Palaeogeography, Palaeoclimatology, Palaeoecology 311 (2011) 230–241

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Palaeogeography, Palaeoclimatology, Palaeoecology

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Variations of the Nile suspended discharges during the last 1.75 Myr

Yulong Zhao a,b,⁎, Zhifei Liu a, Christophe Colin b, Martine Paterne c, Giuseppe Siani b, Xinrong Cheng a,Stéphanie Duchamp-Alphonse b, Xin Xie a

a State Key Laboratory of Marine Geology, Tongji University, Shanghai 200092, Chinab Laboratoire des Interactions et Dynamique des Environnements de Surface (IDES), UMR 8148 CNRS-Université de Paris-Sud 11, Orsay 91405, Francec Laboratoire des Sciences du Climat et de l'Environnement (LSCE), UMR 8212 CEA-CNRS-UVSQ , Gif-sur-Yvette 91198, France

⁎ Corresponding author at: State Key Laboratory of MaShanghai 200092, China.

E-mail address: yeoloon@gmail.com (Y. Zhao).

0031-0182/$ – see front matter © 2011 Elsevier B.V. Alldoi:10.1016/j.palaeo.2011.09.001

a b s t r a c t

a r t i c l e i n f o

Article history:Received 12 March 2011Received in revised form 27 July 2011Accepted 4 September 2011Available online 10 September 2011

Keywords:Eastern Mediterranean SeaThe Nile RiverSuspended dischargeAfrican monsoonObliquity

Planktonic foraminiferal oxygen isotopes, carbonate and organic carbon contents, as well as XRF core scan-ning Ti, V and Ba intensities have been analyzed on sediments of a Mediterranean marine core (CoreMD90-964) to reconstruct changes in the Nile discharges during the last 1.75 Ma. Chronology of CoreMD90-964 is established by correlating the planktonic foraminiferal δ18O record to the Mediterranean Globi-gerinoides ruber δ18O stack. A total number of 42 dark layers are observed under visual observation, amongwhich 21 of them are identified as sapropels by organic carbon contents (N1 wt.%). It is observed that oxida-tion of sapropels is very active during 1000 and 240 kyr, indicating that the local climate is much colder anddryer during this interval. Our results provide new evidences that oxidation of organic material play a majorrole in determining the organic carbon contents in sapropels. Carbonate contents and normalized titaniumand vanadium contents in bulk sediments are suggested to reflect mostly changes in the Nile discharge. Itis found that variation of the African monsoon intensities can strongly affect changes in the Nile suspendeddischarge via both runoff and drainage precipitation. Elevated Nile discharges are generally observed duringAfrican monsoon maxima (deposition of sapropelic layers). Our results suggest that oscillations in the Nilesuspended discharges are more the result of river transport capability than that of erosion potential in sourceareas. It is also observed that cyclicity of the variation of the Nile suspended discharge is generally paced by a78-kyr cycle, the physical meaning of which remains hitherto unclear. It is probably a bundling of 2–3 obliq-uity cycles, as a result of nonlinear responses of Atlantic SST to orbital forcing.

rine Geology, Tongji University,

rights reserved.

© 2011 Elsevier B.V. All rights reserved.

1. Introduction

Known for the organic-rich dark layers deposited between normaldeep-sea marls, called sapropels, the Mediterranean Sea has broadlyappealed interests of paleoclimatologists from a wide range of disci-plines. Since the pioneering work of Rossignol-Strick et al. (1982), acommon consensus is generally reached that formation of sapropels inthe Mediterranean Sea is tightly related to augmentation of the Nilerunoff during maxima of African monsoon precipitation. Annually,huge amounts of fluvial sediments from basaltic regions in the north-eastern Africa are transported into the eastern Mediterranean Basinvia the Nile River (Williams et al., 2006). These sediments have greatimpact, or even decisive effect, on sedimentary and ecological systemin the eastern Mediterranean Sea (e.g. Rossignol-Strick, 1983, 1985;Calvert and Fontugne, 2001; Revel et al., 2010). The total amount, aswell as mineralogical and geochemical composition, of the Nile sedi-ments is thus essential for us to understand climate variability of thenortheastern Africa and long-term changes of African monsoon

(e.g. Foucault and Stanley, 1989; Wehausen and Brumsack, 1999,2000; Foucault and Mélières, 2000; Calvert and Fontugne, 2001; Revelet al., 2010). However, owing to deficiency of a proper proxy, thereare hitherto great disputes on how discharge of the Nile varies in theEarth's history. As suggested byMilliman and Syvitski (1992), sedimentdischarge of a river is primarily controlled by the gross area and topog-raphy of the river catchment. In contrast, average net precipitation andrunoff generally affect sediment discharge of a river to a lesser extent(Milliman and Syvitski, 1992). It is reported that the network oflarge rivers is sometimes able to buffer the changes in riverdischarge and keep the total discharge at the outlets constant (Métivierand Gaudemer, 1999). This effect is, however, probably not very signif-icant for the Nile River because the Nile distributaries are mostly smalland seasonal streams (Milliman and Farnsworth, 2011) and morethan 90% of the Nile suspended discharge is derived from the EthiopianHighlands (Williams et al., 2006). Provided that topography and basinarea of a specific river do not change markedly in a relative stable tec-tonic context, sediment discharge of the Nile is thus dominantly affect-ed by variation of precipitation and river runoff. Both precipitation andriver runoff in the northeastern Africa are closely related with thechanges in African monsoon intensities (Ziegler et al., 2010). Duringmaxima of African monsoon, both the Nile runoff and precipitation in

231Y. Zhao et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 311 (2011) 230–241

the catchment are greatly enhanced. The responses of the Nile dis-charges changes to the African monsoon precipitation are, however,very complicated. On one hand, sediment load capability of the Nile isremarkably enhanced in virtue of the elevated runoff during Africanmonsoon maxima. On the other hand, vegetative cover in source areasis concurrently promoted in a wetter environment, protecting soils insource areas from being eroded (Adamson et al., 1980; Krom et al.,2002; Rohling et al., 2002; Weldeab et al., 2002). Changes of the Niledischarges are generally results of the counterbalance between thetwo opposite procedures. Evidences from Sr and Nd isotopic ratios(Krom et al., 2002;Weldeab et al., 2002) and claymineralogy (Foucaultand Mélières, 2000) have suggested that terrigenous input from theNile is generally reduced during sapropel formation. Studies ofheavy minerals, however, suggest higher discharges from the Nile dur-ing African monsoon maxima (Foucault and Stanley, 1989; Revel et al.,2010). Despite complexity of the govern factors, the deficiency of agree-ment is also attributable to the absence of robust proxies that reflectlong-term changes of the Nile discharge. The existing records that re-flect changes in the Nile sediments inputs are dedicated either to rela-tive short time intervals (Cita et al., 1977; Aksu et al., 1995; Calvertand Fontugne, 2001; Roussakis et al., 2004; Ehrmann et al., 2007;Hamann et al., 2009; Revel et al., 2010) or to the Pliocene epoch(Wehausen and Brumsack, 1999, 2000; Foucault and Mélières, 2000).In this study, we present for the first time a Nile discharge record ofthe last 1.75 Myr based on XRF core scanning vanadium (V) and titani-um (Ti) relative proportions of a deep-sea core in the eastern Mediter-ranean Sea (Core MD90-964). A detailed chronology is established onthe basis of planktonic foraminiferal oxygen isotopes to investigatethe deposition of sapropels and their potential links with variations ofthe Nile discharge. By virtue of their high resolution and long temporalcoverage, the geochemical (Ti and V) records allow us to discuss thelong-term changes of the Nile suspended discharge and its responseto the orbital forcing.

