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Multi-tracer study of continental erosion and sediment transport to the Red Sea and the Gulf of Aden during
the last 20 ka
Virginia Rojas, Laure Meynadier, Christophe Colin, Franck Bassinot, Jean-Pierre Valet, Serge Miska
To cite this version:
Virginia Rojas, Laure Meynadier, Christophe Colin, Franck Bassinot, Jean-Pierre Valet, et al..
Multi-tracer study of continental erosion and sediment transport to the Red Sea and the Gulf of Aden during the last 20 ka. Quaternary Science Reviews, Elsevier, 2019, 212, pp.135-148.
�10.1016/j.quascirev.2019.02.033�. �hal-02149421�
Multi-tracer study of continental erosion and sediment transport to the Red Sea and 1
the Gulf of Aden during the last 20 ka 2
3
VIRGINIA P.ROJAS1*,LAURE MEYNADIER1,CHRISTOPHE COLIN2,FRANCK BASSINOT3, 4
JEAN-PIERRE VALET1,SERGE MISKA2
5
1 Institut de Physique du Globe de Paris, Université Paris Diderot, Sorbonne Paris‐Cité, 6
UMR 7154 CNRS, 1 rue Jussieu, FR‐75238 Paris Cedex 05, France.
7
(virginia_ve@yahoo.com).
8
2 Laboratoire Géosciences Paris-Sud (GEOPS), UMR 8148, CNRS-Université de Paris-Sud, 9
Université Paris-Saclay, Bâtiment 504, 91405 Orsay Cedex, France.
10
3 Laboratoire des Sciences du Climat et de l’Environnement (CEA/CNRS/UVSQ), Domaine 11
du CNRS, Bât. 12, avenue de la Terrasse, F-91198 Gif-sur-Yvette Cedex, France.
12 13
Abstract 14
15
Mineralogical compositions and grain-size distributions combined with 87Sr/86Sr and εNd
16
values of the detrital fraction were studied on cores recovered from the Gulf of Aden (MD92- 17
1002) and the Red Sea (MD92-1008) basins in order to document past changes in Indian 18
monsoon and northwesterly winds during the last glacial-interglacial transition (the last 20 19
ka), encompassing the African Humid Period (AHP). The εNd vs. 87Sr/86Sr plot indicates that 20
sediments result from the mixing of two main sedimentary sources corresponding to the Afar 21
volcanic rocks in Ethiopia and to the Arabian-Nubian Shield. Variations of sediment isotopic 22
and mineralogical composition point to a diminution of the volcanic source contribution 23
during the last deglaciation. Changes of mineral-accumulation rates and grain-size 24
distributions denote a decline in the aridity of the source regions during the Holocene, 25
Version of Record: https://www.sciencedirect.com/science/article/pii/S0277379118301811 Manuscript_c402d5b4650c80fa23684944503e4489
particularly of the Afar volcanic region. In this area, the reduction of detrital supply, from 15 26
cal ka BP, can be explained by an increase of precipitations during the AHP, which resulted in 27
an expansion of the vegetation cover and lake extensions in East Africa. In the Arabian 28
Peninsula, precipitations were confined to the south, allowing sediments to be transported 29
even during the Holocene. Our data suggest that the southwest monsoon was not the main 30
carrier of aeolian sediments to the Red Sea and Gulf of Aden basins, but the Northwesterlies.
31
In the Red Sea, the isotopic and mineralogical tracers reveal a contribution from Saharan dust 32
between 16 and 12 cal ka BP, transported from the Nile catchment after aridification during 33
Heinrich event 1.
34 35
Keywords: radiogenic isotopes, clay mineralogy, siliciclastic grain-size, Red Sea, Gulf of 36
Aden, Last Glacial Period, African Humid Period.
37 38
1. Introduction 39
40
The Red Sea and the Gulf of Aden (GOA) basins are located in the midst of the desert belt 41
between the Saharan and Arabian deserts, with precipitations below 100 mm/year (Locke and 42
Tunnel, 1988) and where aeolian erosion plays a key role, due to the sparse vegetation cover 43
in the area. In this region, precipitations and low-altitude wind circulation show marked 44
seasonal changes, typical of a monsoon system (Webster, 1981; Laing and Evans, 2011).
45
Southwest winds blow over the Arabian Sea, from the Horn of Africa to the Indian 46
subcontinent, during the wet summer monsoon. In contrast, the Northeast winds prevail 47
during the dry winter monsoon (Clemens and Prell, 1991; Sirocko et al., 1991) (Fig. 1).
48
However, low-altitude atmospheric circulation patterns in the region are more complex than a 49
simple SW-NE seasonal reversal. They reflect the influence of the Inter-Tropical 50
Convergence Zone (ITCZ) shifting position and of local topography (Fleitmann et al., 2007;
51
Lezine et al., 2007, 2010). Above the monsoonal wind system, the mid-tropospheric 52
circulation is dominated by northwesterly winds from the Arabian Peninsula, which appear to 53
play an important role in the transport of dust to the western Arabian Sea (Sirocko et al., 54
1991; Sirocko and Lange, 1991; Sirocko et al., 1993; Pourmand et al., 2004), and may also 55
influence the sediment transport to the Red Sea and GOA basins.
56
On orbital timescale and particularly across the last glacial-interglacial transition, many 57
studies have documented major changes in monsoon climates in response to variations in 58
insolation. Continental and marine records from Asia and North Africa suggest that the 59
summer monsoon intensity in the northern hemisphere was weaker 18-20 ka ago and 60
increased to reach a maximum in early-mid Holocene, coeval with the maximum northern 61
hemisphere summer insolation (Overpeck et al., 1996; Gasse and Van Campo, 1994; Gasse, 62
2000; Enzel et al., 1999; Sirocko et al., 1991; Fleitmann et al, 2003, 2007; Revel et al., 2014).
63
Similarly, changes in the flux of aeolian dust brought to the Arabian Sea were documented in 64
relation to variations in Indian monsoon intensity. Sedimentary record from core 74KL, 65
located off the Oman coast, revealed a diminution of windborne dolomite particles since the 66
Last Glacial Maximum (LGM), with a striking minimum corresponding to the early Holocene 67
at the maximum of southwest monsoon intensity (Sirocko et al., 1993).
