

APPENDIX

Related co-author publications and personal contribution

This chapter presents three co-author publications related to the present study and describes the personal contribution of the author of the present thesis.

The first co-author publication “Vegetation and environmental changes in Northern Anatolia between 134 and 119 ka recorded in Black Sea sediments” (L. S. Shumilovskikh, H. W. Arz, A. Wegwerth, D. Fleitmann, F. Marret, N. Nowaczyk, P. Tarasov, and H. Behling; Quaternary Research (2013), 80, 3, 349-360) presents the dynamics of changes in vegetation in Northern Anatolia and the change from freshwater to marine dinoflagellate dominance during Termination II and the Eemian Mediterranean-Black Sea reconnection. Personal contribution included measurements and analyses of TIC and involvement during writing of the manuscript.

The second co-author publication “Orbital- and millennial-scale environmental changes between 64 and 20 ka BP recorded in Black Sea sediments” (L. S. Shumilovskikh, D. Fleitmann, N. Nowaczyk, H. Behling, F. Marret, A. Wegwerth, and H. W. Arz; Climate of the Past (2014), 10, 939-954) provides new insights into the vegetation and environmental changes during long-term and millennial-scale climate variability during the last glacial in the Black Sea “Lake”. Personal contribution involved sampling of the sediment core 25GC-1 for palynological analyses as well as involvement during writing of the manuscript.

The third co-author publication “Environmental and Climate Dynamics During the Last Two Glacial Terminations and Interglacials in the Black Sea/Northern Anatolian Region” (H. W. Arz, L. S. Shumilovskikh, A. Wegwerth, D. Fleitmann, and H. Behling; Integrated Analysis of Interglacial Climate Dynamics (INTERDYNAMIC). M. Schulz and A. Paul, Springer International Publishing, SpringerBriefs in Earth System Sciences, (2015), 121-126) represents an overview of the climate and environmental variability during the last two glacial-interglacial transitions in the Black Sea and is based on the results obtained within the framework of the SPP subproject “Dynamics of Mid-latitude/Mediterranean climate during the last 150 ka: Black Sea/Northern Anatolian Palaeoenvironmental Reconstructions (DynNAP)”. Personal contribution involved interpretation of data as well as participation during writing of the manuscript.

179 Vegetation and environmental changes in Northern Anatolia between 134 and 119 ka recorded in Black Sea sediments

Vegetation and environmental changes in Northern Anatolia between 134 and 119 ka recorded in Black Sea sediments

L. S. Shumilovskikha*, H. W. Arzb, A. Wegwerthb, D. Fleitmannc,d,h, F. Marrete, N. Nowaczykf, P. Tarasovg, and H. Behlinga

a Department of Palynology and Climate Dynamics, University of Göttingen, Untere Karspüle 2, 37073 Göttingen, Germany b Leibniz Institute for Baltic Sea Research Warnemünde, Seestrasse 15, 18119 Rostock- Warnemünde, Germany c Department of Archaeology, School of Archaeology, Geography and Environmental Science, University of Reading, Whiteknights, PO Box 227, Reading RG6 6AB, UK d Institute of Geological Sciences, University of Bern, Baltzerstrasse 3, 3012 Bern, Switzerland e School of Environmental Sciences, University of Liverpool, Liverpool L69 7ZT, UK f Helmholtz Center Potsdam GFZ German Research Centre for Geosciences, Telegrafenberg, 14473 Potsdam, Germany g Institute of Geological Sciences, Palaeontology Section, Freie University Berlin, Malteserstr. 74-100, House D, Berlin 12249, Germany h Oeschger Centre for Climate Change Research, University of Bern, Baltzerstrasse 3, 3012 Bern, Switzerland

