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Quaternary Science Reviews 28 (2009) 2882–2896

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Quaternary Science Reviews

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Climatic variability during the last w90 ka of the southern and northernLevantine Basin as evident from marine records and speleothems

Ahuva Almogi-Labin a,*, Miryam Bar-Matthews a, Dan Shriki a,b, Elina Kolosovsky c, Martine Paterne d,Bettina Schilman a, Avner Ayalon a, Zeev Aizenshtat c, Alan Matthews b

a Geological Survey of Israel, 30 Malchei Yisrael St., Jerusalem 95501, Israelb Earth-Sciences Institute, Hebrew University, Jerusalem 91904, Israelc Institute of Chemistry, Hebrew University, Jerusalem 91904, Israeld LSCE -CNRS-CEA Avenue de la Terrasse 91198 Gif sur Yvette, France

a r t i c l e i n f o

Article history:Received 23 March 2009Received in revised form27 July 2009Accepted 28 July 2009

* Corresponding author. Tel.: þ972 2 531 4232; faxE-mail addresses: almogi@gsi.gov.il (A. Almogi-

(M. Paterne), bschilman@gsi.gov.il (B. Schilman), a

0277-3791/$ – see front matter � 2009 Elsevier Ltd.doi:10.1016/j.quascirev.2009.07.017

a b s t r a c t

The influence of the northern Atlantic and tropical monsoonal systems, as recorded by the River Nile, onthe climate variability of the southeastern Mediterranean was studied in two cores taken by the R/VMarion Dufresne: one core taken SE of Cyprus representing the northern Levantine Basin (core 9501,980 m water depth) and the other located 380 km further south, represents the southeastern LevantineBasin in an area influenced by the River Nile plume (core 9509, 884 m water depth). The study wasperformed at relatively high resolution using several proxies: d18O of Globigerinoides ruber, sedimentcharacteristics and index colour parameters in core sections representing the last 86 ka. A low-resolutionalkenone sea surface temperature record was also measured. The time frame in both cores was mostlyconstrained by ‘wiggle’ matching with the nearby well-dated d18O and d13C record of the Soreq Cave,which is mainly influenced by the eastern Mediterranean water vapor. The sedimentary record of thesouthern core is strongly influenced by the River Nile contribution throughout the last 86 ka, as evi-denced by the higher sedimentation rates compared with the northern core (20 cm/ka vs. 5 cm/ka),continuously darker sediment colour, and higher TOC values (0.6–0.9 vs. 0.25 wt% not including sap-ropels). During sapropels S1 and S3, present in both cores, the influence of the River Nile became morewidespread, reaching as far as Cyprus. Yet, the influence of the River Nile remained stronger in the south,as evident by the higher TOC values in the southern core throughout the entire 90 ka period and thelonger duration of S1 in the southern core. An anomalous low d18O interval that is not recorded inwestern Mediterranean occurred between 58 and 49 ka in the Levantine Basin and is more developed inthe northern core. This period correlates with D-O interstadial 14 and maximum northern hemisphereinsolation during the lastglacial cycle, suggesting that the warming mainly impacted the northernLevant.

The Eastern Mediterranean Sea and land area was considerably warmer than the western Mediter-ranean throughout the LGM – Holocene transition, and the d18OG. ruber drop of 4.5& is significantlygreater than the 3& shift found for the western Mediterranean d18OG. bulloides, both differences reflectingan increased continental effect from the western to eastern Mediterranean. Comparison between themarine and the land d18O records suggests that the origin of rain over the land is composed of mixedsignal from the southern and northern Levantine Basin. The study of Dd18Osea–land variations demon-strates that various factors have influenced the sea–land relationship during the last 90 ka. The ‘amounteffect’ has an important influence on rainfall d18O during interglacial periods (particularly sapropelperiods), whereas during glacial periods, increased land distances and elevation differences arising fromdecrease in sea level may have brought about decrease in d18O of rainfall due to Rayleigh distillationprocesses. These influences were superimposed on those of sea surface water d18O changes broughtabout by continental ice melting, and the strong effects felt in the southern Levantine Basin of the highRiver Nile input during periods of enhanced monsoonal activity.

� 2009 Elsevier Ltd. All rights reserved.

: þ972 2 538 0688.Labin), matthews@gsi.gov.il (M. Bar-Matthews), danch@pob.huji.ac.il (D. Shriki), martine.paterne@lsce.cnrs-gif.fryalon@gsi.gov.il (A. Ayalon), zeev@vms.huji.ac.il (Z. Aizenshtat), alan@vms.huji.ac.il (A. Matthews).

All rights reserved.

A. Almogi-Labin et al. / Quaternary Science Reviews 28 (2009) 2882–2896 2883

1. Introduction

The almost landlocked Mediterranean Sea maintains its soleconnection to the Atlantic Ocean via the Gibraltar Strait. As a result,sea surface conditions are diverse with west to east salinity and seasurface temperature (SST) increase and a corresponding produc-tivity decrease (Antoine et al., 1995; Malanotte-Rizzoli et al., 1999;Pinardi and Masetti, 2000). The higher salinity and water density inthe Levantine Basin result in an increase in sea surface d18O valuesby up to w1& relative to the western Mediterranean (Pierre, 1999).This sharp W–E gradient reflects strong evaporation of theincoming Atlantic Surface water flowing eastward.

The Levantine Basin is mainly influenced by rainfall fronts thatoriginate in the NE Atlantic Ocean, passing over Europe and theMediterranean Sea (Rindsberger et al., 1983; Eshel, 2002); occa-sionally, rainfall is associated with sub-tropic warm air passingalong the southern Mediterranean (Dayan, 1986). Cold dry Arctic airthat penetrates into the eastern Mediterranean region results in anincrease in surface water density, leading to the formation of theLevantine Intermediate Water (LIW) off the island of Rhodes(Malanotte-Rizzoli et al., 1999; Pinardi and Masetti, 2000). Themonsoonal system that originates in the tropical Atlantic or thesouthern Indian Ocean, passes over NE Africa and is associated withthe low-latitude rainfall system (cf. Rossignol-Strick, 1985) thatindirectly influences the hydrography of the Levantine Basin. Thissystem fluctuates in time, and is at its highest strength duringperiods of June maximum insolation in the northern Hemispherethat coincide with elevated River Nile discharge (Rossignol-Stricket al., 1982; Rossignol-Strick, 1985; Haynes, 1987).

The Eastern Mediterranean-Levant region is located at themeeting of the Eurasian continent, the Saharan-Arabian desert andMediterranean Sea. Many studies on the paleoclimate of this regionshow that the Levantine Basin follows the global record but ischaracterized by – larger amplitude – oxygen isotope variations.A difference of ˜5& in d18O values of planktonic foraminifera isfound between the lastglacial maximum and the period of accu-mulation of interglacial sapropels in the deep sea (Fontugne andCalvert 1992; Kallel et al., 1997). This large isotopic shift is alsorecorded on land, as evident from the speleothem d18O records ofthe Eastern Mediterranean caves (e.g. Bar-Matthews et al., 2000,2003). The large glacial–interglacial amplitude not only reflects theice-volume effect, but also to a considerable extent, the semi-enclosed continental conditions of the region. The particularlyimportant role of the freshwater River Nile input on the LevantineBasin was recognized during sapropel periods (e.g. Rossignol-Stricket al., 1982; Rossignol-Strick, 1985; Fontugne et al., 1994; Rohling,1994; Calvert and Fontugne, 2001; Scrivner et al., 2004). Additionalsources of freshwater into the Eastern Mediterranean during sap-ropel periods are rivers from the Tibesti Mountains, which formedas a result of northward shift of the monsoonal belt (Rohling et al.,2002a,b; Osborne et al., 2008) and rainfall on the entire Mediter-ranean Sea from Atlantic sources (Kallel et al., 1997; Bar-Matthewset al., 2000). Other studies have described the influence ofincreased rainfall in the entire Mediterranean Sea during sapropelformation periods (Kallel et al., 1997; Bar-Matthews et al., 2000),the influence of periodic cold northern Atlantic fronts in thenorthern Levantine Basin (Rohling et al., 2002a,b) and the influx offresher Black Sea water (Marino et al., 2007).