2. Materials and methods

Core MD90-964 (33°02.75′N, 32°38.57′E; water depth 1375 m)wascollected during PROMETE III cruise of the R/V Marion Dufresne in

31°30°29°28°

2500

Rosetta

NileDelta

LevantineBasin

Fig. 1. Sitemap of Core MD90-964. Bathymetry of the Levantine Basin is

September 1990. It is located on the distal fan of the Nile in the Levan-tine Basin (Fig. 1). Sediment of Core MD90-964 consists of pale creamto yellowish brown foraminiferal and nannofossil marl ooze, alternatedwith a series of organic-rich dark layers varying in thickness from 2 to41 cm. The core was sampled at a depth interval of 4 cm and a totalnumber of 797 samples were obtained on the upper 32.12 m of thecore (sediments under 32.12 m are disturbed). The samples wereused to analyze planktonic foraminiferal stable oxygen isotopes, bulkcarbonate contents (Carb%), and total organic carbon contents (Corg%).Vanadium (V), titanium (Ti) and barium (Ba) intensities were achievedon core surfaces using an X-ray fluorescence (XRF) Core Scanner.

Stable oxygen isotope analyses of planktonic foraminifers wereperformed on a Finnigan MAT-252 mass spectrometer at the StateKey Laboratory of Marine Geology, Tongji University. Measurementswere taken on well-preserved (clean and intact) planktonic foramin-ifer Globigerinoides ruber (white) specimens, 250–315 μm in diame-ters. The mean external reproducibility of carbonate standards was~0.07‰ for δ18O. Conversion to the international Pee Dee Belemnite(PDB) scale was performed using the NBS-19 standards.

Bulk sediment carbonate contents (Carb%) were determined usingthe gasometric method. Bulk sediments were dried prior to the mea-surement. About 100 mg sediments from each sample were thenweighed out to react with 8 N HCl to determine the carbonate con-tents. Errors of gasometric carbonate tests in this study were general-ly estimated to about 2%. Organic carbon contents (Corg%) wereobtained on an Elementar vario EL III Element Analyzer at the StateKey Laboratory of Pollution Control and Resources Reuse, Tongji Uni-versity. Carbonate fraction in bulk sediments were removed by react-ing with 0.1 N HCl prior to the measurements. The output percentagedata (in carbonate-free fraction) were calculated back to organic car-bon contents in bulk sediments.

XRF core scanning analysis was performed centimeter-by-centimeter on core surfaces using an Avaatech Core Scanner at theState Key Laboratory of Marine Geology, Tongji University. Surfaceof the split core sections was covered with a 4 μm thick Ultralene®foil to avoid contamination of the XRF measurement unit and desicca-tion of the sediment. The sediment surfaces were then carefullysmoothed to get a maximum quality, with all air bubbles under the

MD90-964

34°N

33°

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31°

34°E33°32°

25002500

200

500

1000

1000

1250

1500

1750

2000

2000

Damietta

ODP 967

ODP 966

represented by change of color (revised from Mascle et al., 2006).

232 Y. Zhao et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 311 (2011) 230–241

foil and all wrinkles in the foil aboratively eliminated. All core sec-tions were measured twice at different voltage to obtain intensitiesof V, Ti (10 kV) and Ba (50 kV), respectively. Intensities of each ele-ment were reported versus counts per second (cps), which was pro-portional to the chemical concentration. Total counts of elementsmeasured under 10 kV was adopted as a normalizing parameter to as-sess relative degrees of enrichment/depletion of vanadium and titani-um. The normalized vanadium and titanium intensities, hereafterreferred as V* and Ti* in this study, are employed to represent the va-nadium and titanium contents.

3. Results and discussions

3.1. Age model

Chronology of Core MD90-964 was developed by correlating theplanktonic foraminiferal δ18O record to the Mediterranean G. ruberδ18O stack compiled by Lourens et al. (2004, MedStack). The correlationwas performed semi-automatically using the Macintosh programAnalySeries (Paillard et al., 1996). To evaluate the agemodel, the recordswere Gaussian-filtered at both obliquity (0.02439±0.003 kyr−1) andprecession (0.04762±0.01 kyr−1) bands to assess coherences of thetwo curves. Phases of the two records match very well with each otherat both bands, despite discrepancy of amplitudes observed at sometime intervals (Fig. 2). Another simple but effective way to assess theage model is to perform cross-spectral analysis against a target curvethat reflects characteristics of the orbital cyclicity of the Earth (Hinnov,2004). The ETP (ETP=e+t−p) was defined herein as the target curvewhere e, t, and p represent normalized eccentricity, obliquity, and pre-cession from the Laskar et al. (2004) astronomical solution, respectively.Cross-spectral analyses of Core MD90-964 δ18O record against bothMedStack and ETP were performed. The spectra of Core MD90-964δ18O are highly coherent with both MedStack and ETP at the 41-kyrobliquity band and 23- and 19-kyr precession bands with N95% non-zero coherency (Fig. 3). Non-zero coherency of Core MD90-964 δ18Oagainst MedStack and ETP records at 100-kyr eccentricity band dropsto 80–95% and b80% levels, respectively (Fig. 3). Low coherency of

2 8 10 12 14 1618 20 22 26 2830 34 4240 48 52 56 58 64 66 74 80 82 8S1 S3S4S5 S6 S8 S102 8 10 12 14 1618 20 22 26 2830 34 4240 48 52 56 58 64 66 74 80 82 8

4

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B)

Gau

ssia

n fil

terin

g a

t 41

kyr

band

0 200 400 600 800

Age (ky

S1 S3S4S5 S6 S8 S10

Fig. 2. Age determination of Core MD90-964. Solid bar in gray show positions of sapropelsCodes of sapropels are showed at the top the figure, while the corresponding numbers belois calculated by an average of June–July–August from the Laskar et al. (2004) orbital solutio

MD90-964 δ18O against ETP records at the eccentricity band may beowing to low spectral densities of 100-kyr eccentricity components inthe ETP signal (Fig. 3b). It is showed that δ18O record of Core MD90-964 and the MedStack are generally in phase with each other at allMilankovitch orbital cycles (Fig. 3a). This is in agreementwith theGauss-ian filtering results. Minima of Core MD90-964 δ18O record lag the ETPby ~45° at the obliquity band and ~60° at the precession band (~61° atthe 23-kyr band and ~64° at the 19-kyr band), equivalent to 5.1 kyrand 3.6 kyr respectively (Fig. 3b). Combining together the Gaussian fil-tering and cross-spectral results, we can infer that the chronology ofCore MD90-964 is robust. According to the chronology, the bottom ageof CoreMD90-964 is estimated to be about 1747 ka. The linear sedimen-tation rate (LSR) is highly variable between 0.17 and 16.67 cm/kyr (aver-age 1.84 cm/kyr). The highest (16.67 cm/kyr) and lowest (0.17 cm/kyr)LSR values are observed during MIS 3 and MIS 10, respectively.

δ18O values vary between ~−3 and 3‰ PDB, with the heaviest andlightest values occur during the LGM (3.27‰ PDB) and the Eemianstage (−2.97‰ PDB), respectively. To have a better understanding ofits variation, the δ18O record of Core MD90-964 is overlapped withother G. ruber δ18O records in the eastern Mediterranean to inspecttheir similarities and differences (Fig. 4). It is showed that the oxygenisotope record of Core MD90-964 exhibits a similar δ18O shift at majorglacial/interglacial transitions as other records in the easternMediterra-nean. Relative to the Core MD84-641 (Fontugne and Calvert, 1992),KC01B (Rossignol-Strick et al., 1998) and Vrica/Crotone (Lourens etal., 1996b) records, Core MD90-964, which covers almost the wholeQuaternary, ismuch longer in time coverage. The oxygen isotope recordof ODP Site 967 (Kroon et al., 1998), albeit longer in coverage, is muchlower in resolution and teemed with hiatus. Consequently, our results(Core MD90-964) provide for the first time in the eastern Mediterra-nean Sea a continuous and high-resolution planktonic foraminiferal(G. ruber) oxygen isotope record of the Quaternary period.