180 Vegetation and environmental changes in Northern Anatolia between 134 and 119 ka recorded in Black Sea sediments

Abstract

This multiproxy study on SE Black Sea sediments provides the first detailed reconstruction of vegetation and environmental history of Northern Anatolia between 134 and 119 ka. Here, the glacial–interglacial transition is characterized by several short-lived alternating cold and warm events preceding a meltwater pulse (~130.4– 131.7 ka). The latter is reconstructed as a cold arid period correlated to Heinrich event 11. The initial warming is evidenced at ~130.4 ka by increased primary productivity in the Black Sea, disappearance of ice-rafted detritus, and spreading of oaks in Anatolia. A Younger Dryas-type event is not identifiable. The Eemian vegetation succession corresponds to the main climatic phases in Europe: i) the Quercus–Juniperus phase (128.7–126.4 ka) indicates a dry continental climate; ii) the Ostrya–Corylus–Quercus–Carpinus phase (126.4–122.9 ka) suggests warm summers, mild winters, and high year-round precipitation; iii) the Fagus–Carpinus phase (122.9– 119.5 ka) indicates cooling and high precipitation; and iv) increasing Pinus at ~121 ka marks the onset of cooler/ drier conditions. Generally, pollen reconstructions suggest altitudinal/latitudinal migrations of vegetation belts in Northern Anatolia during the Eemian caused by increased transport of moisture. The evidence for the wide distribution of Fagus around the Black Sea contrasts with the European records and is likely related to climatic and genetic factors.

Keywords

Pollen, Dinoflagellate cysts, Fagus, Vegetation history, Paleoclimate, Black Sea, Penultimate glacial, MIS 5e, Eemian

181 Orbital- and millennial-scale environmental changes between 64 and 20 ka BP recorded in Black Sea sediments

Orbital- and millennial-scale environmental changes between 64 and 20 ka BP recorded in Black Sea sediments

L. S. Shumilovskikh1*, D. Fleitmann2,3,4, N. R. Nowaczyk5, H. Behling1, F. Marret6,

A. Wegwerth7, and H. W. Arz7

1Department of Palynology and Climate Dynamics, University of Göttingen, Göttingen, Germany

2Department of Archaeology, School of Archaeology, Geography and Environmental Science, University of Reading, Reading, UK

3Institute of Geological Sciences, Bern, Switzerland

4Oeschger Centre for Climate Change Research, University of Bern, Bern, Switzerland

5Helmholtz Center Potsdam GFZ German Research Centre for Geosciences, Potsdam, Germany

6School of Environmental Sciences, University of Liverpool, Liverpool, UK

7Leibniz Institute for Baltic Sea Research Warnemünde, Rostock, Germany

Published in Climate of the Past, 10, 939–954 (2014) doi:10.5194/cp-10-939-2014 © Author(s) 2014. CC Attribution 3.0 License. Published by Copernicus Publications on behalf of the European Geosciences Union.

182 Orbital- and millennial-scale environmental changes between 64 and 20 ka BP recorded in Black Sea sediments

Abstract

High-resolution pollen and dinoflagellate cyst records from sediment core M72/5-25- GC1 were used to reconstruct vegetation dynamics in northern Anatolia and surface conditions of the Black Sea between 64 and 20 ka BP. During this period, the dominance of Artemisia in the pollen record indicates a steppe landscape and arid climate conditions. However, the concomitant presence of temperate arboreal pollen suggests the existence of glacial refugia in northern Anatolia. Long-term glacial vegetation dynamics reveal two major arid phases ~ 64–55 and 40–32 ka BP, and two major humid phases ~ 54–45 and 28–20 ka BP, correlating with higher and lower summer insolation, respectively. Dansgaard–Oeschger (D–O) cycles are clearly indicated by the 25-GC1 pollen record. Greenland interstadials are characterized by a marked increase in temperate tree pollen, indicating a spread of forests due to warm/wet conditions in northern Anatolia, whereas Greenland stadials reveal cold and arid conditions as indicated by spread of xerophytic biomes. There is evidence for a phase lag of ~ 500 to 1500 yr between initial warming and forest expansion, possibly due to successive changes in atmospheric circulation in the North Atlantic sector. The dominance of Pyxidinopsis psilata and Spiniferites cruciformis in the dinocyst record indicates brackish Black Sea conditions during the entire glacial period. The decrease of marine indicators (marine dinocysts, acritarchs) at ~ 54 ka BP and increase of freshwater algae (Pediastrum, Botryococcus) from 32 to 25 ka BP reveals freshening of the Black Sea surface water. This freshening is possibly related to humid phases in the region, to connection between Caspian Sea and Black Sea, to seasonal freshening by floating ice, and/or to closer position of river mouths due to low sea level. In the southern Black Sea, Greenland interstadials are clearly indicated by high dinocyst concentrations and calcium carbonate content, as a result of an increase in primary productivity. Heinrich events show a similar impact on the environment in the northern Anatolia/Black Sea region as Greenland stadials.