The most accurately dated high resolution long-term climaterecord of the Eastern Mediterranean region is that derived fromSoreq Cave (Israel) speleothems. These show similar isotopic trendsand amplitude as the eastern Levantine Basin foraminifera (Bar-Matthews et al., 2000, 2003), thus implying that the speleothemsmainly record the Atlantic-Mediterranean cyclonic systems. Thisconclusion is supported by d18O-dD relationships of the fluid

inclusions trapped within the speleothems, which follow theMediterranean Meteoric Water Line (Matthews et al., 2000;McGarry et al., 2004). However, the role of the monsoonal vs. theAtlantic systems on the paleoenvironmental conditions of theLevantine Basin has yet been not fully explored.

The objective of this research is to perform a high resolutionstudy of the oxygen and carbon isotopic composition of theplanktonic foraminifera Globigerinoides ruber, together with a lowerresolution sea surface temperature (SST) study (based on thealkenone unsaturation index Uk

37), on two cores located inthe southern and northern parts of the Levantine Basin (Fig. 1). Thesouthern core (9509) is located beneath the River Nile plume,whose sources are dominantly influenced by the monsoonalsystem: the northern core (9501) is located w380 km further north,SE of Cyprus, and represents the northern Levantine Basin. Corre-spondingly, the southern core is situated closer to the vast Sahara-Arabian desert belt; whereas the northern core is situated whereclimatic conditions are less arid and more typical of the Mediter-ranean type climate. The overall time period covered by these coresis ca 90 ka to the present and includes: two sapropels (S3 and S1),the lastglacial, lastglacial maximum (LGM), deglaciation, and theHolocene. For the purposes of accurate dating of these periods andunraveling the sea–land climate connections, our study placesstrong emphasis on comparison of the isotopic and SST records ofthe two cores with the isotopic and land temperature records of theSoreq Cave speleothems. The aims of the study are to explore thefollowing key issues: the differences in SST and water d18O valuesbetween the northern and southern parts of the Levantine Basinrecorded by the two cores with their climate implications; the sea–land temperature and d18O relationships and their implications forthe isotope hydrology and climatic controls of rainfall systems inthe region; and finally the correlations between the two marinerecords and the isotopic and SST records of western Mediterraneanand global climate change.

2. Material and methods

The two cores, 9501 and 9509, were taken from the VALPAMEDset of cores, collected by the R/V Marion Dufresne during February1995 (Fig. 1). Core 9509 was taken under the River Nile plume in thesoutheastern Levantine Basin at 32�01 N, 34�16 E and 884 m waterdepth. Core 9509 was sampled along its total 17.8 m length in 10 cmintervals and in sapropel layers we sampled every 5 cm. Core 9501was taken SE of Cyprus at 34�32 N, 33�59 E and 980 m water depth inthe northeastern Levantine Basin. This core represents a longer timeperiod, at least the last 250 ka BP based on the presence of 8 distinctsapropels. It was sampled continuously in 2 cm intervals along itstop 5 m (out of total length of 11 m) and in 1 cm thick intervals in thesapropel layers. The colour index of the sediments in the two coreswas measured with a spectrophotometer colour-guide 45/0according to the method described by Chapman and Shackleton(1998). Freeze-dried samples were weighed and placed into poly-ethylene bags and the reflectance was measured at 10 nm incre-ments in the 400–700 nm wavelength range. Quantitative opticalinformation was derived from the differential reflectance withrespect to the wavelength (Balsam and Deaton, 1991). The L*, a*, b*parameters (as defined by the Commission Internationale del’Eclaraigh) provide information about the lightness and the sedi-ment colour: L* characterizes the lightness (approximately equiva-lent to gray scale reflectance) scaled from 0 (black) to 100 (white).Parameters a* and b* are chromaticity coordinates that define objectcolour: a* measures the colour shift from red to green and b*measures the colour shift from yellow to blue. As a* and b* valuesdeviate from zero, the degree of colour saturation becomes moreintense. These parameters are designed in relation to human colour

A. Almogi-Labin et al. / Quaternary Science Reviews 28 (2009) 2882–28962884

perception, but provide a useful synthesis of the complex datacollected by the spectrophotometer.

Sediment samples were washed through a 63 mm sieve and driedat 50 �C. Specimens of the planktonic foraminifera G. ruber werehandpicked from the 250–500 mm size fraction and ultrasonicallycleaned prior to measurement of d18O and d13C values. The stableisotope analyses were performed only on the thin-walled whitevariety of G. ruber. d18O and d13C values of core 9501 were deter-mined at the Geological Survey of Israel on an Isocarb systemattached to the VG SIRA-II Mass Spectrometer after reacting thesamples with 100% H3PO4. d18O and d13C values of G. ruber from core9509 were analyzed at the Laboratorie de Sciences du Climat et del‘environment (CEA-CNRS), Gif-sur-Yvette, France, using a carousel-48 automatic carbonate preparation device attached to FINNIGANMAT 251 mass spectrometer. Calibration of the mass spectrometry inboth laboratories was made relative to the VPDB scale using theinternational NBS 19 carbonate standard (Coplen, 1988).

1 g of dry sediment was taken for total organic carbon (TOC)analyses. The carbonate fraction was removed from the homoge-nized sediment using 50 ml HCl solution (10%) for 3 h while shaking/stirring. The resulting suspension was washed with ultra-pure waterto avoid CaCl2 contamination. The washing was ceased by checkingthe sediment extract with 0.1 N AgNO3. The TOC measurementswere performed using a ThermoFinnigan Elemental Analyzer (EA).

Thirty samples of dry sediments weighing 3–7 g were taken forthe purpose of reconstructing SST from selected time intervals

Fig. 1. (A) The area discussed in this study. (B) River Nile basin including its main tributaLevantine Basin, and core 9509 off southern Israel, under the Nile plume. The Soreq Cave i

including: recent samples from the surface, samples representingthe Holocene (MIS 1), glacial intervals (MIS 2–4), and from inter-glacial MIS 5 including sapropel layers. Past SST values weredetermined from measurements of the alkenone unsaturationindex Uk

37 as defined by Prahl et al. (1988) and Muller et al. (1998).Gas chromatography – mass spectrometry analyses were per-formed on an HP model 5890 II gas chromatograph directly coupledto a HP mass selective detector model G 1800-B. Capillary columnand temperature programming were the same as for the GC anal-ysis. Non-polar substances were eluted from the column in firstfraction, aromatic substances were eluted in second fraction, andpolar substances (including alkenones C37) in third fraction. Therelative abundance of biomarkers of organic matter origin wasdetermined by GC and GC-MS. Presence of high concentration ofshort-chain n-alkenes (C12–C19) indicates the marine origins of theorganic matter (Bouloubassi et al., 1997; Rullkotter et al., 1998)while the long-chain n-alkanes (C20–C40) are characteristic ofterrestrial origin (Eglinton and Hamilton, 1963).

The equations used for SST calculation based on alkenoneswere:

UK37 ¼ ð½C37 : 2� � ½C37 : 4�Þ=ð½C37 : 2� þ ½C37 : 3� þ ½C37 : 4�Þ

SST �C ¼�

Uk37 � 0:039

�.0:034

with an error of 0.5 �C (e.g., Prahl et al., 1988; Muller et al., 1998).

ries. (C) Map showing the locations of core 9501, SE of Cyprus, representing the NEs located on the western flanks of the Judean Mountains.

