Radiogenic isotopes for deciphering terrigenous input provenance in the western Mediterranean

Chemical Geology 410 (2015) 237–250

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Radiogenic isotopes for deciphering terrigenous input provenance in the western Mediterranean M. Rodrigo-Gámiz a,⁎, F. Martínez-Ruiz a, M. Chiaradia b, F.J. Jiménez-Espejo a,c, D. Ariztegui b a b c

Instituto Andaluz de Ciencias de la Tierra (IACT), CSIC-Universidad de Granada, Granada, Spain Department of Earth Sciences, University of Geneva, Geneva, Switzerland Japan Agency for Marine-Earth Science and Technology, Yokosuka, Japan

a r t i c l e

i n f o

Article history: Received 15 January 2015 Received in revised form 4 June 2015 Accepted 5 June 2015 Available online 11 June 2015 Keywords: Marine sediments Westernmost Mediterranean Last Glacial Maximum Radiogenic isotopes Terrigenous provenance Transport mechanisms

a b s t r a c t Radiogenic isotopic signatures in marine sediments can be used to trace terrigenous source areas and transport mechanisms, which are in turn related to climate variability. To date, most of the published studies using this approach have been focused on eastern Mediterranean sediments. In contrast, we study here the terrigenous input provenance in the westernmost Mediterranean (Alboran Sea basin) by using radiogenic isotope proxies and Nd model ages in a marine record spanning the last 20 ka. Nd, Sr and Pb isotopes, obtained from carbonate-free samples from the b37 μm size fraction, were used to characterize terrigenous variations, including eolian input. Substantial shifts in Pb isotopic signatures throughout the studied time interval reveal a change from North African dominated sources during the glacial period to European dominated sources during the Holocene. Nd and Sr shifts likewise indicate two main short-term changes in sediment provenance, during the last Heinrich event and the early–middle Holocene transition (ca. 8.9 ka cal. BP). Nd model ages over 1.45 Ga also support a contribution of an older component in the terrigenous source, likely Archaean material from the present Senegal region, during both periods. Conversely, terrigenous material mainly shows a dominant provenance from present-day Morocco, Mali, Mauritania, Niger, and Algeria, mixed with material from southern Iberia and southern France. Source variations in the westernmost Mediterranean were mainly driven by fluctuations in wind intensity and fluvial discharges. These fluctuations seem to have been modulated by the African monsoon system further conditioned by the ITCZ migrations and the position of the North Atlantic anticyclone system. © 2015 Elsevier B.V. All rights reserved.

1. Introduction The composition of terrigenous constituents in sediments of marginal marine basins mainly derives from riverine and eolian inputs, thus reflecting climatic conditions over adjacent continental regions (e.g., Kolla et al., 1979; Jeandel et al., 2007). In the case of the Mediterranean region, its proximity to the Sahara desert makes this area a key location for tracking terrigenous provenance and variations in source areas. In the western Mediterranean, the main sources of terrigenous material are the eolian fraction from arid and semi-arid regions in North Africa, the suspended fluvial particles from the southwestern Europe and North African runoff, and the suspended particulate matter from the Atlantic and Mediterranean waters (Grousset et al., 1988; Bergametti et al., 1989a, 1989b; Loÿe-Pilot and Martin, 1996; Molinaroli, 1996; Prospero, 1996; Stumpf et al., 2011). The Western Sahara dry land in North Africa is recognized as a major eolian dust source (Prospero, 1999; Goudie and Middleton, 2001; ⁎ Corresponding author at: Avda. de Las Palmeras 4, 18100 Armilla, Granada, Spain. E-mail addresses: [email protected] (M. Rodrigo-Gámiz), [email protected] (F. Martínez-Ruiz), [email protected] (M. Chiaradia), [email protected] (F.J. Jiménez-Espejo), [email protected] (D. Ariztegui).

http://dx.doi.org/10.1016/j.chemgeo.2015.06.004 0009-2541/© 2015 Elsevier B.V. All rights reserved.

Stuut et al., 2009; Prospero and Mayol-Bracero, 2013). Numerical simulations conclude that North Africa is the largest single source of dust on Earth, with up to 8 × 1014 g/a of atmospheric mineral dust, providing 50–70% of total dust emissions (Goudie and Middleton, 2001; Laurent et al., 2008). Extrapolating this figure to the total western Mediterranean basin area (840,000 km2) yields an average total atmospheric flux of 10.9 ± 0.6 × 1012 g/a (Bergametti et al., 1989a, 1989b; Loÿe-Pilot et al., 1989). The most important source areas of Saharan dust are located in Western Sahara, Mauritania and Senegal, northern Mali, Atlas Mountains through Morocco, Algeria and Tunisia, central and eastern Libya, western Chad (Bodélé depression), southern Egypt, and northern Sudan (Molinaroli, 1996; Moreno et al., 2006; Laurent et al., 2008; Stuut et al., 2009; Scheuvens et al., 2013). Dust source areas over Africa have changed over the past depending on the boundaries of different wind systems, their intensities, and the palaeo-positioning and migrations of the Inter-Tropical Convergence Zone (ITCZ) in response to factors such as orbital-induced summer insolation or changes in cross-equatorial temperature gradients (e.g., Nicholson, 2009; McGee et al., 2014). Furthermore, at the scale of Quaternary climatic oscillations, eolian dust fluxes to the ocean may have been higher during stadial (Greenland Stadial, GS) than interstadial periods (Greenland Interstadial, GI: nomenclature based on INTIMATE group, Lowe et al., 2008), as the