183 Environmental and Climate Dynamics During the Last Two Glacial Terminations and Interglacials in the Black Sea/Northern Anatolian Region

Environmental and Climate Dynamics During the Last Two Glacial Terminations and Interglacials in the Black Sea/Northern Anatolian Region

H. W. Arz1*., L. S. Shumilovskikh2, A. Wegwerth1, D. Fleitmann3,4, and H. Behling2

1 Leibniz Institute for Baltic Sea Research Warnemünde, Rostock-Warnemünde, Germany

2 Department of Palynology and Climate Dynamics, Georg-August-University Göttingen, Göttingen, Germany

3 Institute of Geological Sciences and Oeschger Centre for Climate Change Research, University of Bern, Bern, Switzerland

4 Department of Archaeology, School of Human and Environmental Sciences, University of Reading, Reading, UK

Published in M. Schulz and A. Paul (eds.), Integrated Analysis of Interglacial Climate Dynamics (INTERDYNAMIC), SpringerBriefs in Earth System Sciences, pp 121-126 (2015) DOI 10.1007/978-3-319-00693-2_20 © The Author(s) 2015 Springer International Publishing

184

Abstract

This study provides the first detailed multi-proxy paleoenvironmental reconstructions of changes in the aquatic and terrestrial ecosystems during the Holocene, Eemian and the last two glacial/interglacial transitions (Terminations I and II) by studying sediment cores from the southeastern Black Sea and stalagmite studies from Sofular Cave in northwestern Anatolia. The terrestrial proxies document gradual changes from late glacial cold/arid conditions in northern Anatolia, dominated by steppe vegetation, to warm/humid forest dominated landscapes characteristic for interglacial periods. The Holocene and Eemian, however, developed differently, with warmer and moister conditions prevailing during the Eemian. Major fluctuations in the hydrological state of the Black Sea are closely linked to changes of terrestrial environments. Disrupted by large melt water pulses from the disintegrating Fennoscandian Ice Sheet, the limnic glacial Black Sea environment becomes more productive during the postglacial warming. Global sea-level rise finally reconnects the hydrological increasingly active Black Sea basin with the Mediterranean Sea leading to the development of marine, for the Eemian even fully marine, conditions with a stratified water column and sapropelic sedimentation.

Keywords

Black sea, Eemian, Holocene, Glacial-interglacial transitions, Paleoenvironment, Paleoclimatology, Palynology, Geochemistry

185

Eigenständigkeitserklärung

Hiermit erkläre ich, dass diese Arbeit bisher von mir weder an der Mathematisch- Naturwissenschaftlichen Fakultät der Ernst-Moritz-Arndt-Universität Greifswald noch einer anderen wissenschaftlichen Einrichtung zum Zwecke der Promotion eingereicht wurde.

Ferner erkläre ich, dass ich diese Arbeit selbstständig verfasst und keine anderen als die darin angegebenen Hilfsmittel und Hilfen benutzt und keine Textabschnitte eines Dritten ohne Kennzeichnung übernommen habe.