68
In the northern half of the African continent, the humid episode that occurred during the first 69
part of the Holocene is known as the African Humid Period (AHP). Increase of precipitations 70
over this time period resulted in the rise of lake levels and the extension of vegetation cover 71
(Enzel et al, 1999; Gasse, 2000; deMenocal and Tierney, 2012; Shanahan et al., 2015; Lezine 72
et al., 2007; 2010; Revel et al., 2015). A detailed and well-dated sedimentary record at ODP 73
Site 658C (off Cap Blanc, Mauritania) suggests very abrupt, large-scale changes in aeolian 74
sediment transport with a well-defined period of low influx between 12.3 and 5.5 cal. ka BP 75
associated with the AHP, when the Sahara was nearly completely covered by vegetation 76
(Adkins et al., 2006). Continental data from the African continent and the Arabian Peninsula 77
indicate that the beginning and end of this humid period may have been progressive (Lézine 78
et al, 1998; 2010; 2011), whereas the marine windborne records from the region show abrupt 79
transitions that are out-of-phase with the continental evidence (e.g. abrupt decrease of aeolian 80
proxies as early as ~ 15,000 cal yr BP; Lezine et al., 2014). These differences between marine 81
and continental records may be related to the dependence of aeolian sedimentation on changes 82
in wind intensity and/or direction (Lezine et al., 2014).
83
The study of the amount and composition of aeolian dust in oceanic sediments is of great 84
importance because of the influence they have on climatic, pedogenic and ecological 85
conditions in the Earth. Atmospheric dust can have an effect on the Earth’s radioactive 86
balance, since it either absorbs or reflects the solar energy depending on its mineralogy, grain- 87
size and distribution. The erosion of clay minerals and removal from the source areas can also 88
cause the degradation of soils and the supply of nutrients to marine environments, which in 89
turn can have an effect over the removal of carbon dioxide from the atmosphere (Grousset 90
and Biscaye, 2005; Maher et al., 2010; Scheuvens et al., 2013). In order to have a better 91
understanding of the factors affecting dust transport and climatic conditions, methods that can 92
relate aeolian sediment records to its contemporaneous source areas are required. Various 93
types of fingerprint tracers can be applied, including elemental concentrations, mineralogy 94
and grain-size distributions (Grousset and Biscaye, 2005).
95
In the same way, Strontium and Neodymium isotopic ratios can be used as reliable source 96
tracers. The radiogenic isotopic compositions of Nd and Sr are very different in mantle 97
derived vs. crust derived rocks, allowing the distinction to be made between young volcanic 98
areas and old continental shields (Goldstein and Hemming, 2003; Grousset and Biscaye, 99
2005). Geological formations surrounding the Red Sea and GOA basins present a large 100
variety of lithologies, with distinct Nd and Sr isotope signatures for crustal terrains of the late 101
Proterozoic Arabian–Nubian Shield (ANS) (Stoeser and Frost, 2006; Stein and Goldstein, 102
1996), the Saharan Shield (Kuster et al., 2008), Phanerozoic platform sediments and Cenozoic 103
alkali basalts from the Afar region (Betton and Civetta, 1984; Teklay et al., 2009) (Fig. 1a).
104
Previous studies have shown the interest of using mineralogical and geochemical signatures to 105
track changes in the sedimentary material from this region. Sirocko et al (1991) and Sirocko 106
and Lange (1991) characterized the mineralogy of sediments in the Arabian Sea and 107
determined that the Arabian Peninsula is a major source of Arabian Sea deep-sea sediments, 108
transported by northwesterly winds. Aeolian fluxes in the Arabian Sea were found to be 109
higher during glacial periods due to the discharge from northwesterly winds that was more 110
intense when the southwest monsoon was weaker (Sirocko et al. 2000; Pourmand et al., 111
2004). Nd and Sr isotopic measurements in sediments throughout this basin allowed the 112
distinction to be made between western Arabian Sea sediments, which arrived from North and 113
Central Arabian Peninsula, and eastern Arabian Sea, that presented riverine contributions 114
from India (Sirocko, 1994).
115
Stein et al. (2007) measured Sr isotopes in two sediment cores in the Red Sea and the GOA, 116
which covered a time period of 500 ka with a 10 to 30 ka resolution. They identified a 117
hydrothermal Sr component associated with sea-floor spreading in the Red Sea, superimposed 118
with a granitic Arabian source and a loess source. Nd and Sr isotopic compositions from 119
sediment cores in the Northern and Central Red Sea revealed sedimentary sources like the 120
Saharan granitoids, the Arabo-Nubian Shield terrains and basalts from the Ethiopian 121
Highlands (Palchan et al, 2013). Nonetheless, these results do not show appreciable variations 122
between the Holocene and the LGM, suggesting that there were not any significant changes in 123
the northern Red Sea during this period. Similarly, Jung et al. (2004) did not find any source 124
variation in the sediments from core 905 in the Arabian Sea off the Somalian coast during the 125
Holocene (stable Nd isotopic composition), and attributed the variations in the 87Sr/86Sr ratio 126
to weathering changes.
127
In the present study, we examined the Sr and Nd isotopic composition, clay and bulk 128
mineralogy, and grain size changes in sediments from the GOA and southern Red Sea in order 129
to determine the sedimentary sources and transport processes and to establish their variability 130
since the LGM. For this purpose, we studied cores MD92-1002 (GOA) and MD92-1008 (Red 131
Sea), which provide continuous sedimentary records deposited at high sedimentation rates 132
since the LGM (Fig. 1, Bouilloux et al., 2013a, 2013b; Fersi et al., 2016). These cores are 133
positioned near the northern limit of the southwest monsoon winds, whose intensity changed 134
during the last glacial-interglacial transition (Van Campo et al., 1982; Sirocko et al., 1991;
135
Overpeck et al., 1996; Gasse and Van Campo, 1994; Gasse, 2000; Fleitmann et al., 2003, 136
2007; Revel et al., 2014; Fersi et al., 2016), and are surrounded by very contrasted geological 137
formations making them particularly suitable to the study of sediment-source variations. The 138
region was also subjected to the AHP (Gasse, 2000; deMenocal and Tierney, 2012;), which 139
has been documented by changes in the sedimentary record from around 15 ka until 5 ka BP 140
(Juginger and Trauth, 2013; Juginger et al., 2014; Shanahan et al., 2015; Revel et al., 2015).
141
The results make it possible to re-evaluate the role of humidity changes and the relative 142
influence of the northwesterly winds and the southwest monsoon winds on the origin and 143
transport of dust to the GOA and the Red Sea since the LGM, across the deglaciation and the 144
onset of the early Holocene humid period.