A. Almogi-Labin et al. / Quaternary Science Reviews 28 (2009) 2882–2896 2885

2.1. Dating of the cores

Following the studies of Bar-Matthews et al. (2000, 2003),which showed that the well-dated Soreq d18O cave records (by U/Thmethod) are well correlated with low resolution marine recordsfrom the Levantine Basin, we applied the same approach to thehigher resolution marine records of cores 9501 and 9509. Thecomparison between the land and the marine records was based onthe updated d18OSoreq record published in Vaks et al. (2006), basedon many more MC-ICP-MS U-Th ages and speleothems. The d13Cvalues of the speleothems were also used to constrain the timing ofthe marine records, mainly the timing of sapropel S1, because of theinformation they give on periods of intensive hydrological activity(Bar-Matthews et al., 2000). The use of speleothem d13C values toconstrain sapropel timing is based on the fact that enhancedhydrological activity in the eastern Mediterranean occurred inconcert with the monsoonal activity during the early Holocene(Kallel et al., 1997; Arz et al., 2003; Almogi-Labin et al., 2004).

Several other methods were used to constrain the marine timeframework. These include sediment characteristics, such as theidentification of sapropel events from the colour of the sediments,colour index and TOC content (Rossignol-Strick 1985; Calvert andFontugne, 2001). We also dated several intervals using AMS 14Cdating of planktonic foraminifera (made by L.K. Fifield from theAustralian National University in collaboration with M. Paterne)(Table 1). For this purpose, a few hundred specimens of G. ruberfrom core 9501 were collected in samples at seven different timeintervals and in one sample from near the surface, which includedmixed spinose planktonic foraminifera. The ages were converted tocalendar ages following Stuiver and Reimer (1993) and correctedfor a reservoir age of 400 years (Bard, 1988; Siani et al., 2000). Wecompared the calibration with the Soreq speleothem record andthe age calculations using Fairbanks et al. (2005) calibration curve.Generally there was a good agreement within the analytical error,but in two cases at depths of 84 and 142 cm, the difference exceedsthe analytical error. At 84 cm depth, there is a difference ofapproximately 500 y within the analytical error, and at 142 cmdepth the difference is nearly 1000 y. In these two cases we usedthe speleothem record for the correlation, since it better fitted withthe other independent constraints. With respect to the 14C age of37,790 y from the core depth of 322 cm, matching with the spe-leothem record and its location in the core relative to sapropellayers S1 and S3, and fluctuations in the sedimentation rates, allsuggest a much older age of about 54,120 y. This older age makesthe sedimentation rate more uniform and the G. ruber d18O valuesfit better with the speleothem record. The age of 54,120 y is stillwithin the range of correlation with 14C and U/Th as shown inFairbanks et al. (2005). This age offset matches that of samples from

Table 1AMS 14C ages of planktonic foraminifera in core 9501.

Depth (cm) 14C age (y) � 2s 14C agea (y) Correctedagesb (y)

Age incore (y)

Species

2–4 640 180 240 240 240 Mixed Plank54–56 4200 180 3800 4200 5000 G. ruber80–81 7410 190 7010 8000 8200 G. ruber84–85 7600 190 7200 8200 8700c G. ruber88–89 8420 190 8020 9200 9300 G. ruber142–144 13840 210 13440 16000 16880c G. ruber202–204 20100 260 19700 25000 25200 G. ruber322–324 37790 910 37390 50000 54120c G. ruber

a 14C ages: after conversion to calendar ages following Stuiver and Reimer (1993)and correction for a reservoir age of 400 years (Bard, 1988; Siani et al., 2000).

b Corrected ages: calibration after Fairbanks et al. (2005) curve (see text).c The age is after ‘wiggle matching’ with Soreq Cave record

Papua New Guinea (Yokoyama et al., 2000) and carbonate sedimentsamples from Lake Lisan (Schramm et al., 2000).

Core 9509 covers the last w90 ka and for this reason weconcentrated on the top 5 m in core 9501, which covers the sametime interval. The dating in core 9509 was done by wiggle matchingwith Soreq Cave record.

3. Results

3.1. Sediment characteristics

The sediments in the top 5 m of core 9501 are siliciclastics richin planktonic foraminifera. The colour varies between dark gray,light and dark brown with darkest colour intervals occurringbetween 448 and 494 cm and between 80 and 90 cm. The lowerdark interval is interrupted by a 3 cm thick interval of lightersediments at 462–465 cm. In contrast, the colour along the entirelength of core 9509 is dark gray and it was impossible to identifythe sapropel layers from colour alone.

Colour index parameters reveal significant differences betweenthe two cores. Generally, all colour index parameters in core 9501(Fig. 2A) are significantly higher compared to core 9509 (Fig. 2B).They also show a more complex structure in core 9501 with distinctlow L*, a*, and b* values occurring between 4.8 and 4.9 m, partiallycorresponding to the visible dark interval, and between 0.9 and0.8 m, matching the entire visible top dark layer. These two inter-vals in core 9501 seem to coincide with sapropels S3 and S1, asevident also in other parameters. In core 9509, all colour indexparameters are rather uniform except a high peak in L* values at2.8–2.6 m and lower values of a* from 1.6 to 1.2 m; the latterinterval coincides with sapropel S1 (Fig. 2B, a–c).

TOC values in core 9501 are low w0.2–0.3 wt% (Fig. 2A, d)whereas core 9509 is relatively more enriched in TOC at 0.6 to0.9 wt% (Fig. 2B, d). Sapropel layers are identified from the increasein TOC values in both cores. TOC values in sapropel layers in core9501 reach 1.2–1.3 wt% between 4.9 and 4.8 m and between 0.9 and0.8 m. The high TOC values in these layers match the colour indexpeaks. The highest TOC value in core 9509 of w1.7 wt% occursbetween 17.7 and 17.4 m and between 1.6 and 1.2 m, with a sharpdrop to w1.04 wt% at 1.35 m followed by a further increase to1.85 wt% at 1.3 m (Fig. 2B, d).

The average value of >63 mm size fraction in core 9501 is4� 2.3 wt% (Fig. 2A, e). It can be divided into two parts based on thesandy fraction content. The lower part of the core between 5 and2.2 m is characterized by low and uniform contents (w3 wt%). From2.2 m to the top of the core, the sand content increases with largefluctuations between 4 and 18 wt%. The 18% peak value coincideswith highest TOC content and lower colour index parameters in thiscore. In core 9509, the average sand content is low throughout at2.5�1.1 wt% (Fig. 2B, e). A slight increase in the sand fraction occursin the sapropel layers between 1.2 and 1.6 m coinciding with thehighest TOC values.

3.2. Isotopic composition of the planktonic foraminifera G. ruber

The stable isotope values of cores 9501 and 9509, are plotted inFig. 3 as a function of depth. The G. ruber d18O values vary between�0.6& and 3.8& in core 9501 (Fig. 3a) and from�1.65& to 3.1& incore 9509 (Fig. 3c). The higher values are associated with glacialintervals and the lowest values with sapropel events.

The d13C-depth profiles of G. ruber from cores 9501 and 9509 areplotted in Fig. 3b,d respectively. d13C values in core 9501 varybetween 0.5 and 3.0&, with the top 1.2 m being very variable withan amplitude of w2.0&. Between 1.2 and 2.8 m, the amplitude islower, with an average d13C value of w1.5&. The largest increase in

5

4

3

2

1

0

)m

(

ht

pe

D

30 40 50 60

L*

0 2 4 6

a*

5 10 15 20 25

b*

0 0.5 1 1.5 2

TOC (%wt.)