Rostock, 05.11.2015 Antje Wegwerth

186 CURRICULUM VITAE

CURRICULUM VITAE

Antje Wegwerth

2011 – 2014 PhD student at the Leibniz Institute for Baltic Sea Research (IOW) within the framework of the project DynNAP “Dynamics of Mid- latitude/Mediterranean climate during the last 150 ka: Black Sea/Northern Anatolian Paleoenvironmental Reconstructions” as part of the DFG Priority Programme 1266 "Integrated Analysis of Interglacial Climate Dynamics (INTERDYNAMIC)"

2012 and 2013 Biomarker studies at CEREGE, Centre Européen de Recherche et d’Enseignement des Géosciences de l’Environnement, Géochimie et géochronologie, Aix-Marseille University, France (in total two months)

2011 Diploma-Mapping Thesis „Fluxes of dissolved manganese and sulphate across the sediment-water interface of the Baltic Sea“ (IOW/FU Berlin)

2010 Diploma Thesis “Petrography, Biogeochemistry, and Stable Isotope Signatures of Early Cambrian Phosphorites, Meishucun Section, South China“ within DFG Forschergruppe 736 “The Precambrian-Cambrian Ecosphere (R)evolution: Insights from Chinese microcontinents” (FU Berlin/IOW)

2008 – 2011 traineeship and student assistant at the IOW in the Marine Geology department

2007 – 2008 student assistant at the Freie Universität Berlin, Institute of Geological Sciences, Paleontology

2006 traineeship at the Senckenberg Research Institute and Natural History Museum Frankfurt, Messel Pit Fossil Site (one month)

2003 – 2011 Study of Geology-Paleontology at the Freie Universität Berlin

Rostock, 05.11.2015 Antje Wegwerth

187 LIST OF PUBLICATIONS AND OTHER SCIENTIFIC ACTIVITIES

LIST OF PUBLICATIONS AND OTHER SCIENTIFIC ACTIVITIES

Publications

• Wegwerth, A., Ganopolski, A., Ménot, G., Kaiser, J., Dellwig, O., Bard, E., Lamy, F., and Arz, H.W. Black Sea temperature response to glacial millennial-scale climate variability. Geophysical Research Letters (2015), 42, 8147-8154, 2015GL065499 http://dx.doi.org/10.1002/2015GL065499

• Arz, H. W., Shumilovskikh, L. S., Wegwerth, A., Fleitmann, D., Behling, H. Environmental and Climate Dynamics During the Last Two Glacial Terminations and Interglacials in the Black Sea/Northern Anatolian Region. Integrated Analysis of Interglacial Climate Dynamics (INTERDYNAMIC). M. Schulz and A. Paul, Springer International Publishing, SpringerBriefs in Earth System Sciences, (2015), 121-126, http://dx.doi.org/10.1007/978-3-319-00693-2_20

• Wegwerth, A., Dellwig, O., Kaiser, J., Ménot, G., Bard, E., Shumilovskikh, L., Schnetger, B., Kleinhanns, I.C., Wille, M., and Arz, H.W. Meltwater events and the Mediterranean reconnection at the Saalian-Eemian transition in the Black Sea. Earth and Planetary Science Letters (2014), 404, 124-135, http://dx.doi.org/10.1016/j.epsl.2014.07.030

• Shumilovskikh, L.S., Fleitmann, D., Nowaczyk, N., Behling, H., Marret, F., Wegwerth, A., Arz, H.W. Orbital- and millennial-scale environmental changes between 64 and 20 ka BP recorded in Black Sea sediments. Climate of the Past (2014), 10, 939-954, http://www.clim-past.net/10/939/2014/cp-10-939-2014.html

• Shumilovskikh, L. S., Arz, H.W., Wegwerth, A., Fleitmann, D., Marret, F., Nowaczyk, N., Tarasov, P., Behling, H., Vegetation and environmental changes in Northern Anatolia between 134 and 119 ka recorded in Black Sea sediments, Quaternary Research (2013), 80, 3, 349-360, http://dx.doi.org/10.1016/j.yqres.2013.07.005

Manuscripts submitted/in preparation

• Wegwerth, A., Kaiser, J., Dellwig, O., Shumilovskikh, L.S., Nowaczyk, N.R., and Arz, H.W. Northern hemisphere climate control on the environmental dynamics in the glacial Black Sea “Lake”. under review in Quaternary Science Reviews