145
146
Figure 1: Location of cores MD92-1002 (MD2) in the GOA and MD92-1008 (MD8) in the 147
Red Sea. a) Map of the study zone showing the main geologic units (modified from Stein 148
et al., 2007). Location of cores KL23 and KL11 in the Red Sea from Palchan et al. (2013) 149
and core NIOP 905 in the Arabian Sea from Jung et al. (2004) is also shown. Surface wind 150
conditions (1000 hP) for (b) winter and (c) summer months (ECMWF analyses 1990- 151
1997) from Lezine et al., (2014). (For interpretation of the references to colour in this 152
figure legend, the reader is referred to the Web version of this article).
153
154
2. Materials and Methods 155
156
2.1. Chronology of the cores 157
158
Cores MD92-1002 and MD92-1008 were retrieved during the MD73 REDSED campaign of 159
R/V Marion Dufresne in 1992. We will refer to core MD92-1002 (GOA) as MD2 and to core 160
MD92-1008 (Red Sea) as MD8 to simplify the notation. Core MD2 is located at 12°01’19’’
161
N and 44°12’01’’ E and at 1327 m of water depth. Core MD8 is located 230 km to the north 162
of the Bab-el-Mandeb strait at a position of 14°25’86’’ N, 42°13’62’’ E and at a water depth 163
of 708 m (Fig. 1).
164
The carbonate content of samples is very similar for both cores and varies between 25% and 165
70%. 14C dating was performed on planktonic foraminifera (Globigerinoides ruber and 166
Globigerinoides sacculifer) and published in Bouilloux et al. (2013a, 2013b). The 18O values 167
were analyzed on planktonic foraminifera Globigerinoides ruber every 5 cm by Bouilloux et 168
al. (2013a and 2013b), with additional data from Fersi et al. (2016). δ 18O curve combined 169
with 14C dating made it possible to establish the age model for the cores. Both cores reached 170
the LGM and recorded the deglaciation and the Holocene period. The resulting age models 171
indicate that core MD2 covers a period from 20 cal ka BP to the present and has a mean 172
sedimentation rate of 52 cm/ka. Core MD8 was studied from 30 cal ka BP to the present, 173
displaying a mean sedimentation rate of 30 cm/ka during this period (Fig 2).
174 175
176
Figure 2: δ18O planktonic foraminifera (G. ruber) values and age model for cores in this 177
study: a) MD2 (GOA) from Fersi et al. (2016) and Bouilloux et al. (2013a) b) MD8 (Red Sea) 178
from Bouilloux et al. (2013b). Climatic periods include the LGM (Last Glacial Maximum), 179
Bølling Allerød (BA), Younger Dryas (YD) and the Holocene. c) Depth in core as a function 180
of Age cal ka BP from the age model based on 14C measurements on planktonic foraminifera.
181 182
2.2. Non-calcareous bulk and clay mineralogical analyses 183
184
X-ray diffraction (XRD) analyses were performed using a PANalytical X’Pert Pro 185
Diffractometer (GEOPS Laboratory, Université Paris-Sud) with CuKα radiation and Ni filter, 186
under 40 kV and a current intensity of 25 mA. Bulk sediments were leached using acetic acid 187
3.5 M to remove the carbonate fraction. The samples were then rinsed several times to 188
remove the acid residues. Part of the carbonate-free fraction was dried, finely grinded in an 189
agate mortar and the resulting powders were analyzed by XRD to identify the mineralogy and 190
determine the proportion of the main non-carbonate minerals in the sediments. Errors in 191
mineral percentages were estimated to be around 1-2%.
192
Identification of clay minerals was conducted on oriented mounts of the non-calcareous clay 193
fraction (<2 μm), following the procedure described by Holtzapffel (1985). Particles smaller 194
than 2 µm were separated by gravitational settling (from Stockes law) and centrifugation. The 195
measurements were performed under three different conditions: air-dried, ethylene glycol- 196
saturated and heated (490 °C for 2 hours). The samples were X-rayed in the range 3–30° 2Θ 197
with a step size of 0.03° 2Θ and a measuring time of ~ 1 second/step. The proportions of 198
different clay groups were determined semi-quantitatively using peak areas of basal 199
reflections for the main clay mineral groups on ethylene glycol saturated samples. Replicate 200
analyses of selected samples show that errors for the semi-quantitative estimation of clay 201
minerals smectite, palygorskite, illite, kaolinite and chlorite are about 2%. Relative 202
proportions of kaolinite and chlorite were determined by area ratios for the (002) peak of 203
kaolinite (3.57 Å) and the (004) peak of chlorite (3.54 Å). In cases where the proportions of 204
kaolinite or chlorite are less than 10%, errors for these two minerals can reach 3-4%. Relative 205
proportions of different clay groups were given in percentages of the total clay assemblage.
206
Determinations were performed using MacDiff software.
207 208
2.3 Laser grain-size analyses 209
210
Grain-size distribution measurements of the carbonate-free, terrigenous particles (from 0.02 to 211
2000 μm) were carried out on a Malvern Mastersizer 2000 Particle Size Analyzer at the 212
GEOPS Laboratory. Prior to the analyses, the bulk fraction of the sediments was pre-treated 213
with hydrogen peroxide 0.2 M and acetic acid 3.5 M to remove the organic matter and the 214
carbonate fraction, respectively. The mixture was then rinsed several times in order to remove 215
any acid residue. This aqueous suspension was then poured into the fluid module of the 216
particle-size analyzer. Ultrasonic oscillation was not used during these measurements as 217
previous studies had shown that ultrasonic dispersion could break brittle minerals (e.g. micas) 218
(Trentesaux et al., 2001). In order to diminish errors, each sample was measured twice and the 219
mean of both measurements was taken to represent the grain-size distribution for each sample.
220 221
2.3. Sr and Nd isotopic compositions 222
223
Sr and Nd isotope analyses were carried out on 22 samples from each core at a sampling 224
resolution of around 40 cm. εNd values were obtained following the analytical procedure 225
described in detail by Le Houedec et al. (2012). The carbonate fraction was dissolved by 226
chemical leaching using acetic acid in excess (8 mL of acetic acid 1.6 M) to ensure total 227
dissolution of 400 mg of dried samples. Then, the residue was leached with HBr 1 M and 228
rinsed repeatedly with an alternation of HBr and distilled water in order to eliminate possible 229
authigenic phases.
230
Carbonate-free samples were then dissolved by chemical attack using a mixture of 231
concentrated HF and HNO3 in equal proportions (the amount was adjusted depending on the 232
sample weight). A subsequent attack of H3BO3 and HNO3 was performed in order to dissolve 233
the calcium fluorides that could have formed. The Rare Earth Elements (REE) were then 234
separated from major elements in the sample by ionic chromatography, using Triskem TRU 235
resin. Nd was obtained by separation from the other REE using Triskem LN resin. Finally, Sr 236
was separated from the major element fraction using SR-Spec resin.