3.2. Sapropel chronology

Totally 42 dark layers are recognized from Core MD90-964 undervisual observations. Details of these dark layers as well as their

8 90 9694 100 104 110112 120118 122126 130 134 140 142144148150152 156 160 164 170fnoqruvdS *

8 90 9694 100 104 110112 120118 122126 130 134 140 142144148150152 156 160 164 170

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at 65ºN (W

/m2)

Gaussian filteringat 23 kyr band

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r)

fnoqruvdS *

. Dashed and dotted lines show positions of ghost and hidden sapropels, respectively.w refer to their corresponding insolation peaks (i-cycles). The 65°N summer insolationn.

(a) (b)

Fig. 3. Cross-spectral results of MD90-964 δ18O against (a) MedStack and (b) the ETP records using the ARAND software. Horizontal dashed lines labeled 80% and 95% show the non-zero coherency between the two records.

233Y. Zhao et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 311 (2011) 230–241

counterpart sapropels in other eastern Mediterranean cores are listedin Table 1 to have a general view of sapropel chronology in the east-ern Mediterranean Sea. The term “sapropel” in paleoceanography is

2-4 6 8 10 12 14 16 18 20 22

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MD90-964 MD84-641 KC0

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uber

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er δ

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Age

Fig. 4. A compilation of G. ruber δ18O rec

generally defined by organic carbon contents in bulk marine sedi-ments, nevertheless criterions adopted by different authors are highlyvariable (Kidd et al., 1978; Hilgen, 1991; Fontugne and Calvert, 1992).

24 2628 30 3234 36 38 4042 44 46 48 50 52 54 5658 60

4

2

0

-2

-4

1B Vrica/Crotone ODP 967

G.ruber δ

18O (‰

PD

B)

1000 1200 16001400

(kyr)

ords in the eastern Mediterranean.

Table 1Compilation of sapropel chronology in the eastern Mediterranean.

ODP 967 a KC01B b Vrica c This study

Sapropels i-Cycle Age(ka)

Sapropels Si-Cycle Age(ka)

Sapropels i-Cycle Age(ka)

Top (m) Bottom(m)

Midpoints (m) Lithology i-Cycle Age(ka)

SAP1 2 8 S1? Si2 0.47 0.80 0.64 S1 2 8* 4 35a 6 55SAP2 8 81 S3? Si8 81 4.99 5.12 5.06 S3 8 81SAP3 10 102 S4 Si10 102 6.02 6.08 6.05 S4 10 103SAP5 12 124 S5 Si12 124 6.66 7.00 6.83 S5 12 124* 14 148SAP6 16 172 S6 Si16 172 7.72 7.78 7.75 S6 16 172SAP7 18 195 S7 Si18 195 7.89 7.92 7.91 DGL d 18 195SAP8 20 217 S8 Si20 217 8.03 8.12 8.08 S8 20 217* 24 262* 26 288 S′ si26 288 9.98 10.07 10.03 DGL 26 288SAP10 30 331 S10 Si30 331 11.36 11.46 11.41 S10 30 331* 34 369b 38 407 S11 Si38 407

12.63 12.71 12.67 DGL 40 42512.92 12.98 12.95 DGL 42 446

SAP11 44 461 S12 Si44 461SAP12 46 483SAP13 48 503 13.56 13.80 13.68 DGL 48 503* 50 529 Sa Si50 529* 52 553 14.30 14.32 14.31 GL e 52 553SAP14 54 575 S* si54 575SAP15 56 597 Sb Si56 597 15.14 15.33 15.24 GL 56 597c 58 618 * si58 618* 60 647SAP16 62 668SAP17 64 690 16.94 17.00 16.97 DGL 64 690

17.13 17.18 17.16 DGL 66 712* 74 785 * si74 785 17.68 17.84 17.76 DGL 74 785* 76 804* 78 822* 80 841* 82 862 * si82 862 18.21 18.43 18.32 DGL 82 862* 84 882* 86 908 * si86 908SAP18 88 934 19.24 19.35 19.30 GL 88 934SAP19 90 955 S" si90 955 19.54 19.62 19.58 DGL 90 955SAP20 92 977 Sc Si92 976SAP21 94 997 * si94 997 20.22 20.31 20.27 DGL 96? 1015SAP22 96 1027 * si96 1027SAP23 98 1048 * si98 1048SAP24 100 1070 Sd Si100 1070 20.76 20.92 20.84 Sd 100 1069SAP25 102 1091 * si102 1091SAP26 104 1111 * si104 1111 21.20 21.23 21.22 SAP-104 104 1111* 108 1144 * si108 1144SAP27 110 1164 21.80 21.90 21.85 SAP-110 110 1164SAP28 112 1185 21.92 21.97 21.95 SAP-112 112 1185* 114 1203* 116 1222* 118 1240* 120 1261 22.90 22.93 22.92 SAP-120 120 1261SAP29 122 1280 v 122 1280 23.14 23.38 23.26 v 122 1280SAP30 126 1315 u 126 1315 23.77 23.96 23.87 u 126 1315SAP31 128 1335SAP32 130 1356 t 130 1356 25.15 25.16 25.16 GL 130 1356SAP33 132 1376SAP34 134 1398 * 134 1398 25.68 25.86 25.77 * 134 1399SAP35 138 1429 s 138 1429SAP36 140 1449 r 140 1449 26.62 27.07 26.85 r, q 140-142SAP37 142 1471 q (C13) 142 1471* 144 1490 27.29 27.37 27.33 GL 144 1490* 148 1524 27.82 27.93 27.88 GL 148 1524* 150 1544 28.42 28.46 28.44 GL 150 1544SAP38 152 1564 p (C12) 152 1564 28.76 28.81 28.79 GL 152 1564SAP39 154 1584SAP40 156 1603 o (C11) 156 1603 29.34 29.40 29.37 o 156 1603SAP41 158 1622SAP42 160 1642 n (C10) 160 1642 30.07 30.13 30.10 n 160 1642* 162 1661

30.67 30.79 30.73 SAP-164 164 1681

234 Y. Zhao et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 311 (2011) 230–241

ODP 967 a KC01B b Vrica c This study

Sapropels i-Cycle Age(ka)

Sapropels Si-Cycle Age(ka)

Sapropels i-Cycle Age(ka)

Top (m) Bottom(m)

Midpoints (m) Lithology i-Cycle Age(ka)

SAP43 166 1694SAP44 168 1715 h (C9) 168 1715SAP45 170 1736 f (C8) 170 1736 31.84 31.89 31.87 f 170 1736

a Emeis et al., 2000.b Langereis et al., 1997.c Lourens et al., 1996b.d DGL refers to dark gray layers.e GL refers to gray layers.

Table 1 (continued)

235Y. Zhao et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 311 (2011) 230–241

In the present study, sediment layers with Corg% values higher than1 wt.% are defined as sapropels (Fontugne and Calvert, 1992). Accord-ingly, 22 of all dark layers are recognized as sapropels, whereas therest are simply described by their colors as “dark gray layers” (DGL)or “gray layers” (GL) (Table 1). Coding of sapropels is principallybased on the nomenclature in Lourens et al. (1996a). The code “Sd”,which is firstly introduced by Langereis et al. (1997), is also adoptedherein to describe the sapropel at insolation cycle (i-cycle) 100.Sapropels that beyond the nomenclature of Lourens et al. (1996a)and Langereis et al. (1997) are informally coded by the prefix “SAP”with their insolation cycles (e.g. “SAP-34” stands for the sapropel atinsolation cycle 34).