0 5 10 15 20

% >63µm

size fraction

B

a b c d e

18

16

14

12

10

8

6

4

2

0

)m

(

ht

pe

D

A

a b c d e

Fig. 2. Colour index showing parameters L*, a*, b*, TOC (% wt) and % of size fraction >63 mm in 9501 (A) and 9509 (B) vs. depth. The upper horizontal gray bar indicates sapropel S1and the lower bar sapropel S3.

A. Almogi-Labin et al. / Quaternary Science Reviews 28 (2009) 2882–28962886

d13C value occurs between 3.0 and 4.0 m, reaching w3.0& (Fig. 3b).In contrast, d13C values in core 9509 are lower and more homoge-neous with an average of w0.7&, but with frequent variationsranging between 0.5 and 1.5& (Fig. 3d).

3.3. Sea surface temperatures based on alkenones

Alkenone SST values are given in Table 2. They show significanttemperature variations over the lastglacial–interglacial cycle. SSTestimates for the Holocene and MIS 5.1 in core 9509 are generallyabove 18 �C. Minimum SST occurred during the LGM with temper-atures of 11–12 �C while temperatures varied between 13 and 16�Cduring the earlier part of the glacial. Because of low organic mattercontent, SST estimates for core 9501 were determined only for theHolocene where temperature values range between 17 and 21�C.

3.4. Time framework and sedimentation rates

As described in the methods, the time framework for both coresis largely based on the comparison with the adjacent well-datedSoreq Cave record (Bar-Matthews et al., 2003; updated in Vaks

et al., 2006) The ages at the base of the studied 5 m section in core9501 and the base of the 9509 core are 86.6 and 86 ka BP, respec-tively (Fig. 4a, b). This time framework of the marine cores aspresented in Fig. 4 is augmented by other criteria and observations:

1. The sapropels are shown to be time markers by most proxies(Fig. 2). The age of Holocene sapropel S1 in core 9501 wasdefined from 9.5 to 8.2 ka BP using the AMS 14C ages (Table 1),the change in sediment colour, colour index parameters,increase in TOC values and the increase of the percentage of the>63 mm size fraction (Fig. 2). The occurrence of the base of S1 incore 9501 also coincides with the sharp increase in the d13C ofthe speleothems (Fig. 4c), which is an indicator for the begin-ning of the deluge rain conditions on land that lasted until7.4 ka (cf. Bar-Matthews et al., 2000, 2003). The top of thesapropel in core 9501 based mainly on AMS 14C ages wasdetermined to 8.2 ka. This age coincides with a sharp inter-mittent decrease in d13C and increase in d18O of the speleothemrecord (Fig. 4a,c).

2. Core 9509, which was not dated by 14C, is correlated with theoxygen isotopic record of the speleothems in Fig. 4b and

5

4

3

2

1

0

De

pth

(m

)

4 3 2 1 0 -1 -2

18O (‰ VPDB)

0 1 2 3

13C (‰ VPDB)

18

16

14

12

10

8

6

4

2

0

4 3 2 1 0 -1 -2

18O (‰ VPDB)

0 1 2 3

13C (‰ VPDB)

9501

a b c d

9509

S1

S3

S1

S3

Fig. 3. d18O and d13C values vs. depth of cores 9501 (a, b) and 9509 (c, d).

A. Almogi-Labin et al. / Quaternary Science Reviews 28 (2009) 2882–2896 2887

compared with the corresponding isotopic record of core 9501in Fig. 4d. The age of the base of sapropel S1 in this core wasalso dated to 9.5 ka BP using this matching of isotopic profiles.The TOC profile of sapropel S1 in core 9509 differs from that ofcore 9501 (Fig. 2d) by having a bi-modal peak which is absentfrom core 9501. The low TOC ‘valley’ in core 9509 coincideswith the w8 ka event clearly documented in sapropel S1 ofeastern Mediterranean cores (Casford et al., 2001). The absenceof a second TOC peak in sapropel S1 of core 9501 is supportedby the fact that the 14C age of the top of sapropel S1 is 8.2 ka(Table 1). Therefore we suggest that duration of sapropel S1 incore 9509 was longer than that in core 9501.

3. The base of sapropel S3 in both cores is dated as 86 ka BP usingthe contemporaneous decrease in d18OG. ruber, increase in TOCvalues, and decrease in d18O and d13C of the speleothems(Fig. 2A(d), B(d) and Fig. 4). The interval of high TOC in core9501 lasted from 86 to 85 ka, while in core 9509 this intervalwas from 86 to 84.5 ka. The gradual increase in d18O in core9509 from a value�1& at 86 ka, continued until about w80 ka,when d18O values stabilized in both cores at about 2&.

4. The speleothem d18O record clearly shows that the glacialperiod can be divided into two main modes. The first is from 80to 49 ka where the average d18O¼�3.0� 0.35&. Both marinecores show this same trend, but it is more prominently definedin core 9501 (Fig. 4a). We correlated the lowest d18O values ofthe marine cores during the lastglacial with the lowest d18Ovalues of the speleothems, starting at 55 ka BP (Fig. 4a,b). Thistime period is also equivalent to the lowest d18O values for MIS3.3 in the SPECMAP record (Imbrie et al., 1984; Martinson et al.,1987). The end of this lowest d18O event was dated in themarine record to 49 ka BP, taking into account the timing of thetermination of the event in the speleothem record, butassuming more constant sedimentation rates.

5. We matched the younger part (second mode) of the lastglacialof the marine record using 3 distinct events in the speleothemrecord: the highest d18O values at 49 and 19 ka and the lowestd18O value occurring at 35.5 ka (Fig. 4a,b). This correlation alsoagrees with the corrected 14C age of core 9501 at 2.0 m of 25 ka(Table 1).

6. During deglaciation, there is a very clear trend of a gradual d18Odecrease of the order of 5& occurring from the LGM to the base

of sapropel S1. Superimposed on this trend it is possible toidentify the Bølling/Allerod low d18O values and the highervalues of the Younger Dryas (Fig. 4a,b). The post-sapropelperiod in core 9501 was constrained by the use of the two 14Cages and assuming a constant sedimentation rate. In core 9509,we matched the marine record with the speleothem isotopicrecord since they show similar trends.

The results of our age model calibration show that the averagesedimentation rate in core 9509 is 20.7 cm/ka, compared to 5.75 cm/ka in core 9501. The sedimentation rates are not constant throughtime with the variations being most marked in core 9509 (Fig. 5). It isimportant to note that the sedimentation rates in this core are muchhigher between 43 and 57 ka, up to 34.5 cm/ka, whereas during thelast 35 ka and in sediments older than 74 ka the sedimentation ratesare below 20 cm/ka. In core 9501 the sedimentation rates are higherduring the interglacials, up to 11.9 cm/ka, compared to the glacialswhere sedimentation rates are w4.3 cm/ka.

4. Discussion

4.1. The relationship between the southeasternand northern Levantine Basin

4.1.1. Sediment characteristicsThe two cores, taken w380 km apart, represent sedimentation

under the influences of two different climate regimes. The southerncore 9509, taken off southern Israel, is strongly influenced by theRiver Nile input, which originates in the low-latitude monsoonalsystem (Rossignol-Strick, 1985) whereas the northern core 9501,taken SE off Cyprus, is mostly under the influence of the NorthAtlantic/Mediterranean climate system.

The greater influence of the Nile on the southern Levant Basinsedimentation is evident from the sedimentation rates, colourparameters, TOC and d13CG. ruber. Although both cores lie at similarwater depths of 880 and 980 m, the average sedimentation rate incore 9509 located under the Nile plume is 20.7 cm/ka, which isnearly 4 times higher than that of core 9501. This can be attributedto the higher sediment load carried seasonally during the floodingperiod of the Nile. Hamann et al. (2008) also found high sedi-mentation rates in another core located under the Nile plume,

Table 2Organic matter (OM) characterization and SST calculations based on alkenones incores 9501 and 9509.