237
Nd and Sr isotopic compositions were analyzed on carbonate-free samples using a Thermo 238
Finnigan Neptune Multiple Collector Inductively Coupled Plasma Mass Spectrometer (MC- 239
ICP-MS) at the GEE Laboratory, IPGP (Paris, France). The Sr and Nd isotope ratios were 240
corrected from mass bias according to the exponential law relative to 146Nd/144Nd = 0.7219 241
and 86Sr/88Sr = 0.1194. Sr values were corrected for Kr and Rb isobaric interferences by 242
measuring the 83Kr/86Kr, 83Kr/84Kr and 85Rb/87Rb ratios.
243
The accuracy was monitored by sequential analyses of NIST 3135a standard (May and 244
Rumble, 2003) for Neodymium and NBS SRM 987 standard for Strontium (1 standard every 245
3 samples) and subsequent correction of the 143Nd/144Nd ratio (0.511418) and 87Sr/86Sr ratio 246
(0.71025). Repeated measurements of standard NBS SRM 987 yielded a mean 87Sr/86Sr value 247
of 0.71030 ± 1 (2σ, n=21). The measured value for standard NIST 3135a was of 0.511409 ± 248
8 (2σ, n=16). Reference standard JNdi (143Nd/144Nd ratio = 0.512115) was also used for 249
comparison, being the measured value 0.512109 ± 6 (2σ, n=7). The 143Nd/144Nd ratios are 250
expressed as εNd values, calculated from the difference between the sample and the CHUR 251
standard (143Nd/144Nd= 0.512638) as:
252
εNd= {[(143Nd/144Nd)sample /(143Nd/144Nd)CHUR] -1}× 10000 253
The precision was evaluated by sequential measurements of the same sample. The estimated 254
error (2σ) on εNd values is ± 0.2, and ± 2.10-5 for 87Sr/86Sr.
255 256
3. Results 257
258
3.1. Mineralogy 259
260
3.1.1. Mineralogical composition of carbonate-free bulk samples 261
262
In general, the bulk mineralogical composition of both cores is very similar, with a 263
predominance of quartz and plagioclase, suggesting a mixing of granitoid and basaltic 264
sources. In core MD2 (GOA) main minerals are quartz (10%-30%), plagioclase (5%-30%), 265
micas (biotite/muscovite) (15%-30%), potassium feldspar (5%-15%), chlorite (5%-15%) and 266
amphibole (5%-15%). Pyroxene is present in some samples, ranging from 5% to 20%. The 267
noteworthy changes in mineralogy include a ~ 10% increase of quartz content from the glacial 268
period to the Holocene (from 20% to 30%), and the presence of pyroxene only in glacial 269
sediments (Fig. 3a).
270
In core MD8 (Red Sea) sediments comprise quartz (15%-30%), plagioclase (5%-30%), micas 271
(biotite/muscovite) (5%-30%), potassium feldspar (5%-20%), chlorite (5%-15%) and 272
amphibole (5%-10%). Pyroxene is present in some samples from 5% to 10%. Unlike what 273
was observed for the GOA record, the Red Sea core does not display significant variations in 274
quartz content from the glacial period to the Holocene. Pyroxene is only observed in glacial 275
period sediments, in the same way as for the GOA (Fig. 3a).
276
The presence of ferromagnesian minerals like amphibole and pyroxene, along with 277
plagioclase as the dominant type of feldspar in the sediments, is an indication of the basaltic 278
character of the samples. Such mineralogy also reflects the low degree of chemical alteration 279
that some of the minerals underwent, probably because of the limited precipitations in the 280
source regions. The low weathering degree could also reflect a short-transport distance of the 281
sediments after erosion by glaciers present in the mountains of East Africa (as observed by 282
Kelly et al., 2014 and Hendrickx et al., 2015 during the Pleistocene).
283 284
285
Figure 3: Mineral composition as measured by XRD analysis for GOA and Red Sea 286
sediments a) Mineral percentages of bulk decarbonated sample, b) clay mineral percentages 287
from fraction < 2 μm. Variations in the GOA include an increase of palygorskite and quartz 288
and a diminution of smectite and pyroxene during the Holocene. In the Red Sea pyroxene is 289
present only during the glacial period and smectite diminishes from 16 to 12 cal ka BP. (For 290
interpretation of the references to colour in this figure legend, the reader is referred to the 291
Web version of this article).
292 293
3.1.2. Clay mineralogy (fraction <2 μm) 294
295
Clay minerals are a minor component of the sediments (<5%), however they can be 296
particularly helpful to reconstruct past environmental conditions since they derive from the 297
alteration of primary minerals and their genesis and evolution depend on lithological, climatic 298
and morphological factors (Chamley, 1989). In the clay size fraction of core MD2 (GOA), 299
smectite (50-70%) is the dominant clay mineral while palygorskite (5-20%.), illite (10-15%), 300
chlorite (10-15%) and kaolinite (5-10%) are present in lesser quantities (Fig. 3b). In general, 301
smectite variations are inversely correlated to those of palygorskite and the other clay 302
minerals (See graphs in supplementary material 1). We notice a diminution of smectite and an 303
increase of palygorskite from MIS 2 to the Holocene (Fig. 4d and 4e). Illite shows a small 304
increase in the last glacial-interglacial transition, while chlorite and kaolinite do not display 305
significant variations over this time period (Fig 3b).
306
Regarding the clay mineralogical composition of core MD8 (Red Sea), smectite (55-80%) is 307
also the dominant clay mineral along with a significant amount of palygorskite (5-10%). Illite, 308
chlorite and kaolinite range from 5% to 15% (Fig 3b). The clay composition does not show 309
the same variation from MIS 2 to the Holocene as in the GOA basin, but there is a strong 310
decrease in smectite content counterbalanced by an increase in the relative proportion of the 311
other clay minerals during the time interval from 16 to 12 cal ka BP (Fig 3b and 4e).