• Häusler, K., Dellwig, O., Schnetger, B., Leipe, T., Moros, M., Pohl, C., Pollehne, F., Wegwerth, A., and Arz, H.W. Exceptional manganese authigenesis in deep basins of the Baltic Sea: Temporal allocation, hydrographic prerequisites and budget calculations. under review in Geochimica et Cosmochimica Acta

188 LIST OF PUBLICATIONS AND OTHER SCIENTIFIC ACTIVITIES

• Wegwerth, A., Eckert, S., Dellwig, O., Köster, J., Schnetger, B., Severmann, S., Weyer, S., Brüske, A., Kaiser, J., Arz, H.W., and Brumsack, H.-J.. Redox evolution during Eemian and Holocene sapropel formation in the Black Sea. in preparation

Presentations

• Lipka, M., Liu, B., Wegwerth, A., Dellwig, Winde, V., Al-Raei, A. M., and Böttcher, M. E. (2015) Carbon transformation and the sources of dissolved inorganic carbonate in sediments of a temperate coastal sea, the Baltic Sea: A stable isotope and modelling approach, EGU General Assembly, April 2015, Vienna, Austria

• Lipka, M., Wegwerth, A., Dellwig, Winde, V., Al-Raei, A. M., and Böttcher, M. E. (2015) Element transformation rates and fluxes across the sediment-water interface in a temperate coastal sea, the Baltic Sea, 2015 Aquatic Sciences Meeting, Aquatic Sciences: Global And Regional Perspectives - North Meets South, ASLO 2015, February 2015, Granada, Spain

• Lipka, M., Wegwerth, A., Dellwig, O., Al-Raei, A. M., Schoster, F., and Böttcher, M. E. (2014) Element transformation rates and fluxes across the sediment-water interface of the Baltic Sea, The 12th Colloquium on Baltic Sea Marine Geology, BALTIC 2014, September 2014, Rostock-Warnemünde, Germany

• Arz, H.W., Shumilovskikh, L., Wegwerth, A., Fleitmann, D., Kaiser, J., Nowaczyk, N., Behling, H., Dellwig, O., Ménot, G., Bard, E. (2014) Environmental changes in the Black Sea: A Mediterranean - Black Sea connection Perspective, Interdisciplinary meeting on climate change and seismic hazards during the Holocene in the mediterranean, July 2014, Aix-en-Provence, France

• Lipka, M., Wegwerth, A., Dellwig, O., Al-Raei, A., Schoster, F., and Böttcher, M.E. (2014) Element transformation rates and fluxes across the sediment-water interface of the Baltic Sea, EGU General Assembly, April/May 2014, Vienna, Austria

• Arz, H.W., Shumilovskikh, L., Wegwerth, A., Dellwig, O., Kaiser, J., Nowaczyk, N., Plessen, B., Prange, M., Ménot, G., Bard, E., Behling, H., Fleitmann, D. (2013) Pronounced millennial scale variability in the Black Sea Lake during Marine Isotope Stage 3. 14th RCMNS, September 2013, Istanbul, Turkey

• Wegwerth, A., Dellwig, O., Kaiser, J., Bard, E., Ménot, G., Nowaczyk, N., Plessen, B., Shumilovskikh, L., Fink, S., Arz, H. W. (2013) The impact of D-O Climate Oscillations during the last Glacial in the Black Sea. 11th INTERNATIONAL CONFERENCE ON PALEOCEANOGRAPHY, September, 2013, Sitges - Barcelona, Spain

• Wegwerth, A., Dellwig O., Kaiser, J., Bard ,E., Ménot, G., Nowaczyk, N., Plessen, B., Schnetger, B., Shumilovskikh, L., and Arz, H. (2013) Palaeoceanography & Palaeoclimate during the penultimate Glacial-Interglacial transition in the Black Sea - Termination II. EGU General Assembly, April 2013, Vienna, Austria