Depth (cm) Age (ka) SST (�C) OM origin

Core 95090–2 0.0 19.5 Mixed10–12 0.6 Terrestrial100–102 6.0 18.4 Marine110–112 6.9 Marine120–122 7.4 Marine130–132 7.9 20.1 Marine135–136 8.1 23.5 Marine140–142 8.4 19.4 Marine150–152 8.9 18 Marine160–162 9.3 Terrestrial170–172 9.7 a

225–226 12.1 15.2 Mixed250–252 13.2 15.5 Mixed330–332 19.0 Marine335–336 19.3 11.3 Marine340–342 19.7 a

350–352 20.3 Marine360–362 21.0 a

655–656 34.9 12.7 Marine700–702 36.6 13.2 Marine725–726 38.6 15.2 Marine1150–1152 52.9 14.3 Marine1220–1222 54.8 15.6 Marine1230–1232 55.1 Marine1240–1242 55.4 Terrestrial1250–1252 55.7 16.5 Marine1500–1502 69.2 14.2 Mixed1520–1522 70.3 18.3 Marine1775–1776 85.8 19.2 Marine1780–1782 86.0 18.1 Marine

Core 95010–2 0.0 Mixed2–4 0.2 Mixed4–6 0.4 Mixed6–8 0.6 17.2 Marine8–10 1.0 Mixed10–12 1.1 Mixed80–81 8.2 Marine81 -83 8.4 Marine83–85 8.6 18.2 Marine85–87 8.9 Marine87–89 9.2 19.3 Marine89–90 9.5 Marine90–92 9.9 a

92–94 10.4 a

98–100 11.7 20.6 Marine100–102 12.0 21.5 Marine102–104 12.4 Marine202–204 25.5 14.4 Marine

a Sediments with low content of organic matter.

A. Almogi-Labin et al. / Quaternary Science Reviews 28 (2009) 2882–28962888

based on grain-size analysis. Higher sedimentation rates of85–130 cm/ka were found by Schilman et al. (2001b) in latestHolocene cores located under the Nile plume, but in shallowerwater depths of 470 and 670 m. The high input of Nile material intothe southern Levantine Basin is also evident from the colour andthe TOC content. Core 9509 is continuously dark, both visually andas quantified by the three different colour parameters (Fig. 6), to theextent that sapropels could not be visually distinguished. Despitethe fact that in both cores the fine fraction comprise the majority ofthe sediments, the TOC wt% content is higher by 2–3 timescompared with the northern core. The dark colour and the high TOCdemonstrate that the River Nile was an important source ofnutrients that could enhance the local primary productivity,culminating with the accumulation of organic matter mainly ofmarine origin (Table 2). The darker colour of the sediments couldalso result from better preservation of the organic matter due to

higher sedimentation rates and/or fluvial contribution carryingoxides-rich sediments. The continuous low content of the sandyfraction in both cores indicates that the core sequences were notdisturbed by any turbidites. The increase of the sandy fraction inthe sapropels is due to an artifact caused by coagulation of organicmatter with clays as evident from the difficulties in separating thesandy from the fine fraction in the sapropel layers.

The influence of the River Nile dissipates rapidly to the northand can be hardly recognized in core 9501 southeast off Cyprus.This is consistent with other isotopic records from the easternMediterranean, such as 87Sr/86Sr and 143Nd/144Nd, which alsoindicate a limited River Nile contribution further north (e.g.,Warning and Brumsack, 2000; Weldeab et al., 2002; Box et al.,2008). However, during the periods when sapropels accumulated,the difference in sediment characteristics as well as 87Sr/86Sr (Boxet al., 2008) between the southern and the northern part of theLevantine Basin diminishes because of the increasing hydrologicalactivity in the Levantine Basin and the increasing flow of the RiverNile due to intensification of the monsoon system (Kallel et al.,1997; Emeis et al., 1998; Bar-Matthews et al., 2000; Calvert andFontugne, 2001; Almogi-Labin et al., 2004; Ehrmann et al., 2007).

4.1.2. Sea surface temperaturesThe Levantine Basin SST w11 �C for the LGM, based on the

alkenone measurements performed in this study and by Emeis et al.(1998, 2000, 2003) were lower by w10 �C compared to the earlyHolocene SST values of˜22 �C (Fig. 7a). The nearby land tempera-tures, that were retrieved from fluid inclusion data and D47

(‘clumped isotopes’) values in speleothems (McGarry et al., 2004;Affek et al., 2008) are similar to alkenone SST estimates. The land-model temperatures from fluid inclusion data during the LGM arew2 �C colder than the marine record, but within the error. Duringthe glacial between 70 and 20 ka BP, the SST record thoughdiscontinuous, shows a general cooling trend toward the LGM from18 �C down to 12 �C. Essallami et al. (2007) measured alkenone SSTvalues for the last 30 ka in core MD84-632 from the southern partof the Levantine Basin. Their SST values fluctuate between 15 and20 �C in the period from 30 to 25 ka BP. For this time period we haveno temperature records. However, for the LGM their SST values areslightly higher (w2 �C) than the SST values determined in thisstudy and by Emeis et al. (1998, 2000, 2003). This trend of slightlyhigher temperatures continues until the early Holocene.

Alkenone SST estimates from the western Mediterranean(Martrat et al., 2004) generally show a similar trend to the Levan-tine Basin (Fig. 7b). Because of the low resolution of our samplingfor SST measurements we were not able to define any N–S gradientin past SST in the Levantine Basin. Presently there is a 1–2 �Cdifference between the cooler northern and the warmer southernpart of the Levantine Basin (Marullo et al., 1999). It is also impos-sible to estimate W-E gradients in the Mediterranean. However,comparison of the western (Martrat et al., 2004) and the easternMediterranean alkenone SST records (Essallami et al., 2007) showsthat the eastern Mediterranean was warmer by w2 �C during theHolocene and by up to 6 �C between 30 and 15 ka. Presently thewest-east Mediterranean SST difference is in the order of 4 �C. Thusit seems that the eastern Mediterranean was warmer than itswestern part but the SST differences vary.

SST records from the whole Mediterranean indicate that duringthe lastglacial SST were significantly lower, by up to 10 �C than inthe Holocene (Fig. 7). This large amplitude contrasts with the muchsmaller 3.5 �C glacial/interglacial difference in the eastern tropicalAtlantic (Schneider et al., 1995; Nurnberg et al., 2000), and isconsistent with a stronger landlocked effect in the MediterraneanSea and a reduction in the moderating effect of the Atlantic Oceancompared to the Holocene.

-2

-3

-4

-5

-6

-7

18O

(‰

VP

DB

)S

oreq

9509

Soreq

4

3

2

1

0

-1

18

O (‰

V

PD

B) - 9509

-4

-6

-8

-10

-12

13C

(‰

V

PD

B)S

oreq

Soreq

b

c

-2

-3

-4

-5

-6

-7

-8

18O

(‰

V

PD

B)S

ore

q 9501

Soreq

4

3

2

1

0

-1

-2

18

O (‰

V

PD

B) - 950

1a

0 10 20 30 40 50 60 70 80 90

Age (kyr)

4

3

2

1

0

-1

18O

(‰

V

PD

B)G

. ru

ber

9501

9509

d

S1 S3

Fig. 4. d18O values of Soreq Cave and core 9501 vs. age (a), Soreq and core 9509 (b) d13C values of Soreq Cave vs. age (c) and d18O of cores 9501 and 9509 vs. age superimposed oneach other (d). The vertical gray bars indicate sapropels S1 and S3. Note that the time span of S1 in 9501 is shorter, ending at 8.2 ka compared with 9509.