312
313
Figure 4: Isotope and mineralogical data for the GOA and the Red Sea basins. Left: GOA 314
(MD2), right: Red Sea (MD8). a) δ18O curve from planktonic foraminifera, blue diamonds 315
indicate samples with bimodal grain-size distributions b) 87Sr/86Sr of detrital fraction, c) εNd
316
detrital fraction, d) % palygorskite in the clay fraction, e) % smectite in the clay fraction, f) 317
siliciclastic accumulation rate. The period of time studied includes MIS 1 and MIS 2. The 318
African Humid Period (AHP) (from Revel et al., 2015; Juginger and Trauth, 2013; Juginger et 319
al., 2014) is represented by a shaded block. H1 and H2 refer to the Heinrich events 1 and 2 320
respectively. Error bars in 87Sr/86Sr and εNd are within the points sizes. Sample AV22 (ash 321
layer) is not shown in the εNd curve so that the focus can be on the glacial-interglacial 322
variation. In both basins there is a change in Sr and Nd isotopes to a lower basaltic 323
composition during the Holocene, though this variation is less pronounced in the Red Sea.
324
Palygorskite percentage increases during the Holocene, whereas smectite and siliciclastic 325
accumulation rates diminish during this period in the GOA. In the Red Sea, palygorskite and 326
accumulation rates increase from 16 to 12 cal ka BP, while smectite decreases.
327 328
3.1.3. Mass accumulation rates 329
330
Mass accumulation rates can provide information on changes in sediment supply, reflecting 331
changes in source area, type of transport, and/or oceanic circulation (Maher et al, 2010). Mass 332
accumulation rates were calculated as:
333
Mass Accumulation Rate (g.cm-2.ka-1) = SR x DBD x f 334
where SR is the sedimentation rate (in cm.ka-1), DBD is the dry bulk density (in g.cm-3) and f 335
the percentage of the siliciclastic fraction, i.e. the fraction of the sediments comprised of 336
silicate minerals of detrital origin. The dry bulk density, mainly controlled by porosity in the 337
sediments, was estimated by measuring the dry weight of samples of known volume 338
(Bouilloux et al., 2013a, 2013b). The siliciclastic accumulation rate is shown in figure 4f for 339
the GOA and the Red Sea sediments. For core MD2 (GOA), accumulation rates are higher 340
(around 50 g.cm-2.ka-1) during the glacial period from 20 to 14 cal ka BP, then they diminish 341
from 14 to 12 cal ka BP during what is known as the Bølling Allerød (BA) period in the 342
higher latitudes, followed by a rise during the Younger Dryas (YD) event at 11 ka BP.
343
Accumulation rates then drop during the Holocene, to around 10 to 20 g.cm-2.ka-1. In core 344
MD8 (Red Sea), accumulation rates are low (around 15 g.cm-2.ka-1) and stable between the 345
glacial period and the Holocene, except for two peaks at 16-12 ka and 24-28 cal ka BP, 346
where accumulation rates increased to around 20 to 30 g.cm-2.ka-1. 347
348
3.2 Siliciclastic grain-size 349
350
Grain-size of siliciclastic particles from both cores range from 1 to 200 µm with significantly 351
coarser material in glacial period sediments (between 1 and 200 µm) than in Holocene 352
sediments (between 1 and 100 µm). The predominant size fraction in both basins is fine silt 353
(from 8 to 16 μm) (Fig. 5), and represents 60% to 80% of samples. This grain-size 354
distribution denotes a high grade of sediment sorting that is characteristic of wind-blown dust 355
(Chamley, 1989; Wang & Lai, 2014). The grain-size distributions are unimodal during the 356
Holocene (with a mode centered around 15 μm) and bimodal for some of the samples during 357
the glacial period, with a coarser size mode centered at about 50 μm. In both cores the coarser 358
mode is observed in the glacial period sediments and until around 14-15 cal ka BP (blue 359
diamonds in figure 4a). The finer mode is present all over the core and appears to increase 360
from the LGM to the Holocene in both basins. This finer mode is generally smaller for the 361
Red Sea sediments (8-22 μm) than for those of the GOA (10-40 μm) (Fig. 5).
362
XRD measurements of the fractions coarser and finer than 20 μm were performed for one 363
glacial period sample in each core in order to characterize the mineralogical composition of 364
both grain-size modes (grain-size subpopulations). The coarser size mode contains a larger 365
quantity of quartz and amphibole whereas the finer mode is dominated by phyllosilicates.
366
Pyroxene is only present in the coarser mode, possibly denoting a short-distance transport of 367
this mineral after physical weathering.
368
369
Figure 5: Grain-size distributions for the GOA and the Red Sea samples present similar 370
downcore behavior. Yellow line: Holocene (3 cal ka BP) in the GOA. Yellow dashed line:
371
Glacial period (15 cal ka BP) in the Gulf of Aden. Red line: Holocene (3 cal ka BP) in the 372
Red Sea. Red dashed line: Glacial period (15 cal ka BP) in the Red Sea. There is a change 373
from bimodal (mode 1 = 15 μm and mode 2 = 50 μm) to unimodal (10-40 μm) grain size 374
distributions from the glacial period to the Holocene in both basins.
375 376
3.3 Sr and Nd isotopic compositions 377
378
87Sr/86Sr ratios and εNd values measured on the carbonate-free fraction of cores MD2 and 379
MD8 are reported in Table 1. 87Sr/86Sr ratio values display glacial-interglacial variations for 380
both cores, with lower ratios in glacial sediments and higher values during the Holocene. In 381
core MD2 (GOA), the 87Sr/86Sr ratio gradually increases from 0.7070 between 18 and 14 cal 382
ka BP to 0.7085 during the Holocene. For core MD8 (Red Sea) the 87Sr/86Sr values vary from 383
0.7065 during glacial periods to 0.7075 during the Holocene, with an abrupt change at 12.8 384
cal ka BP (Fig 4b).
385
The εNd values are higher for the last glacial period sediments than for the Holocene sediments 386
in both cores. In core MD2 (GOA), the εNd values vary from -1.7 during glacial period to -3.6 387
during the Holocene (Fig 4c). The values become more negative after 14 cal ka BP and reach 388
minimum values (around -3.5) at 11.2 cal ka BP, remaining stable during the rest of the 389
Holocene. On this general trend, εNd values display a large positive peak reaching +3.8 from 390
17 to 15.6 cal ka BP. This anomalous εNd peak corresponds to a volcanic ash layer (Bouilloux 391
et al., 2013a) and probably originated from explosive volcanism in the Turkana basin and the 392
Hadar region in East Africa (Sirocko et al., 2000). Sr isotopes do not display this large 393
volcanic peak, possibly because Sr and Rb are more easily disturbed by weathering and 394
diagenetic processes (Goldstein and Hemming, 2003). Indeed, Sr is more easily lost from 395
minerals than Rb, whereas Nd and Sm show similar responses to weathering (Feng et al., 396
2009; van der Does et al., 2018). Primary minerals from the volcanic ash layer could have 397
been more prone to weathering (e.g. olivine) than the minerals from other sections of the 398
sediments, thus resulting in a decoupling of the 87Sr/86Sr and εNd data.