Dark layers with Corg% values b1 wt.% are visually observed at theequivalent positions of S7, t, p (Lourens et al., 1996a) and S′, Sb, S″(Langereis et al., 1997), whereas S9, S11, S12 (Lourens et al., 1996a)and Sa, Sc (Langereis et al., 1997) are absent from both visual observa-tion and Corg% profiles. It indicates that these sapropels have probablyundergone severe post-deposition oxidation, because organic materialsin sapropels can be largely reduced or even completely removed dur-ing post-depositional oxidation (Higgs et al., 1994; Thomson et al.,1995; van Santvoort et al., 1996, 1997; Emeis et al., 2000; Calvert and

2 8 1012 14161820 22 262830 34 4240 48 52 56 58 6466 74 8082

S1 S3S4S5 S6 S8 S10

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Age (

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δ18O

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(%

)

Fig. 5. Variation of carbonate contents, organic carbon c

Fontugne, 2001; van Santvoort et al., 2002). Therefore, it is necessaryto develop an alternate criterion for identification of sapropels in envi-ronments of severe post-depositional oxidation. Since enrichment ofBa is frequently observed in Mediterranean sapropels (e.g. Thomsonet al., 1995; Wehausen and Brumsack, 1999; Calvert and Fontugne,2001; Rinna et al., 2002; Arnaboldi and Meyers, 2003; Weldeab et al.,2003; Arnaboldi and Meyers, 2007), the Ba contents is introduced insome studies to identify the “missing” or “ghost” sapropels (Langereiset al., 1997; Emeis et al., 2000; Calvert and Fontugne, 2001). A possibleprocedure that can mislead the sapropel identification is the mobiliza-tion of barite (BaSO4) in marine sediments, which frequently happensunder sulfate-reducing conditions (Dymond et al., 1992). This phe-nomenon, however, is demonstrated insignificant in the Mediterra-nean settings, because Ba enrichment is commonly observed in deep-sea sapropels, regardless of the anoxia intensity (Passier et al., 1999;Wehausen and Brumsack, 1999, 2000). In Core MD90-964, Ba enrich-ment is observed in all visible dark layers (including sapropels, GLs,and DGLs in Table 1) except for the horizons at i-cycle 52, 66, and152 (Fig. 5). To distinguish from the “normal” sapropels, the GLs andDGLs are called “ghost sapropels” (Larrasoaña et al., 2006). Otherthan the dark layers, enrichment of Ba is also found at some intervals

88 90 9694 100 104 110 112 120118 122126 130 134 140142144148150152156160 164 170

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ontents, and barium intensities in Core MD90-964.

236 Y. Zhao et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 311 (2011) 230–241

(i-cycle 14, 22, 58, 80, 94, and 118) in which neither dark layers norCorg% peaks are observed. These layers, analytical detectable but notvisible, are equivalent to the “hidden sapropels” in Larrasoaña et al.(2006). The “sapropels” (with Corg%N1 wt.%), “ghost sapropels”, and“hidden sapropels” are generally considered to form under the samemechanism but have undergone different levels of post-deposition ox-idation (Larrasoaña et al., 2006). They all together constitute the“sapropel chronology” of Core MD90-964 (Fig. 5). These layers will bementioned indiscriminatively as “sapropelic layers” hereafter unlessspecifically stated. To summarize, sapropel chronology of Core MD90-964 consists of 48 sapropelic layers, among which 22 layers are recog-nized as sapropels (with Corg%N1 wt.%).

It is also observed that both the Corg% and the Ba intensitiesare extremely low between 1000 and 400 ka, indicating that post-depositional oxidation of organic material is remarkably extensive dur-ing this period (Fig. 5). Such kinds of periods, called red intervals, indi-cate that the deep-sea basin is flushed by water masses that carriedenough oxygen to erase the recently deposited sapropels at certaintimes (Emeis et al., 2000). Occurrence of the red interval is generally a

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0123456

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Age

Fig. 6. Variation of titanium and vanadium contents as

basin-wide phenomenon (Emeis et al., 2000), with the intensities ofwhich seems closely related with the water depth. All sapropelic layersin sediments of ODP Site 966 (Eratosthenes Seamount, water depth927 m) during 1000 and 240 ka are severely oxidized (Emeis et al.,2000). The red interval in CoreMD90-964 can actually also be extendedto 240 ka, if S10 and SAP-34 is envisaged as a short interruption (Fig. 5).A similar red interval is also observed in ODP Site 969 (MediterraneanRidge, water depth 2200 m), interrupted by SAP-46, SAP-54, and SAP-56 (Emeis et al., 2000). At the location of ODP Site 967 (LevantineBasin), which is much deeper in water depth (2553 m), the isochronalred interval is more frequently interrupted by sapropels, with totally 9sapropels survive (Emeis et al., 2000). The red interval in ODP Site 964(Ionian Sea, water depth 3658m), however, only develops during theinterval of 900–700 ka (Emeis et al., 2000). Combining together all re-sults from Core MD90-964 and the eastern Mediterranean ODP sites,we can infer that the intensities of organicmatter degradation in sapro-pelic layers are closely related with the water depth, with the deepersamples undergoes stronger degradation. This is in accordance withMurat and Got (2000), who suggest that Corg% in sapropels is linearly

6

4

2

0

0

20

40

60 Carb%

(%)

V* (×

10-3)

7

5

3

1

Ti* (×

10-2)

88 90 9694 100 104 110 112 120118 122126 130 134 140142144148150152156160 164 170

fnoqruvdS *

1000 1200 16001400

(kyr)

well as total terrigenous fluxes in Core MD90-964.

0

1

414 100 78 41 23 19S

pect

ra D

ensi

ty(a)

Ti*

0

Spe

ctra

Den

sity

1(b)

V*

Frequency (1/kyr)Period (kyr)

0.00 0.01100.0 50.0 33.3 25.0 20.0 16.7

0.02 0.03 0.04 0.05 0.06

0.00 0.01100.0 50.0 33.3 25.0 20.0 16.7

0.02 0.03 0.04 0.05 0.06

Fig. 7. Estimated power spectrum of the MD90-964 Ti* and V* record using the ARANDsoftware. The vertical dashed lines in gray present the main Milankovitch orbitalcomponents.

237Y. Zhao et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 311 (2011) 230–241

correlatedwith water depth. Strong degradation of organic materials insapropels is usually the result of active deep-water ventilation in theeastern Mediterranean, which is closely related with cold and dryevents in the northeastern Africa (De Rijk et al., 1999). It implies thatduring the interval 1000–240 kyr, at least 1000–400 kyr, the climatein the northeastern Africa is much colder and dryer than the intervalsbefore and after. However, further investigation of long-term SST andhumidity changes in the easternMediterranean is needed to understandthe exact duration and mechanism of these red intervals. Nevertheless,our results provide new evidences which suggest that post-depositional oxidation of organic material is very important in deter-mining the Corg% in sapropels.

3.3. Variations of biogenic carbonate contents

Carbonate contents of Core MD90-964 sediments are highly vari-able, with the highest (~63%) and lowest (~1%) values occur in Holo-cene and sapropel “v”, respectively (Fig. 5). Carbonate contents aregenerally decreased during deposition of most sapropelic layers.Sapropel S10, “ghost sapropels” at i-cycle 18, 26, 66, 90, 144, and150, and “hidden sapropel” at i-cycle 22, however, possess elevatedCarb% comparing to the “normal” marls. It is reported that carbonatein the Mediterranean sediments is mostly biogenic, and carbonatefraction in detrital sediments is generally negligible (Van Os et al.,1994). Biogenic carbonate in deep-sea sediments, to the best of ourknowledge, is governed by three factors: terrigenous dilution, calci-um carbonate dissolution, and productivity fluctuations of calcareousorganisms (Volat et al., 1980). Diagenesis can also strongly affect thecarbonate contents in marine sediments (Sprovieri et al., 1986). Towell understand the Carb% record of Core MD90-964, it is thereforenecessary to gain a clear idea of the control factors.

Diagenesis of loose sediments is generally negligible in shallow-burial marine environments. It has been demonstrated by previous au-thors that diagenesis is not a dominant factor of carbonate contents inloose Mediterranean sediments (Van Os et al., 1994). Water depth ofCore MD90-964 (1375 m) is well above the lysocline (present-dayequals to ~3000m), dissolution of carbonate is thus considered ofminor importance in themodern oceanographic settings. However, dis-solution may become significant during deposition of the sapropeliclayers (Van Os et al., 1994). To evaluate the degrees of diagenesis anddissolution, all samples are carefully inspected under opticalmicroscopeduring preparation of oxygen isotope analyses. Neither diagenetic alter-nation nor carbonate dissolutionwas observed on foraminiferal shells inall samples. Dissolution of carbonate in sapropel S5, S6 (Weldeab et al.,2003) and Pliocene sapropels (Thunell et al., 1991; Van Os et al., 1994;Nijenhuis and de Lange, 2000) also prove insignificant in determiningthe carbonate content. Accordingly, we suggest that diagenesis and dis-solution are less important in determining the carbonate contents inCore MD90-964.