0 20 40 60 80

Age (kyr)

5

4

3

2

1

0

De

pth

(m

) - 9

50

1

18

16

14

12

10

8

6

4

2

0

Dep

th

(m

) - 9

509

9501

9509

MIS 1 2 3 4 5.1

Fig. 5. Age model showing depth (m) of cores 9501 (gray) and 9509 (black) vs. age. Thenumber at the top marks marine isotope stages (MIS) 1–5.1 (Martinson et al., 1987).

A. Almogi-Labin et al. / Quaternary Science Reviews 28 (2009) 2882–2896 2889

4.1.3. Stable isotope profilesWhereas the differences in sediment characteristics between

the two cores are clear, and indicative of the greater contribution ofRiver Nile sediment load to the southern core, the d18OG. ruber

relationships are more complex. Comparison of d18OG. ruber values ofthe two marine records reveals an offset that varies with time(Fig. 8). Present day d18OG. ruber values along the northern andsouthern cores are w0.5& and �0.4&, respectively. This modernDd18O difference of 0.9& was also observed by Kallel et al. (1997). Itcannot be explained by the present day difference of surfacetemperature or surface salinity between the two areas. It isequivalent to about 4 �C difference in SST or about 4.5& differenceof salinity considering the d18O water/salinity relationship for theEastern Mediterranean (Kallel et al., 1997). The present day Dd18Omost probably reflects both a 1–2 �C increase in SST and a slightd18Ow decrease from 1.6& in the northern part of the LevantineBasin to w1.4–1.5& off the Israeli coast (Marullo et al., 1999; Pierre,1999; Schilman et al., 2001a). Kallel et al. (1997) also observedhigher surface water d18O north of Cyprus during sapropel S1,similar to the difference found today between the southern andnorthern parts of the Levantine Basin and in contrast with the restof the Mediterranean Sea.

During the last 90 ka the d18O difference between 9501 and9509 (Dd18O9501–9509) can be divided into several periods. Theaverage Dd18O9501–9509 is 0.24� 0.58& (1s) (Fig. 8) but fluctuations

of this parameter range from �1.4 to 1.5& and show evidence ofperiodicity.

During sapropel S3, the Dd18O9501–9509 was small and closer tothe average value compared to sapropel S1 (Fig. 8). SST differenceswere also small (Emeis et al., 1998; Fig. 7) indicating that during this

90

80

70

60

50

40

30

20

10

0

Ag

e (k

yr)

30 40 50 60

L*

95019509

0 2 4 6

a*

5 10 15 20 25

b*

0 0.5 1 1.5 2

TOC (%wt.)

0 5 10 15 20

% >63 m

size fraction

a b c d e

Fig. 6. Comparison of colour index parameters L*, a*, b*, TOC (% wt) and % of size fraction >63 mm in cores 9501 (gray) and 9509 (black) vs. age. The upper horizontal gray barindicates sapropel S1 and the lower bar sapropel S3.

A. Almogi-Labin et al. / Quaternary Science Reviews 28 (2009) 2882–28962890

period the N–S gradient was minor. The lastglacial is characterizedby three different Dd18O9501–9509 modes, which can serve as anindicator for the intensity of the Nile discharge and other variablesof the climate system and their influence on the EM. The first mode

Fig. 7. SST determined by alkenones in cores 9501 and 9509 compared with oth

is characterized by similar d18O values in both cores, (Dd18O9501–9509

is close to the average value), and it occurred from 43 to w32 ka,and during the end of the lastglacial between 23 and 16 ka (Fig. 8);the second mode is characterized by Dd18O9501–9509 values that are

er SST records from the eastern (a) and western (b) Mediterranean vs. age.

S1 S3

0 10 20 30 40 50 60 70 80 90

Age (ka)

-1

0

1

2

18O

(‰

V

PD

B)G

. ru

ber

400

420

440

460

480

In

so

latio

n (65

N Ju

ly

)

18O 9501- 9509 3-pt. running averageInsolation (65 N July)

18O (‰ VPDB) 9501 - 9509= 0.24±0.58

Fig. 8. Dd18O9501–9509 values, their 3-point running average and insolation at 65�N plotted vs. age. The vertical gray bars define sapropels S1 and S3. Minimum Dd18O periods

associated with maximum insolation are indicated by the shaded area.

A. Almogi-Labin et al. / Quaternary Science Reviews 28 (2009) 2882–2896 2891

significantly higher than the average, and occurred at several shortintervals at w67 ka, 48–43 ka, w32 to w24 ka. The third mode,characterized by negative Dd18O9501–9509 values, occurred at thebeginning of the lastglacial at 77 to 68 ka, 58 to 49 ka and duringthe deglaciation from 16 to 12 ka.

The period between 58 and 49 ka, representing the third mode(Figs. 4 and 8), is characterized by low values of Dd18O9501–9509

(as low as�1.2&) and low d18OG. ruber values (w1 to 0&), in the twocores. During this period the sedimentation rate increased signifi-cantly in core 9509 (Fig. 5). SST during this interval varied between13 and 17 �C (Emeis et al., 2003; and this study), the highest duringthe entire lastglacial. These low d18O values coincide with Dans-gaard-Oeschger interstadial 14, and with maximal insolation valuesin 65�N (Figs. 8 and 9). Another period with low Dd18O9501–9509 thatoccurs between 16 and 12 ka coincides with maximal insolationvalues at 65�N. Thus, we suggest that the lower d18O in the northernpart of the Levant is a response to high latitude warming relative tothe southern basin.

The unique period at w30 and w25 ka BP with significantlyhigher Dd18O9501–9509 values (sometimes exceeding 1&; Fig. 8)occurred at a time when the SST in the SE basin were extremelyhigh, above 17 �C (Essallami et al., 2007). There is a lack of highresolution SST data for this period from the northern part of theLevant (Fig. 7). However, the few existing SST values measured offCrete (Emeis et al., 2003) show that temperatures were as low asw13 �C, similar to the Alboran Sea SST (Martrat et al., 2004;Essallami et al., 2007). If these colder temperatures indeed onlyprevailed in the northern Levant, they would result in the largeDd18O9501–9509 observed between the two parts of this basin duringthis period.

During the time intervals typical of the first mode whenDd18O9501–9509 is small and close to the average value it is not clearif the climate in the north and south was similar, or if the contri-bution of the Nile was counterbalanced by lower temperatures inthe north. Due to the limited SST record, it is difficult to evaluate thereason for the near zero Dd18O9501–9509 values (Fig. 8). However, thecontinuous contribution of the River Nile to the SE Levantine Basinthroughout the lastglacial until w17 ka, is evident from the higherTOC contents in the southern core, the higher sedimentation ratesof this core, and its continuous dark colour (Fig. 6). The dischargefrom the White Nile was apparently negligible during the lastglacialperiod (Talbot et al., 2000), therefore we suggest that maincontributors of freshwater to the Mediterranean Sea were the BlueNile and the Atbara throughout the Glacial and only during most ofthe deglaciation there was a major decrease in their discharge.

The Younger Dryas period is characterized by negativeDd18O9501–9509 values similar to the time period between 58 and49 ka. The lower values in the northern part of the Levantine Basinreflect the warmer SST (4–5 �C), in the northern basin (Fig. 7). Thiswarming coincides with the northern hemisphere increase ininsolation (Fig. 9) and it emphasizes that the immediate effects ofwarming were felt earlier in the northern part of the LevantineBasin than in the south.