399
For core MD8 (Red Sea), εNd data show generally greater radiogenic values and a less 400
pronounced glacial-interglacial change than for the GOA sediments . The εNd values decrease 401
from -0.5 during the glacial period to -1.4 during the Holocene (Fig 4c). The εNd from this 402
core is also characterized by a small peak of -2.5 at 12.8 cal ka BP that is lower than the 403
Holocene values.
404 405
Table 1: εNd and 87Sr/86Sr data from the sediment cores MD2 and MD8. Samples marked with * 406
are chemical duplicates. Calibrated ages from Bouilloux et al. (2013a, 2013b) and Fersi et al.
407
(2016) 408
Site Core Lab number
Depth in core (cm)
Age (cal ka
BP)
143Nd/144Nd 2σ (106)
εNd 2σ 87Sr/86Sr 2σ (105)
Gulf of
MD92- 1002
AV17 66 0.99 0.512471 4 -3.31 0.07
AV18 115 1.90 0.512461 4 -3.50 0.08
Aden AV01 172 3.03 0.512471 3 -3.26 0.07 0.70857 2
AV13 257 5.12 0.512454 4 -3.59 0.08 0.70834
2
AV13* 257 5.12 0.512456 5 -3.54 0.10 0.70835
1
AV02 322 7.09 0.512458 6 -3.50 0.12 0.70813
2
AV14 356 8.30 0.512483 4 -3.03 0.08 0.70788
2
AV15 393 9.63 0.512474 3 -3.19 0.06 0.70788
1
AV03 434 10.45 0.512477 4 -3.13 0.07 0.70795
1
AV04 457 10.72 0.512458 4 -3.52 0.08 0.70781 2
AV05 497 11.18 0.512459 5 -3.49 0.10 0.70786
2
AV06 550 11.96 0.512482 4 -3.04 0.07 0.70763 1
AV07 602 12.76 0.512525 4 -2.20 0.07 0.70721 2
AV08 645 13.67 0.512528 3 -2.15 0.06 0.70712
1
AV09 675 14.46 0.512553 3 -1.66 0.06 0.70695
1
AV10 725 15.41 0.512545 4 -1.82 0.07 0.70707
2
AV19 735 15.60 0.512554 5 -1.68 0.10
AV20 745 15.78 0.512622 4 -0.35 0.08
AV16 765 16.15 0.512708 3 1.37 0.06 0.70692
2
AV16* 765 16.15 0.512702 5 1.24 0.10
AV21 785 16.53 0.512737 5 1.90 0.09
AV22 802 16.84 0.512834 5 3.78 0.09
AV11 815 17.09 0.512550 4 -1.72 0.08 0.70710
1
AV12 864 18.00 0.512543 4 -1.85 0.09 0.70701
1
AV12* 864 18.00 0.70701
1 Red
Sea
MD92- 1008
AR18 25 0.72 0.512564 4 -1.44 0.08
AR19 66 1.62 0.512567 4 -1.39 0.08
AR20 109 2.66 0.512565 4 -1.42 0.09
AR01 162 4.04 0.512564 3 -1.44 0.07 0.70748 1
AR14 203 5.52 0.512562 4 -1.47 0.08 0.70740
1
AR13 253 7.38 0.512576 5 -1.20 0.09 0.70723 2
AR15 292.5 9.00 0.512567 3 -1.38 0.07 0.70709 1
AR16 360 11.13 0.512559 4 -1.54 0.08 0.70719 1
AR21 377 11.59 0.512570 4 -1.33 0.08
AR02 420 12.79 0.512509 4 -2.51 0.08 0.70749 2
AR03 470 14.12 0.512560 4 -1.52 0.08 0.70674
1
AR24 477 14.29 0.512568 4 -1.37 0.08
AR25 498 14.82 0.512587 5 -1.00 0.09
AR04 530 15.63 0.512606 3 -0.62 0.07 0.70659 2
AR05 537 15.91 0.512631 3 -0.13 0.07 0.70657
1
AR26 538 15.96 0.512619 4 -0.37 0.08
AR07 595 18.66 0.512589 4 -0.96 0.08 0.70655 1
AR08 621 19.89 0.512595 4 -0.85 0.01 0.70646
1
AR09 650 23.38 0.512613 4 -0.48 0.07 0.70651
1
AR27 674 25.25 0.512581 4 -1.11 0.08
AR17 733 27.47 0.512584 3 -1.05 0.07 0.70669
1
4. Discussion 409
410
4.1 Identification of the detrital sediments sources 411
412
The isotopic and mineralogical data suggests that sediments from the Red Sea and the GOA 413
are originated from a mixture of basaltic and granitoid sources. Potential sedimentary sources 414
in this region are summarized in figure 6. The εNd vs. 87Sr/86Sr plots are used to determine 415
the specific Nd and Sr fingerprint of sediments and distinguish the contribution of different 416
lithologies. One of the main lithological units surrounding the Red Sea and GOA basins are 417
the Cenozoic volcanic rocks from the Afar region. Most of these terrains are covered by a 418
thick succession of flood basalts, which are rich in plagioclase and clinopyroxene (Kieffer et 419
al., 2004; Hagos et al, 2016). The triple junction causes the Afar depression, which covers an 420
area of 200,000 km2 and is flanked by the Ethiopian Plateau and Highlands (Bosworth et al., 421
2005). These territories display very radiogenic εNd values (0 to +10) and low 87Sr/86Sr ratios 422
(0.703 to 0.705) (Betton and Civetta, 1984; Teklay et al., 2009; Ayalew et al., 2019). From 423
figure 6, we can see that these rocks are likely to be the basaltic end-member of the sediment 424
source mixture. Sea-floor hydrothermal sources in the Red Sea and the GOA have already 425
been used to explain the 87Sr/86Sr values of sediments in the basins (Stein et al. 2007).
426
However, a hydrothermal origin for the sediments seems unlikely since the values for the Red 427
Sea Deeps (Pierret et al., 2009) have higher 87Sr/86Sr than our samples. This is also supported 428
by the fact that Nd and Sr isotopic values from cores KL23 (Palchan et al., 2013) in the North 429
of the Red Sea and KL15 (Stein et al., 2007) in the eastern GOA present 87Sr/86Sr values that 430
are more radiogenic than our samples, thus indicating that the basaltic character diminishes as 431
we move farther away from the Afar region.