Change in carbonate production may have also played an importantrole in interpreting the temporal variations of carbonate contents. Assuggested by previous authors, variations of Carb% in sapropels aremainly attributed to changes in carbonate productivities (Van Os etal., 1994; Weldeab et al., 2003). This judgment is drawn mostly basedon the assumption that detrital dilution is decreased during sapropelformation (Weldeab et al., 2003). Variations of the detrital input duringsapropel formation are, however, hitherto unclear. To evaluate the im-pact of carbonate productivity on carbonate contents in sediments, it istherefore necessary to investigate the carbonate production itself. Inmodern marine environments, carbonate is produced biologically,mostly in surface water, in association with photosynthesis of cocco-lithophorid phytoplankton. If variation of Carb% reflects mostly the sig-nals of carbonate production, a general decrease in abundance ofcoccoliths in sapropelic layers shall be observed. However, it is shownthat total abundances of coccoliths in sapropelic layers are not statisti-cally different with coccolith abundances in normal marls (Negri et al.,

1999), implying that no decreases of coccolith abundances actuallyoccur in sapropelic layers. Considering that total abundances of plank-tonic foraminifers are usually increased in sapropelic layers (Muerdterand Kennett, 1984; Cita et al., 1996; Negri et al., 1999), carbonate pro-duction is unlikely to decrease remarkably during deposition of mostsapropelic layers. Therefore, changes in carbonate production cannotinterpret the main variations of carbonate contents.

Detrital dilution is probably of great importance to the determinationof carbonate content in Core MD90-964, as the core is located just be-neath the Nile deep-sea fan (Fig. 1). Total terrigenous burial flux (terr.flux) is thus calculated to evaluate the significance of terrigenous dilu-tion.Downcore terr. flux canbe calculated as the product of the fractionalweight percent of detrital components [100−Carb%−Corg%], dry bulkdensity (g/cm3), and LSR (cm/kyr). The dry bulk density of the adjacentODP Site 967 sediments (Emeis et al., 1996) is adopted to represent thebulk density of Core MD90-964. It is showed that terr. flux is closely re-lated with variation of the Carb%— elevated terr. flux observed in sapro-pelic layers that possess decreased Carb%, and vice versa (Fig. 6). Itindicates that terrigenous dilution is significant in controlling the varia-tion of carbonate contents in Core MD90-964. Provided that carbonatediagenesis and dissolution are almost negligible, we suggest that car-bonate contents in Core MD90-964 sediments is mostly controlled byterrigenous dilution, with changes in carbonate productivity probablyof minor importance.

238 Y. Zhao et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 311 (2011) 230–241

3.4. Oscillations of the Nile discharges during the last 1.7 Ma

Terrigenous input to the eastern Mediterranean is generally con-sidered as a simple two “end-member” mixing system — the Nileand the Sahara (Venkatarathnam and Ryan, 1971; Krom et al.,1999; Weldeab et al., 2002). For the easternmost part of the Medi-terranean Basin, eolian dusts input from Sahara is of minor impor-tance relative to suspended materials from the Nile (Hamann etal., 2009). It has been reported that the overwhelming majority ofsuspended sediment load of the Nile is transported by the BlueNile and Atbara River, whose sources are located in the basaltic re-gion in the Ethiopian Highlands (Williams et al., 2006). Therefore, itis possible to reconstruct changes of the Nile suspended dischargeby variations of specific minerals or elements which are character-istic to sediments from the Blue Nile. As basaltic rocks are charac-terized by higher Ti contents (Wilson, 1989), changes in Ticontents in sediments of the Nile Delta are suggested to havereflected variations of Nile suspended discharge (Revel et al.,2010). Both Ti intensities and Ti* records of Core MD90-964 showstrong cyclicity, with elevated Ti contents in sapropelic layers(Fig. 6). Considering that Saharan dusts input (another potentialsource of Ti, Wehausen and Brumsack, 2000) to the Mediterraneanis greatly decreased during formation of sapropelic layers (Wehau-sen and Brumsack, 2000; Weldeab et al., 2002), we can infer that Ticontents in Core MD90-964 sediments also reflects mostly varia-tions of Nile suspended discharge.

0.01

0.03

Ti*

with

21-

kyr

filte

rT

i* w

ith 7

8-ky

r fil

ter

0.05

0.07

0.01

0.03

0.05

0.07

0 200 400 600 800

Age

Fig. 8. Comparison of the Gaussian-filtered Ti* at precessional (0.04762±0.01 kyr−1), obliquuity record from the Laskar et al. (2004) orbital solution.

Because V in seawater is very sensitive to changes in redox condi-tion, records of V contents in the Mediterranean sediments are usuallyconsidered to have reflected the remobilization of V during post-depositional degradation of organic materials (Thomson et al., 1995;Van Santvoort et al., 1997; Wehausen and Brumsack, 1999; Calvertand Fontugne, 2001). Since oxidation of organic materials is generallyconstrained to the presence of sapropelic layers, the majority of V intheMediterranean sediments is highly accumulated in sapropelic layersor at the boundaries of sapropelic layers (Thomson et al., 1995;Van Santvoort et al., 1997; Wehausen and Brumsack, 1999; Calvertand Fontugne, 2001). V contents in “normal”marl sediments are usuallyvery low, peaks of V in the sapropelic layers usually occur as “pluses” ofextreme values (Thomson et al., 1995; Van Santvoort et al., 1997;Wehausen and Brumsack, 1999; Calvert and Fontugne, 2001). Changesof V contents in Core MD90-964, however, does not follow this pattern.Oscillation of V contents generally paces the Ti contents (Fig. 6). Evi-dently, the strong cyclicity observed in oscillation of V contents cannotbe simply explained by the redox-sensitive remobilization of V duringdegradation of organic materials (Fig. 6). In contrast, the synchronousoscillation of Ti and V contents strongly implies that changes of V inCore MD90-964 are probably closely associated with the Nile sus-pended discharge. It is reported that V in Fe- or Ti-bearing oxideheavy minerals (e.g. titaniferous magnetite, vanadiferous magnetite,and rutile) can reach a content of as much as 1.5%–3.0% V (Dill, 2010).Becausemost Fe and Ti oxides are very conservative to chemical weath-ering, V species in forms of Fe and Ti oxides are usually confined in

Ti* w

ith 41-kyr filterT

i* with O

bliquity

0.01

0.03

0.05

0.07

0.01

0.03

0.05

0.07

1000 1200 16001400

(kyr)

ity (0.02439±0.004 kyr−1), and 78-kyr (0.01289±0.003 kyr−1) bands with the obliq-

239Y. Zhao et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 311 (2011) 230–241

crystal structures of these oxides. Considering that V atoms in crystalstructures are constrained from entering the seawater, redox-inducedremobilization of V in marine sediments is unlikely to occur in this sit-uation. The synchronous oscillation of V and Ti contents indicates thatV in Core MD90-964 sediments is most likely to exist in the forms ofFe and Ti oxides, rather than the mobile quinquevalent H2VO4 species(Thomson et al., 1995). Changes of V contents in Core MD90-964hence reflect mainly variations of Fe- and Ti-bearing minerals in sedi-ments. It is also observed that variations of Ti* and V* are negatively cor-related to changes of carbonate contents, with increasing/decreasing Ti*and V* correspond to decreasing/increasing carbonate contents (Fig. 6).On the contrary, intervals with elevated terr. flux are generally corre-sponding to an increasing in the Ti* and V* (Fig. 6). Provided that vari-ation of carbonate contents in Core MD90-964 is mostly controlled byterrigenous dilution, the negative correction implies that changes inTi* and V* reflect mostly variations of the Nile suspended discharge.