During most of the Holocene, including the early Holocene sap-ropel S1 period, Dd18O9501–9509 values are significantly above theaverage value with a difference of up to 1.3& (Fig. 8). Similarly,Fontugne et al. (1994) showed that during sapropel S1 a difference ofup to 0.9& is found between d18OG. ruber sampled off the Nile deltaand that sampled close to Cyprus. These authors attributed thedifference to the isotopic composition of the sea water (d18Ow)(assuming constant SST of 25�C) and inferred a significant contri-bution of lower d18Ow water, probably from the River Nile, into thevicinity of the Nile plume area. Alkenone SST was measured in thesouthern part of the basin by Essallami et al. (2007), and limited dataexist for the north (Emeis et al., 2000). Comparison between theserecords (Fig. 7a) shows that SST in the north was 1–2 �C coolerduring sapropel S1, indicating that the large Dd18O9501–9509 value canonly partially be attributed to the SST gradient, and that the RiverNile must have contributed to differences in d18Ow by making it moredepleted in the SE Mediterranean. Thus, based on the sedimentcharacteristics and Dd18O9501–9509 values of this work and the data ofFontugne et al. (1994), the SE Levantine Basin during sapropel S1 waslargely influenced by water derived from monsoonal activity.

The Dd18O9501–9509 values remain relatively high in the later partof the Holocene. Based on trace element studies, Calvert and Fon-tugne (2001) inferred a significant decrease in the Nile input duringthis period. Despite the decrease in River Nile input, its influence onthe southeastern Mediterranean was higher than in the NE Medi-terranean, as evident from the Dd18O9509–9501 together with thesedimentological data (Fig. 6)

4.2. Sea–Land d18O relationships

In a comparison between the Soreq d18O speleothem record andthe low resolution EM record (Vergnaud-Grazzini et al., 1977; Cal-vert and Fontugne, 2001), Bar-Matthews et al. (2003) showed thatEastern Mediterranean speleothems record have both regional andglobal paleoclimate implications. In general, the d18O values of thecave water from which speleothems form reflect the d18O of the seasurface source of atmospheric vapor that leads to cloud formation,

4

2

0

18O

(‰

V

PD

B)G

. ru

be

r

-44

-40

-36

18O

N

GR

IP

(‰

)

400

420

440

460

480

In

so

latio

n (6

5N

J

uly

)

9501

9509

a

c

0 10 20 30 40 50 60 70 80 90

Age (kyr)

4

2

0

18

O (‰

V

PD

B)

9501

ODP-977A

d

b

Fig. 9. Comparison between insolation, derived from Berger and Loutre (1991) (a), NGRIP (2004) record (b), d18OG. ruber of cores 9501 and 9509 (c) and d18OG. bulloides from thewestern Mediterranean (ODP-977A) after Martrat et al. (2004) with d18OG. ruber of core 9501 (d). The dashed lines connect between high d18OG. ruber values and low d18O NGRIPvalues. The solid lines connect between low d18OG. ruber and high d18O NGRIP values. The vertical gray bar indicates the time correlation between D-O interstadial 14, insolation andminimum d18OG. ruber in the Levantine Basin.

A. Almogi-Labin et al. / Quaternary Science Reviews 28 (2009) 2882–28962892

and the Dd18O between the marine sea surface source and rainfallthat results from evaporation processes above the sea, distancetraveled by the clouds over land, cooling effects due to increasedelevation, and an affect reflecting the amount of rainfall (‘amounteffect’). Based on a comparison of the speleothems record to thelow resolution marine record Bar-Matthews et al. (2003) showedthat the average Dd18Osea–land values were relatively constantduring the last 185 ka. Kolodny et al. (2005) interpreted this relativeconstancy across the lastglacial –present interglacial transition toindicate that the EM speleothems d18O record primarily reflects seasurface d18O changes. Bar-Matthews et al. (2003), however, inter-preted d18O oscillations during interglacial periods to largely reflectthe amount effect of decreasing d18O of rainfall with increasingamount.

We explore these questions using the higher resolution Dd18O-sea–land records obtained in this study from the two marine cores(Fig. 10). This comparison shows quite large systematic deviationsfrom the average Dd18Osea–land value of 5.6� 0.7& found by Bar-Matthews et al. (2003). The average Dd18Osea–land value for core9501 is 5.68& and for core 9509 it is 5.46&, showing that generallythe northern core value is closer to the earlier estimate value, whichwas based on the comparison of the Soreq speleothem record withthe d18OG. ruber record of Fontugne and Calvert (1992).

This could suggest that the major fingerprint of the vapor sourceis derived from the central/northern Levantine Basin. However,detailed examination of Dd18Osea–land values shows that smallerDd18O differences, of the order of w4–5& generally occurredduring interglacials MIS 5.1 and MIS 1 as well as between w60 and45 ka (i.e., MIS 3). In contrast, during most of the glacial period, andespecially during the transition from MIS 5.1 to 4 and during theLGM and early deglaciation, the Dd18Osea–land values are higher thanaverage, reaching a maximum value of almost 7&. The higherDd18Osea–land values during the glacial period potentially can beexplained by the increased sea–land distance and elevationincrease due to the lowered sea level relative to other periods. Boththese changes will lead to greater isotopic depletion in rainfallreaching the cave, through Rayleigh effects, and a consequentw1.5& deviation from the long-term average. On the other hand,periods with Dd18Osea–land values lower than the average, mostprobably reflect heavy rainfall events resulting in smaller sea–landfractionation (the amount effect). These coincide with sapropels S1and S3 events, periods that were suggested to have higher rainfallall over the Mediterranean (Kallel et al., 1997; Bar-Matthews et al.,2000). The amount effect during these events is amplified due tothe fact that the sea level was still lower than at present (Wael-broeck et al., 2002) and the land was larger. This suggests that the

A. Almogi-Labin et al. / Quaternary Science Reviews 28 (2009) 2882–2896 2893

smaller Dd18Osea–land between w60 and 45 ka also representsa more humid phase in the northern and southern Levantine Basin.Rossignol-Strick (1985) proposed that this period is associated withmaximum northern hemisphere insolation with negative preces-sion values. Bar-Matthews et al. (2000) showed also, based on theisotopic record of the speleothems, that this indeed was the wettestperiod during the lastglacial.

The transition from the LGM to the Holocene shows an overalldecrease from higher, to modern-day average values of Dd18Osea–

land in both cores. This most likely represents warming and theincoming of the amount effect as a major control of rainfall d18Oduring the Holocene and today.

The Dd18Osea–land values of the two marine cores (Fig. 10) showsimilar long-term oscillations despite the fact that the climate ineach of the sub-basins was not always the same. The similar long-term trends of Dd18Osea–land clearly indicate that the sources ofrainfall in the eastern Mediterranean landmass comprise a mixedregional signal from the northern and southern Levantine Basin.

The Dd18Osea–land variations demonstrate that various factorshave influenced the sea–land relationship during the last 90 ka. Theamount effect seems to be an important control of rainfall d18Oduring interglacial periods, and especially sapropel periods,whereas during glacial periods land distances and elevationdifferences arising from decrease in sea level may have broughtabout decrease in d18O of rainfall due to Rayleigh distillationprocesses. It is also important to add that changes in d18O of seasurface water will also have affected the d18O values of rainfall, andhence that of cave waters. This source effect is not directly recordedin Dd18Osea–land values, which would remained constant if thesource effect were the only control of rainwater d18O. The influ-ences of rainfall, the amount, sea–land distance and elevationchanges are thus superimposed on sea surface water d18O changesbrought about by continental ice melting, and the strong effectsexpressed in the southern Levantine Basin of the high River Nileinput during periods of enhanced monsoonal activity.

4

5

6

7

18

O (‰

V

PD

B) 9501-S

oreq

0 10 20 30 40

Age

18 O 9501-Soreq 3-pt. running average

7

6

5

4

18O

(‰

V

PD

B) 9509-S

oreq 18 O 9509-Soreq

3-pt. running average

a

b

18 O (‰ VP

18O (‰ VPS1

Fig. 10. Comparison between d18OG. ruber and d18OSoreq Cave record expressed as Dd18O9501

average (b). The vertical bars represent sapropels S1 and S3.