432
Other important geological features in the region are the Saharan granitoids and the Arabo- 433
Nubian Shield (ANS) granitoids. The Mid- to Late Proterozoic Saharan granitoids present 434
very negative εNd values (-15 to -5) and high 87Sr/86Sr values (0.71 to 0.73) (Kuster et al., 435
2008). The ANS granitoids are Late Proterozoic in age and their isotopic values are 436
intermediary between the Sahara and the Afar region (εNd = 0 to -2.5; 87Sr/86Sr = 0.708 to 437
0.73) (Stoeser and Frost, 2006; Stein and Goldstein, 1996). Both of these regions could 438
correspond to the second end-member of the sediment mixture.
439
Most of the isotopic studies carried out in cores from the Red Sea and the GOA basins, 440
attributed the composition of sediments to a mixture between the volcanic rocks from the 441
Ethiopian Highlands and a granitic component that is either a combination of ANS and 442
Saharan sediments or only the ANS. Stein et al. (2007) awarded the most unradiogenic 443
87Sr/86Sr values in their samples from the GOA to a granitic component from the ANS 444
(≈0.714). They also found a loess component that yields a 87Sr/86Sr value of ≈0.7085, which 445
influenced both the Red Sea and the GOA. Palchan et al. (2013) found the sediments from the 446
northern Red Sea to be a mixture of granitoids from the Sahara masif, granitoids from the 447
ANS (in the Arabian Peninsula and eastern Africa) and the Cenozoic basalts from the Afar 448
region. These samples present values of around 0.710-0.711 87Sr/86Sr and from -2 to -8 εNd
449
(KL23 core from Fig. 6). Core KL11 in the central Red Sea, from the same authors, presented 450
values more similar to our samples of around 0.706 for 87Sr/86Sr and around -2 to 0 εNd. The 451
composition of sediments from this core was attributed to a basaltic-granitic mixture 452
originated in the Ethiopian highlands. Even if core KL11 was not measured for the same time 453
period as our samples, the similarities between the values suggest resembling origins for the 454
sediments in the central and southern Red Sea (core MD8 from this study). Data from core 455
NIOP 905 in the Arabian Sea (Jung et al., 2004) show higher 87Sr/86Sr values (0.714-0.716) 456
and lower εNd values (-5 to -7). The Sr isotopic composition was explained by an increase of 457
weathering intensity during the Holocene, whereas Nd isotopic compositions were attributed 458
to a northeast African source, i.e. the ANS, which remained constant. Core MD2 (from this 459
study) in the GOA shows values that are slightly higher in 87Sr/86Sr and lower in εNd than the 460
Red Sea values, so they are more influenced by this “granitic source” in comparison with the 461
Southern Red Sea basin (Fig. 6). This could indicate that the second end-member corresponds 462
to the ANS granitoids, since this source is closer to the GOA basin, while the southern Red 463
Sea receives a greater contribution from the Afar basalts.
464
Scheuvens et al., (2013) recently published a compilation of compositional data available for 465
northern African dust. They compiled a set of mineralogical and geochemical compositions 466
from the main African source regions established by Prospero et al. (2002). Some of this data 467
is plotted on figure 6 for comparison with our sedimentary cores. Sediments from Libya and 468
Tunisia from Grousset and Biscaye (2005) yield εNd values between -10 and -15, with 469
87Sr/86Sr ratios around 0.715. Sediments from Niger and Chad display similar values of 470
around -10 to -13 for εNd and 0.715 87Sr/86Sr. This area includes the Bodélé depression, 471
which is one of the most active dust sources on Earth. Finally, sediments from Egypt from 472
Grousset et al. (1998) present more moderate isotopic values, with εNd values between -3 and 473
-11 and 87Sr/86Sr ratios between 0.716 and 0.718, which are more similar to our sample 474
values.
475
Clay mineralogy of the different cores and dust sources in this region can also provide 476
information to help understand the provenance of sediments from cores MD2 and MD8. From 477
the clay correlation plots (supplementary material 1) we notice that percentages of 478
palygorskite and illite are negatively correlated to smectite concentrations in our samples, 479
suggesting that they are originated from different sources. Jung et al. (2004) found that the 480
most prominent clay mineral in their samples from core NIOP 905 was palygorskite, with 481
lower concentrations of kaolinite and smectite. Palchan et al. (2013), on the other hand, found 482
lower percentages of palygorskite and a predominance of illite-smectite in their samples from 483
core KL23 in the northern Red Sea. They also did not find variations in palygorskite 484
percentages between the last glacial period and the Holocene, similar to palygorskite values 485
from core MD8 (from this study) in the southern Red Sea. Stein et al. (2007) did not report 486
the presence of palygorskite in core KL11 in the central Red Sea, but did find clay minerals 487
like illite-smectite, kaolinite and chlorite.
488
Clay mineralogical data from Sirocko and Lange (1991) suggest that the Arabian Peninsula 489
soils are rich in both smectite and palygorskite. From the clay composition compilation of 490
Scheuvens et al. (2013), we can assume that sources from Chad and the Bodélé depression are 491
probably not contributing to the sediment mixture, since no palygorskite is observed and the 492
illite/kaolinite ratios are very low (<0.5). Mineralogical composition from Libyan and 493
Egyptian sediments are more in agreement with our results, with the occasional presence of 494
palygorskite and illite/kaolinite ratios between 1 and 3.
495
From all of this evidence we can establish that sediments from cores MD2 (GOA) and MD8 496
(Red Sea) result from a mixture of Afar volcanic rocks (end-member 1) and a granitic source 497
that is most similar to the ANS in Arabian and East African sediments (end-member 2).
498
However, we cannot completely rule out a contribution from the Saharan sediments based on 499
this isotopic and mineralogical data.
500 501
502
Figure 6: Nd-Sr plot with values from GOA and Red Sea sediments and possible source 503
regions. Samples from cores MD2 and MD8 are a mixture of the Afar region and Arabian- 504
Nubian Shield (ANS) sources. Cores KL11 and KL23 (Palchan et al., 2013) are located in 505
central and northern Red Sea, respectively. Core NIOP 905 (Jung et al., 2004) is located in the 506
Arabian Sea, off the Ethiopian coast.