Spectral analysis are performed on Ti* and V* records usingthe ARAND program. Strong spectral densities at long eccentricity(414-kyr), obliquity (41-kyr), and precessional (23- and 19-kyr) cy-cles are observed in the spectra of Ti* record (Fig. 7). The V* presentsimilar spectral pattern with the Ti* record, except that periodicityat 19-kyr is insignificant (Fig. 7). The precessional signal inMediterraneanpaleoclimatic records is usually attributed to an immediate response toorbital insolation of the African monsoon (Rossignol-Strick, 1983,1985). It is also suggested that the responses of the African mon-soon to orbital insolation can also be affected by the changes inobliquity (Lourens et al., 1996a), mostly the responses to insolationforcing at high northern latitudes (Tuenter et al., 2003). The long ec-centricity signals are generally considered to correlate with presenceof large sapropel groups (Hilgen, 1991). The spectral results stronglyimply that changes of Ti* and V* are closely related with variations ofthe African monsoon. As changes of Ti* and V* reflect mostly signalsof the Nile suspended discharge, we can infer that variations of theNile suspended discharge is generally dominated by the intensitiesof Africanmonsoon, with greater Nile suspended discharge during Afri-can monsoon maxima. Changes in the intensities of African mon-soon affect both the total runoff and precipitation in the drainage ofthe Nile River, which can both have an effect on sediment discharge.However, the influences of precipitation and runoff on the sedimentdischarge are different. During maxima of African monsoon, precipita-tion in the northeastern Africa is greatly elevated, leading to a corre-sponding elevation in the Nile runoff (Ziegler et al., 2010). Anenhanced runoff in source areas is undoubtedly favorable to transporta-tion of the suspended discharges. Influences of enhanced precipitationon the suspended discharge, however, can be very complicated.On one hand, erosion potential in source areas can be largely en-hanced; on the other hand, a consequent longer and wetter rainyseason in the source areas can bring about better vegetative cover whichprotects soils from being eroded (Adamson et al., 1980; Krom et al.,2002; Rohling et al., 2002; Weldeab et al., 2002). Provided that Ti*and V* are elevated during African monsoon maxima, we can inferthat changes in suspended discharges of the Nile are more the resultof river transport capability than of erosion potential. Changes in vegeta-tive cover in the drainage cannot be the dominant factor of theNile suspended discharges variations.

In addition to the orbital cycles stated above, strong spectral den-sities at ~78 kyr periods are also observed in the Ti* and V* record(Fig. 7). Gaussian filtering of the Ti* record at different periods (pre-cessional, obliquity, and 78-kyr) further shows that the 78-kyr cyclebehaves as the dominant pacemaker of Ti* record (Fig. 8). Consider-ing that 78-kyr is almost twice the length of the obliquity cycles(41-kyr), the obliquity record is overlapped with the Ti* record to in-spect their phase relationship. It is indicated that maxima of Ti* aremostly correlated with, however not every, maxima of obliquity(Fig. 8d). Large cycles (confined by the 78-kyr cycles) of the Ti* recordare generally paced by bundles of 2–3 obliquity cycles (Fig. 8d). The

correlation between maxima of the Nile suspended discharges andmaxima of obliquity can be attributed to the responses of the Africanmonsoon to orbital insolation at high northern latitudes (Lourens etal., 1996a; Tuenter et al., 2003). However, it is hitherto not well un-derstood why the cyclicity of the Nile suspended discharge occurs inthe form of bundles of 2–3 obliquity cycles and why the bundling cy-cles are only observed in records of terrigenous sediments in theMediterranean Sea. The similar bundling of two/three obliquity cyclesinto grand cycles (obliquity subharmonics) has also been observed inthe early Pleistocene tropical SST records (Liu et al., 2008). These au-thors generally attribute the bundling to be a nonlinear response ofSST to orbital obliquity (Liu et al., 2008). Monsoon is traditionally de-fined as seasonal changes in atmospheric circulation and precipita-tion associated with the asymmetric heating of land and sea. Giventhat temperature in northeastern Africa is almost stable all throughthe year, intensity of African monsoon is thus strongly affected bythe changes in tropical Atlantic SST. Therefore, the ~78-kyr periodin Core MD90-964 V* and Ti* record is probably also a nonlinear re-sponse of Atlantic SST to orbital forcing. Nevertheless, further under-standing of the responses of African monsoon system to orbitalforcing is necessary to provide a more reasonable and detailedexplanation.

4. Conclusion

Planktonic foraminiferal oxygen isotopes, carbonate and organiccarbon contents, as well as XRF core scanning Ti, V and Ba intensitieshave been analyzed on sediments of Core MD90-964 to reconstructchanges in the Nile discharges during the last 1.75 Ma. We concludethat:

(1) A high-resolution oxygen isotopes chronology of the easternMediterranean is established based on the variation of plank-tonic foraminifer G. ruber δ18O in Core MD90-964. The bottomage of this core is estimated to be ~1747 ka, with the highest(16.67 cm/kyr) and lowest (0.17 cm/kyr) LSR values observedduring MIS 3 and MIS 10, respectively.

(2) Totally 42 dark layers are visually observed on sediment sur-face, among which 21 of them are recognized as sapropels(Corg%N1 wt.%). Dark layers with Corg%b1 wt.% are probablyresidues of oxidized sapropels. These layers can be recog-nized by the barium contents. Our results also show thatpost-depositional oxidation play a major role in determiningthe Corg% in sapropels.

(3) Variation of the Nile suspended discharge, as indicated by car-bonate contents, Ti* and V* records, is closely related with theAfrican monsoon, with higher discharge observed during Afri-can monsoon maxima. The African monsoon can affect theNile discharge via both runoff and drainage precipitation. Ourresults indicate that changes in the Nile suspended dischargesare more the result of river transport capability than of erosionpotential in source areas. It is also observed that the oscillationof the Nile suspended discharge is dominantly paced by a~78 kyr cycle. This cyclicity is probably owing to the bundlingof two/three obliquity cycles, which is a result of the nonlinearresponse of Atlantic SST to orbital forcing.

Acknowledgments

We thank Hao Wang for his technical assistance in laboratories.Dr. Lucas J. Lourens is acknowledged for providing of the MedStackdata. Samples are provided by R/V Marion Dufresne PROMETE III cruise,all participants and scientists on board are appreciated. This study is sup-ported by theNational Basic Research Programof China (2007CB815906)and the National Natural Science Foundation of China (40925008 and

240 Y. Zhao et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 311 (2011) 230–241

40876024). Yulong Zhao also gives thanks to the China ScholarshipCouncil for financial support during his habitation in France.

References

Adamson, D.A., Gasse, F., Street, F.A., Williams, M.A.J., 1980. Late Quaternary history ofthe Nile. Nature 288, 50–55.

Aksu, A.E., Yasar, D., Mudie, P.J., 1995. Origin of late glacial-Holocene hemipelagic sed-iments in the Aegean Sea — clay mineralogy and carbonate cementation. MarineGeology 123, 33–59.

Arnaboldi, M., Meyers, P.A., 2003. Geochemical evidence for paleoclimatic variationsduring deposition of two Late Pliocene sapropels from the Vrica section, Calabria.Palaeogeography, Palaeoclimatology, Palaeoecology 190, 257–271.

Arnaboldi, M., Meyers, P.A., 2007. Trace element indicators of increased primary pro-duction and decreased water-column ventilation during deposition of latest Plio-cene sapropels at five locations across the Mediterranean Sea. Palaeogeography,Palaeoclimatology, Palaeoecology 249, 425–443.

Calvert, S.E., Fontugne, M., 2001. On the late Pleistocene-Holocene sapropel record ofclimatic and oceanographic variability in the eastern Mediterranean. Paleoceano-graphy 16, 78–94.

Cita, M.B., Vergnaud-Grazzini, C., Robert, C., Chamley, C., Ciaranfi, N., D'Onofrio, S.,1977. Paleoclimatic record of a long deep sea core from the Eastern Mediterranean.Quaternary Research 8, 205–235.