189 LIST OF PUBLICATIONS AND OTHER SCIENTIFIC ACTIVITIES

• Wegwerth, A., Struck, U., Segl, M., Vennemann, T.W., Gehlken, P.L., Heubeck, C., and Böttcher M.E. (2010) Stable isotope record of coexisting apatite and dolomite in Early Cambrian phosphorites, Meishucun section, South China. EGU General Assembly, May 2010, Vienna, Austria

• Wegwerth, A., Dellwig, O., Struck, U., Hippler, D., Gehlken, P.-L., Trappe, J., Franz, G., Böttcher, M.E., and Heubeck, C. (2009) Biogeochemistry and petrography of Early Cambrian phosphorites from the Meishucun section, Yangtze Platform, South China. GV Annual Meeting, October 2009, Göttingen, Germany

Research cruises

• 2013 RV "Maria S. Merian": MSM33, Black Sea • 2012 RV “Elisabeth Mann Borgese“: 06EZ1215-REDOX, Central Baltic Sea • 2009 RV „Südfall“: North Sea • 2008 RV „Prof. A. Penck“: 07/08/24 Central Baltic Sea • 2008 RV „Prof. A. Penck“: 07/08/14 Central Baltic Sea

190 ACKNOWLEDGEMENTS

ACKNOWLEDGEMENTS

First of all, I wish to thank my supervisor Helge W. Arz for enabling the study and activities around the Black Sea history and for giving me the freedom to find my own way into the Quaternary. He always offered great help and advice at the right moment and I am grateful for his inspiring way of thinking. Thank you for letting me travel into the ice age world. I appreciate the enduring support by Olaf Dellwig during development of methods for inorganic geochemical analyses on the ostracods and sediments and I am thankful for his advisory support during manuscript writing, for his contagious intrinsic motivation, and for bringing me sometimes down to earth. I kindly acknowledge Jérôme Kaiser for his instructive discussions about organic biomarkers and Quaternary history. Lyudmila S. Shumilovskikh is gratefully acknowledged for her infectious enthusiasm and kind support during our Black Sea cruise and conversations. I am thankful to all the co-authors, who greatly contributed to the manuscripts. Martin Wille, thank you for encouraging the fruitful Sr-isotope analyses. My colleagues and friends of the Marine Geology Section at IOW and my thesis committee are acknowledged for being available for any advice at all times. Many thanks also to Dagmar Benesch, Anne Köhler, Nadine Hollmann, Ines Scherff, Sascha Plewe, Olga Körting, Katharina Häusler, and Sibylle Fink for kind assistance during work in laboratory and onboard. I thank Hans-Jürgen Brumsack, Bernhard Schnetger, and Sebastian Eckert for their confidence and sharing valuable unpublished data. This study also greatly benefited from analyses of biomarkers and friendly assistance by Yoann Fagault during two stays at Cerege as well as helpful comments by Guillemette Ménot and Edouard Bard. Thank you, Olga, for your permanent support during literature research. I am much indebted to our ‘team Gitte’ for spending some lively times during pauses for thoughts. Thank you, Katharina, for your word of encouragement and kind atmosphere everyday. I also gratefully acknowledge my family and friends for mental support during this time and for letting me write this work during moments we usually spend together. Thank you, Felix, for your inspiring and lively interest in natural sciences. Thank you, Olaf, for your empathy from every point of view, for uncountable inspiring long-playing record evenings, for our sorted chaos during drafting ideas and pieces of papers, for your patience and impatience just at the right moment, and for being with me. Finally, I wish to thank the captain and crew of RV “Meteor” for their support during the M72/5 Black Sea cruise in 2007. I am also grateful for participating during the Black Sea cruise with RV “Maria S. Merian” in 2013. This work was funded by the Deutsche Forschungsgemeinschaft (BE2116/20-1, AR 367/9-1, AR 367/9-2, and FL 710/1-1) within the priority program 1266 INTERDYNAMIC, by the Gary Comer Science and Education Foundation, by the diverse institutions of the co-authors, and by the Leibniz Institute for Baltic Sea Research.

191