507 508
4.2 Temporal evolution of the sedimentary tracers 509
510
A positive correlation (r=0.78) exists between εNd values and the percentages of smectite in 511
the clay fraction (supplementary material 2). An increase of smectite percentage is associated 512
with more positive εNd values. In the GOA, εNd values diminish from the glacial period to the 513
Holocene in the same way as smectite content decreases. In the Red Sea the smectite 514
percentages do not vary between the glacial period and the Holocene but there is a drop of 515
these values between 16 and 12 cal ka BP, which coincides with the εNd diminution to -2.5.
516
Smectite is a major product of volcanic rocks weathering (Singer, 1984) and its presence in 517
MD2 and MD8 cores is in agreement with a basaltic origin i.e. from the erosion of soils that 518
formed from the alteration of basalts in the Afar region (Yemane et al., 1987; Van Den 519
Eeckhaut et al., 2009). Smectite could also originate from in-situ weathering of basalts on the 520
mid-ocean ridges. However the study by Sirocko & Lange (1991) found no evidence of this 521
process in sediments of the Arabian Sea and the GOA during the late Quaternary and related 522
the spatial distribution of smectite contents to a continental origin.
523
Lower smectite percentages in sediments from MD2 and MD8 cores are associated with 524
higher palygorskite proportions. Past studies of clay percentages and accumulation rates on 525
sediments throughout the Arabian Sea, attributed the presence of palygorskite to a central 526
Arabian provenance (Sirocko et al, 1991; Sirocko & Lange, 1991). This mineral is transported 527
principally by winds (Chamley, 1989) and is particularly abundant in the Arabian Peninsula 528
soils (Mackenzie, 1984). From the world distribution map of palygorskite in soils (Velde, 529
1995), this mineral is also found in northeast Africa and the Sahara. In the GOA, palygorskite 530
reaches 20% of the clay fraction and is most abundant during the Holocene (after 10 cal ka 531
BP). Illite also shows a small increase from glacial period to the Holocene, while smectite 532
decreases. In the Red Sea, the highest smectite proportions are consistent with the greater 533
basaltic composition of rocks surrounding the basin. The diminution of εNd (and increase of 534
87Sr/86Sr) from the glacial period to the Holocene is not recorded by the smectite changes, 535
suggesting that the positive εNd values during glacial period are produced by the presence of 536
primary ferromagnesian minerals like pyroxene and not by clay minerals (Fig 3).
537
The magnetic susceptibility of cores MD2 and MD8 was studied by Bouilloux et al. (2013a 538
and 2013b). Magnetic susceptibility carries an intrinsic signal that depends on concentration, 539
grain size and type of magnetic mineral; which can reflect subtle lithological variations 540
(Stoner et al., 1996). In the GOA, first order magnetic data variations were attributed to 541
changes in the redox conditions in the basins caused by variations in bottom-water ventilation.
542
The low susceptibility values were explained by magnetite reduction, supported by the 543
presence of framboidal pyrite and high Total Organic Carbon (TOC) values. If we only take 544
into account the values not influenced by the magnetite preservation episode, magnetic 545
susceptibility correlates with changes in the εNd curve, so that positive εNd values correspond 546
to higher magnetic susceptibility and these variations could be source-driven. In the Red Sea, 547
high magnetic susceptibility values during the glacial period were attributed to a large amount 548
of minerals like goethite and hematite due to increased aridity and consequently a higher 549
supply of aeolian particles (Bouilloux et al., 2013b). Fe and Ti oxides, which are the most 550
important minerals that produce a magnetic signal, are more abundant in mafic volcanic rocks 551
than in silicic rocks (Tauxe, 2010). In both cores, intervals with greater radiogenic εNd values 552
(and less radiogenic 87Sr/86Sr values) are associated with an increase in the smectite 553
percentage and magnetic susceptibility values, confirming the higher relative contribution of 554
the basaltic material to the studied sites during the glacial period.In spite of the differences 555
between the Red Sea and the GOA, we found an overall coherence between the mineralogical 556
and geochemical data from both cores. They both show 2 modes of grain-size distribution 557
during the glacial period and they also fall in the same trend in the εNd correlation plots vs.
558
87Sr/86Sr and smectite, demonstrating that they experienced similar sedimentary regimes 559
(similar sediment sources and transport vectors). Changes in mineralogical and geochemical 560
compositions during the glacial-interglacial transition denote a variation in the relative 561
proportion of both sources.
562 563 564
4.3 Transport and aridity: two key components that explain variations in detrital 565
supplies to the Red Sea and the Gulf of Aden since the LGM 566
567
In the GOA, enhanced siliciclastic accumulation rates during the glacial period and the YD 568
are in agreement with a dominant physical weathering and an increase in source area aridity 569
(Clemens and Prell, 1991; Sun et al., 2008; Maher et al, 2010; Singer, 1984). Clemens and 570
Prell (1990) correlated the peaks in dust flux in the Arabian Sea to maxima of global ice 571
volume and found the orbital periodicities of 100, 41 and 23 ka, typical of global climate 572
changes. High aridity results in decreased vegetation cover and, therefore, increased deflation 573
of lithogenic material. This is consistent with the humidity index calculated by Fersi et al.
574
(2016) from pollen assemblages in sediments from core MD2. The driest period was 575
identified from 20 to 13.5 cal ka BP when siliciclastic accumulation rates are higher. The 576
highest humidity index values (from 10 to 4 cal ka BP) coincide with the lowest accumulation 577
rates of around 10 g.cm-2.ka-1, representing the most humid conditions in the sedimentary 578
source area.
579
In the Red Sea, mass accumulation rates stay stable around 10 g.cm-2.ka-1, except for two 580
peaks of 30 g.cm-2.ka-1 between 14 and 16 ka and of 20 g.cm-2.ka-1 between 24 and 28 ka.
581
Effects of sea level rise as a reason for higher sediment supply was proposed by Bouilloux et 582
al. (2013b). However, this explanation seems unlikely since remobilization of aeolian material 583
previously deposited on the continental margins would yield isotopic values more similar to 584
the ones observed during the glacial period (i.e. positive εNd and low 87Sr/86Sr ratios).
585
Instead, these peaks could be related to the Heinrich events 1 and 2, when conditions were 586
extremely arid in the region (Palchan et al., 2013; Shahanan et al., 2006).
587
Sediments from the glacial period display a bimodal distribution with an additional mode 588
value at 50 μm. Particle size variations may reflect changes in sediment transport and/or 589
sedimentary sources. Several studies have used dust size as an indicator of wind strength 590
(Clemens and Prell, 1991; Stuut et al., 2014) or changes in transport pathways, with larger 591
particles indicating shorter transport distances (Lambert et al., 2008; Maher et al., 2010). In 592