Cita, M.B., Erba, E., Lucchi, R., Pott, M., Van der Meer, R., Nieto, L., 1996. Stratigraphy andsedimentation in the Mediterranean Ridge diapiric belt. Marine Geology 132,131–150.

De Rijk, S., Hayes, A., Rohling, E.J., 1999. Eastern Mediterranean sapropel S1 interrup-tion: an expression of the onset of climatic deterioration around 7 ka BP. MarineGeology 153, 337–343.

Dill, H.G., 2010. The “chessboard” classification scheme of mineral deposits: mineralo-gy and geology from aluminum to zirconium. Earth-Science Reviews 100, 1–420.

Dymond, J., Suess, E., Lyle, M., 1992. Barium in deep-sea sediment: a geochemical proxyfor paleoproductivity. Paleoceanography 7, 163–181.

Ehrmann, W., Schmiedl, G., Hamann, Y., Kuhnt, T., Hemleben, C., Siebel, W., 2007. Clayminerals in late glacial and Holocene sediments of the northern and southern Ae-gean Sea. Palaeogeography, Palaeoclimatology, Palaeoecology 249, 36–57.

Emeis, K.-C., Robertson, A.H.F., Richter, C., Blanc-Valleron, M.-M., Bouloubassi, I., Brum-sack, H.-J., Cramp, A., de Lange, G.J., di Stefano, E., Flecker, R., Frankel, E., Howell, M.W., Janecek, T.R., Jurado-Roriguez, M.-J., Kemp, A.E.S., Koizumi, I., Kopf, A., Major,C.O., Mart, Y., Pribnow, D.F.C., Rabaute, A., Roberts, A.P., Rullkötter, J.H., Sakamoto,T., Spezzaferri, S., Staerker, T.S., Stoner, J.S., Whiting, B.M., Woodside, J.M., 1996.Proceedings of the Ocean Drilling Program, Initial Reports, 160, pp. 215–287.

Emeis, K.-C., Sakamoto, T., Wehausen, R., Brumsack, H.-J., 2000. The sapropel record ofthe eastern Mediterranean Sea — results of Ocean Drilling Program Leg 160.Palaeogeography, Palaeoclimatology, Palaeoecology 158, 371–395.

Fontugne, M., Calvert, S.E., 1992. Late Pleistocene variability of the carbon isotopic com-position of organic matter in the eastern Mediterranean: monitor of changes in car-bon sources and atmospheric CO2 concentrations. Palaeoceanography 7, 1–20.

Foucault, A., Mélières, F., 2000. Palaeoclimatic cyclicity in central Mediterranean Plio-cene sediments: the mineralogical signal. Palaeogeography, Palaeoclimatology,Palaeoecology 158, 311–323.

Foucault, A., Stanley, D.J., 1989. Late Quaternary paleoclimatic oscillations in East Africarecorded by heavy minerals in the Nile delta. Nature 339, 44–46.

Hamann, Y., Ehrmann, W., Schmiedl, G., Kuhnt, T., 2009. Modern and late Quaternaryclay mineral distribution in the area of the SE Mediterranean Sea. Quaternary Re-search 71, 453–464.

Higgs, N.C., Thomson, J., Wilson, T.R.S., Croudace, I.W., 1994. Modification and completeremoval of eastern Mediterranean sapropels by postdepositional oxidation. Geology22, 423–426.

Hilgen, F.J., 1991. Extension of the astronomically calibrated (polarity) time scaleto the Miocene/Pliocene boundary. Earth and Planetary Science Letters 107,349–368.

Hinnov, L.A., 2004. Earth's orbital parameters and cycle stratigraphy. In: Gradstein, F.M., Ogg, J.G., Smith, A.G. (Eds.), A geological time scale 2004. Cambridge UniversityPress, Cambridge, pp. 55–62.

Kidd, R.B., Cita, M.B., Ryan, W.B.F., 1978. Stratigraphy of eastern Mediterranean sapro-pel sequences recovered during DSDP Leg 42A and their paleoenvironmentalsignificance. In: Hsü, K.J., Montadert, L., Garrison, R.B., Fabricius, F.H., Kidd, R.B.,Müller, C., Cita, M.B., Bizon, G., Wright, R.C., Erickson, A.J., Bernoulli, D., Mélière,F. (Eds.), InitialReports of the Deep Sea Drilling Program, volume 42. NationalScience Foundation, Washington DC, pp. 421–443 (Part. 1).

Krom, M.D., Cliff, R.A., Eijsink, L.M., Herut, B., Chester, R., 1999. The characterisation ofSaharan dusts and Nile particulate matter in surface sediments from the Levantinebasin using Sr isotopes. Marine Geology 155, 319–330.

Krom, M.D., Stanley, D.J., Cliff, R.A., Woodward, J.C., 2002. Nile River sediment fluctua-tions over the past 7000 yr and their key role in sapropel development. Geology30, 71–74.

Kroon, D., Alexander, I., Little, M., Lourens, L.J., Matthewson, A., Robertson, A.H.F.,Sakamoto, T., 1998. Oxygen isotope and sapropel stratigraphy in the easternMediterranean during the last 3.2 million years. In: Robertson, A.H.F., Emeis, K.-C.,Richter, C., Camerlenghi, A. (Eds.), Proceedings of the Ocean Drilling Program, Scien-tific Results 160. College Station (Ocean Drilling Program), Texas, pp. 181–190.

Langereis, C.G., Dekkers, M.J., de Lange, G.J., Paterne, M., van Santvoort, P.J.M., 1997.Magnetostratigraphy and astronomical calibration of the last 1.1 Myr from an

eastern Mediterranean piston core and dating of short events in the Brunhes. Geo-physical Journal International 129, 75–94.

Larrasoaña, J.C., Roberts, A.P., Hayes, A., Wehausen, R., Rohling, E.J., 2006. Detectingmissing beats in the Mediterranean climate rhythm from magnetic identificationof oxidized sapropels (Ocean Drilling Program Leg 160). Physics of the Earth andPlanetary Interiors 156, 283–293.

Laskar, J., Robutel, P., Joutel, F., Gastineau, M., Correia, A.C.M., Levrard, B., 2004. A longterm numerical solution for the insolation quantities of the Earth. Astronomy andAstrophysics 428, 261–285.

Liu, Z., Cleaveland, L.C., Herbert, T.D., 2008. Early onset and origin of 100-kyr cyclesin Pleistocene tropical SST records. Earth and Planetary Science Letters 265,703–715.

Lourens, L.J., Antonarakou, A., Hilgen, F.J., Van Hoof, A.A.M., Vergnaud-Grazzini, C.,Zachariasse, W.J., 1996a. Evaluation of the Plio-Pleistocene astronomical timescale.Paleoceanography 11, 391–413.

Lourens, L.J., Hilgen, F.J., Raffi, I., Vergnaud-Grazzini, C., 1996b. Early Pleistocene chro-nology of the Vrica section (Calabria, Italy). Paleoceanography 11, 797–812.

Lourens, L.J., Hilgen, F., Shackleton, N.J., Laskar, J., Wilson, D., 2004. The Neogene period.In: Gradstein, F.M., Ogg, J.G., Smith, A.G. (Eds.), A geological time scale 2004.Cambridge University Press, Cambridge, pp. 409–440.

Mascle, J., Sardou, O., Loncke, L., Migeon, S., Caméra, L., Gaullier, V., 2006. Morphostruc-ture of the Egyptian continental margin: insights from swath bathymetry surveys.Marine Geophysical Researches 27, 49–59.

Métivier, F., Gaudemer, Y., 1999. Stability of output fluxes of large rivers in South andEast Asia during the last 2 million years: implications on floodplain processes.Basin Research 11, 293–303.

Milliman, J.D., Farnsworth, K.L., 2011. River Discharge to the Coastal Ocean — A GlobalSynthesis. Cambridge University Press, Cambridge. 384 pp.

Milliman, J.D., Syvitski, J.P.M., 1992. Geomorphic/tectonic control of sediment dis-charge to the ocean: the importance of small mountainous rivers. Journal of Geology100, 525–544.