4.3. Comparison with western Mediterranean and global records

The study has demonstrated that the d18OG. ruber values of the SEand the NE Levantine Basin were generally similar, although thesouthern basin was located under a much stronger influence of theRiver Nile. However, there is a significantly larger differencebetween these eastern Mediterranean d18OG. ruber profiles and thoseof the western Mediterranean Sea (Martrat et al., 2004). Presently,there is a 1.0& difference in d18Ow between the western and theeastern Mediterranean Sea due to the strong W-E evaporationgradient (Pierre, 1999). As a result, the d18OG. ruber values from theEM Sea and d18OG. bulloides values from the western MediterraneanSea are similar in the top sediments, despite the higher SST in theeastern Mediterranean (Pinardi and Masetti, 2000).

During most of the last 90 ka, the d18OG. bulloides isotopic recordfrom the western Mediterranean is higher than that of d18OG. ruber

from the EM (Fig. 9d). This mainly reflects the lower SST in thewestern Mediterranean (Martrat et al., 2004; Fig. 7b). Between 80and 70 ka, the isotopic values all over the Mediterranean merge andbecame similar, suggesting that a negligible W-E gradient existed atthis time, probably due to similar evaporation and temperatureconditions over the entire basin (Kallel et al., 2000). The Dd18Oglacial-

interglacial difference (of d18OG. ruber), between 4.3 and 4.5& shown byour data and that of Vergnaud-Grazzini et al. (1977), Rossignol-Strickand Paterne (1999), Calvert and Fontugne (2001), and Essallami et al.(2007) is larger in the EM compared with the western Mediterra-nean where the Dd18Oglacial-interglacial is w3& (Cacho et al., 1999;Martrat et al., 2004). This difference reflects the increasing conti-nental affect on the eastern Mediterranean d18Ow, and the moder-ating influence of the Atlantic Ocean on the western Mediterranean.The moderating affect in the western Mediterranean also results inlower amplitudes in d18O during the glacial and the Holocene (Fig. 9).

A comparison of the two studied cores with the NGRIP (2004)record (Fig. 9) shows that a close correspondence exists, eventhough the time scale of the eastern Mediterranean record was

50 60 70 80 90

(kyr)

DB) 9501 - Soreq = 5.68±0.66

DB) 9509 - Soreq = 5.46±0.64 S3

-Soreq and its 3-point running average (a) and Dd18O9509-Soreq and its 3-point running

A. Almogi-Labin et al. / Quaternary Science Reviews 28 (2009) 2882–28962894

determined independently from that of the SPECMAP and theNGRIP records. An interval of constant high d18O values occursbetween 78 and 68 ka in the marine record with values around 2&

(Figs. 4 and 9c), whereas NGRIP (Fig. 9b) indicates that this period ofrather constant d18O values occurs later at between 70 and 62 ka.Evidently, the high amplitude d18O variations observed in NGRIPbetween w78 and w70 are more moderate in the eastern Medi-terranean land and sea records, either because of a reduced northhemisphere influence on the climate change in the eastern Medi-terranean and/or because of an offset in the dating of NGRIP. Suchan offset between these two records could be due to the differentdating method of the eastern Mediterranean marine records, whichstrongly relies on the highly accurate U/Th dating of the Soreq Cave.From 55 ka onwards there is a good agreement between therecords, as is evident by the general shape and especially during thecolder stadial events as shown in Fig. 9.

The period between 58 and 49 ka is considered to be globallyand regionally relatively warm (Figs. 7 and 9b), especially in thenorthern part of the EM basin. It also appears that this period wasthe warmest during the lastglacial in the western Mediterranean(Martrat et al., 2004), and coincides with Dansgaard-Oeschgerinterstadial 14, with maximal insolation values in 65�N.

5. Conclusions

This study has focused on the last 90 ka d18O and alkenone SSTrecords of two marine cores situated in the southern and northernparts of the Levantine Basin and the relationship of these records toeach other and to the on-land d18O record of the Soreq Cave. Theconnection between d18O compositions and sediment characteris-tics in two marine cores were used for identifying the contributionof the monsoonal system, via the River Nile and the northernAtlantic/Mediterranean on the paleoclimate of the Levantine Basin.The age model applied to the marine cores by their correlation withthe well-dated speleothems of the Soreq Cave is independentlyconfirmed through the correlation with NGRIP record, especially forthe lastglacial period.

Sediment characteristics indicate a distinct N–S differentiationin the Eastern Mediterranean. The fingerprint of the River Nile,regulated by the fluctuations of the monsoonal system, is mainlyrestricted to the southern Levantine Basin, and diminishes towardsCyprus. This is evident from the higher sedimentation rates, higherTOC and index colour parameters. During sapropel periods, thesedimentary characteristics of eastern Levantine Basin becamemore uniform owing to higher discharge of the River Nile reachingfurther north.

The d18O difference between the southern and the northern partof the Levantine Basin occurred in three different modes reflectingthe response of each sub-basin to local and global forcing. Duringsapropel S1 the distinctly higher Dd18O values result from signifi-cantly lower d18Ow values in the southern part of the basin, due tothe contribution of freshwater from the River Nile. Higher Dd18Ovalues at w25 and w30 ka BP were related to lower SST in thenorthern part of the basin. Lower Dd18O values between 58 and49 ka reflect a warming event in the north that coincided with theinsolation maximum in the northern hemisphere and D-O inter-stadial 14. The impact of this event became reduced southwards.Low Dd18O also occur from w15 to 12 ka BP coinciding withmaximum northern hemisphere insolation.

Sapropel S1 terminated earlier in the northern part of theLevantine Basin compared to the southern part of the basin. Thereasons for this are not clear, but may be related to earlier venti-lation in the northern Levantine Basin.

Sea–land Dd18O is similar in the two parts of the basin, varyingbetween 4 and 7&. Periods with smaller Dd18O occur during

warmer and more humid periods of the sapropels and between 58and 49 kyr. The hydrology during these periods is strongly influ-enced by the amount effect. In contrast, higher Dd18O values arerestricted to the end of the lastglacial when colder period prevailedand sea level was low. Here the data suggests that stronger Rayleighdistillation effects arising from greater land distance and higherelevation brought about lower d18O values in rain and cave waters.The influences of the amount effect, sea–land distance, and eleva-tion changes on the isotopic composition of rainfall are thussuperimposed on sea surface water d18O changes. The similartrends of the Dd18Osea–land exhibited by both cores suggest that theorigin of rain over the land is composed of mixed signal from thesouthern and northern Levantine Basin.

A comparison of SST values for the eastern and western Medi-terranean Sea shows that there was an apparently similar temper-ature rise from the LGM to the Holocene of about 10 �C. This largetemperature change reflects the fact that the Mediterranean Sea isisolated from the moderating effects of the Atlantic Ocean. Never-theless the Eastern Mediterranean Sea and land area is considerablywarmer than the western Mediterranean throughout the LGM –Holocene transition, and the d18OG. ruber drop of 4.5& is significantlygreater than the 3& shift found for the western Mediterraneand18OG. bulloides. It is thus clear that there is an increased continentaleffect from the western to eastern Mediterranean.

Acknowledgments

We thank the master, crew and scientists aboard R/V MarionDufresne for their help and assistance. We gratefully acknowledgethe technical assistance received from M. Kitin, Geological Survey ofIsrael. Thanks are due to the AMS 14C dating lab in Australia. Theresearch was supported by grants no. 20/01-13.0 and 910/05 fromthe Israel Science Foundation. N. Kallel and anonymous reviewerare thanked for suggestions resulting in a significantly improvedmanuscript and N. Roberts for handling the manuscript.

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