Muerdter, D.R., Kennett, J.P., 1984. Late Quaternary planktonic foraminiferal biostratig-raphy, Strait of Sicily, Mediterranean Sea. Marine Micropaleontology 8, 339–359.

Murat, A., Got, H., 2000. Organic carbon variations of the eastern Mediterranean Holo-cene sapropel: a key for understanding formation processes. Palaeogeography,Palaeoclimatology, Palaeoecology 158, 241–257.

Negri, A., Capotondi, L., Keller, J., 1999. Calcareous nannofossils, planktonic foraminif-era and oxygen isotopes in the late Quaternary sapropels of the Ionian Sea. MarineGeology 157, 89–103.

Nijenhuis, I.A., de Lange, G.J., 2000. Geochemical constraints on Pliocene sapropel for-mation in the eastern Mediterranean. Marine Geology 163, 41–63.

Paillard, D., Labeyrie, L., Yiou, P., 1996. Macintosh program performs time-series anal-ysis. Eos, Transactions American Geophysical Union 77, 379.

Passier, H.F., Bosch, H.-J., Nijenhius, I.A., Lourens, L.J., Böttcher, M.E., Leenders, A.,Damstè, J.S., De Lange, G.J., De Leeuw, J.W., 1999. Sulphidic Mediterranean surfacewaters during Pliocene sapropel formation. Nature 397, 146–149.

Revel, M., Ducassou, E., Grousset, F.E., Bernasconi, S.M., Migeon, S., Revillon, S., Mascle,J., Murat, A., Zaragosi, S., Bosch, D., 2010. 100,000 years of African monsoon vari-ability recorded in sediments of the Nile margin. Quaternary Science Reviews 29,1342–1362.

Rinna, J., Warning, B., Meyers, P.A., Brumsack, H.-J., Rullkötter, J., 2002. Combined or-ganic and inorganic geochemical reconstruction of paleodepositional conditionsof a Pliocene sapropel from the eastern Mediterranean Sea. Geochimica et Cosmo-chimica Acta 66, 1969–1986.

Rohling, E., Cane, T.R., Cooke, S., Sprovieri, M., Bouloubassi, I., Emeis, K.C., Schiebel, R.,Kroon, D., Jorissen, F.J., Lorre, A., Kemp, A.E.S., 2002. African monsoon variabilityduring the previous interglacial maximum. Earth and Planetary Science Letters202, 61–75.

Rossignol-Strick, M., 1983. African monsoons, an immediate climate response to orbitalinsolation. Nature 304, 46–49.

Rossignol-Strick, M., 1985. Mediterranean Quaternary sapropels, an immediate re-sponse of the African monsoon to variation of insolation. Palaeogeography, Palaeo-climatology, Palaeoecology 49, 237–263.

Rossignol-Strick, M., Nesteroff, W.D., Olive, P., Vergnaud- Grazzini, C., 1982. After thedeluge: Mediterranean stagnation and sapropel formation. Nature 295, 105–110.

Rossignol-Strick, M., Paterne, M., Bassinot, F., Emeis, K.-C., de Lange, G.J., 1998. Anunusual mid-Pleistocene monsoon period over Africa and Asia. Nature 392, 269–272.

Roussakis, G., Karageorgis, A.P., Conispoliatis, N., Lykousis, V., 2004. Last glacial-Holocene sediment sequences in N. Aegean basins: structure, accumulation ratesand clay mineral distribution. Geo-Marine Letters 24, 97–111.

Sprovieri, R., Thunell, R.C., Howell, M.W., 1986. Paleontological and geochemical ana-lyses of three laminated sedimentary units of the late Pliocene-early Pleistoceneage from the Monta San Nicola section in Sisily. Riv. Ital. Paleontol. 92, 401–434.

Thomson, J., Higgs, N.C., Wilson, T.R.S., Croudace, I.W., de Lange, G.J., van Santvoort, P.J.M., 1995. Redistribution and geochemical behavior of redox-sensitive elementsaround S1, the most recent eastern Mediterranean sapropel. Geochimica et Cosmo-chimica Acta 59, 3487–3501.

Thunell, R., Rio, D., Sprovieri, R., Raffi, I., 1991. Limestone-marl couplets: origin of theearly Pliocene Trubi marls in Calabria, southern Italy. Journal of Sedimentary Re-search 61, 1109–1122.

Tuenter, E., Weber, S.L., Hilgen, F.J., Lourens, L.J., 2003. The response of the African sum-mer monsoon to remote and local forcing due to precession and obliquity. Globaland Planetary Change 36, 219–235.

Van Os, B.J.H., Lourens, L.J., Hilgen, F.J., de Lange, G.J., Beaufort, L., 1994. The formationof Pliocene sapropels and carbonate cycles in the Mediterranean: diagenesis, dilu-tion, and productivity. Paleoceanography 9, 601–617.

241Y. Zhao et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 311 (2011) 230–241

Van Santvoort, P.J.M., de Lange, G.J., Thomson, J., Cussen, H., Wilson, T.R.S., Krom, M.D.,Strohle, K., 1996. Active post-depositional oxidation of the most recent sapropel(S1) in sediments of the eastern Mediterranean Sea. Geochimica et CosmochimicaActa 60, 4007–4024.

Van Santvoort, P.J.M., de Lange, G.J., Langereis, C.G., Dekkers, M.J., Paterne, M., 1997.Geochemical and paleomagnetic evidence for the occurrence of “'missing“' sapro-pels in eastern Mediterranean sediments. Paleoceanography 12, 773–786.

Van Santvoort, P.J.M., De Lange, G.J., Thomson, J., Colley, S., Meysman, F.J.R., Slomp, C.P.,2002. Oxidation and origin of organic matter in surficial Eastern Mediterraneanhemipelagic sediments. Aquatic Geochemistry 8, 153–175.

Venkatarathnam, K., Ryan, W.B.F., 1971. Dispersal patterns of clay minerals in the sed-iments of the eastern Mediterranean Sea. Marine Geology 11, 261–282.

Volat, J.-L., Pastouret, L., Vergnaud-Grazzini, C., 1980. Dissolution and carbonate fluctu-ations in Pleistocene deep-sea cores: a review. Marine Geology 34, 1–28.

Wehausen, R., Brumsack, H.-J., 1999. Cyclic variations in the chemical composition ofthe eastern Mediterranean Pliocene sediments: a key for understanding sapropelformation. Marine Geology 153, 161–176.

Wehausen, R., Brumsack, H.-J., 2000. Chemical cycles in the Pliocene sapropel-bearingand sapropel-barren eastern Mediterranean sediments. Palaeogeography, Palaeo-climatology, Palaeoecology 158, 325–352.

Weldeab, S., Emeis, K.-C., Hemleben, C., Schulz, H., Vennemann, T.W., 2002. Sr, Nd iso-tope composition of late Pleistocene sapropels and non-sapropel sediments fromthe eastern Mediterranean Sea: implications for detrital influx and climatic condi-tions in the source areas. Geochimica et Cosmochimica Acta 66, 3585–3598.

Weldeab, S., Emeis, K.-C., Hemleben, C., Schmiedl, G., Schulz, H., 2003. Spatial produc-tivity variations during formation of sapropels S5 and S6 in the MediterraneanSea: evidence from Ba contents. Palaeogeography, Palaeoclimatology, Palaeoecol-ogy 191, 169–190.

Williams, M., Talbot, M., Aharon, P., Salaam, Y.A., Williams, F., Brendeland, K.I., 2006.Abrupt return of the summer monsoon 15,000 years ago: new supporting evidencefrom the lower White Nile valley and Lake Albert. Quaternary Science Reviews 25,2651–2665.

Wilson, M., 1989. Igneous Petrogenesis. Unwin Hyman, London.Ziegler, M., Tuenter, E., Lourens, L.J., 2010. The precession phase of the boreal summer

monsoon as viewed from the eastern Mediterranean (ODP Site 968). QuaternaryScience Reviews 29, 1481–1490.