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result of a general southward displacement of the ITCZ during boreal summer (e.g., Reader et al., 1999; Elenga et al., 2000). At present, the relative amount of dust input from the Sahara and Sahelian regions to the western Mediterranean is low and occurs in winter (Bergametti et al., 1989b; Dulac et al., 1992), while in the eastern Mediterranean it is favoured during stronger summer westerlies (Jilbert et al., 2010). The terrigenous particles deflated from the surface depend on several factors including wind speed, atmospheric instability, height of the source area, particle size, particle exposure, soil moisture, vegetative cover, and mineralogical composition (e.g., deMenocal, 1995, 2004; Moreno et al., 2006; Mulitza et al., 2008). The determination of terrigenous particle provenance, of both eolian or riverine input, in Mediterranean marine sediments has, in general, been based on geochemical and mineralogical data (e.g., Coude-Gaussen et al., 1987; Foucault and Mélières, 2000; Wehausen and Brumsack, 2000; Caquineau et al., 2002; Weldeab et al., 2002a, 2002b, 2003; Díaz-Hernández et al., 2011; Formenti et al., 2011) as well as on remote sensing methods, analysis of surface dust observations, back-trajectory analysis, and the use of mineral tracers (e.g., Goudie and Middleton, 2001; Laurent et al., 2008). Radiogenic isotopes (Sr, Nd, Pb) are likewise powerful tracers for identifying and characterizing source areas of terrigenous material, which may in turn give us additional information about the provenance and transport mechanisms (e.g., Grousset et al., 1988, 1992, 1998; Revel et al., 1996; Tütken et al., 2002; Grousset and Biscaye, 2005; Jullien et al., 2007; Cole et al., 2009; Box et al., 2011; Meyer et al., 2011; Stumpf et al., 2011; Scheuvens et al., 2013). Previous research in a West–East Mediterranean transect characterized the spatial distribution of detrital flux and terrigenous provenance in Late Pleistocene and Holocene sediments, obtaining as main radiogenic end members Saharan dust and Nile particulate matter (e.g., Krom et al., 1999a, 1999b; Weldeab et al., 2002a, 2002b, 2003; Revel et al., 2010; Box et al., 2011; Blanchet et al., 2013). In the westernmost Mediterranean area, however, less work has focused on the identification of terrigenous provenance and transport patterns. Thus, here we provide a novel and detailed characterization of the terrigenous material and eolian input provenance using radiogenic signatures in marine sediments from a sediment record in the Alboran Sea basin. Previous mineralogical, geochemical and sedimentological analyses from this record have been used to reconstruct the climate variability in terms of atmospheric and oceanic responses as well as to identify eolian input variations since the Last Glacial Maximum (LGM) (Rodrigo-Gámiz et al., 2011, 2014a,b). However, the terrigenous provenance had not been described yet. In this study, we use Sr, Nd, and Pb isotopes and Nd model ages (TDM) to constrain for the first time the geographic provenance of terrigenous material and eolian dust, and their transport mechanisms during the last 20 ka. 1.1. Radiogenic isotopes as proxies for unravelling chemical weathering or terrigenous source The variability in radiogenic isotope composition of rocks, which are ultimately the source of the particulate matter suspended and transported by winds or rivers, is essentially the result of chemical fractioning between radioactive parents and radiogenic daughters operated by large-scale geological and ageing processes (Frank, 2002). Consequently, rocks of different ages and provenance in terms of large timescale reservoirs, e.g. mantle vs. crust, have significantly and measurably different isotopic compositions. The isotope composition of the terrigenous material that is the weathering product of the source rocks may coincide or not with that of the source rock, as discussed below. Nd is a relatively immobile element during weathering and therefore the Nd isotope composition in weathering products is stable during changes from wet to dry periods (e.g., Nesbitt et al., 1980; Dickin, 1997; Braun et al., 1998; Frank, 2002). In contrast, the Sr isotope composition is associated with changes in weathering intensity (Nesbitt et al., 1980; Frank, 2002), and thus radiogenic Sr is

preferentially removed from the source region, leaving a residue with low Sr isotope ratios (Blum and Erel, 1997). During prolonged chemical weathering the radiogenic Sr fraction of a rock is extracted from the source faster than the non-radiogenic Sr fraction, due to the preferential breakdown of Rb-rich phases such as mica and K-feldspar (Nesbitt et al., 1980; Frank, 2002). Therefore the combination of Sr–Nd isotope variations could reflect two different signatures, i.e. changes in the source area, and intense chemical weathering periods involving enhanced rainfall and riverine runoff (Frank, 2002). The extent of chemical weathering is strongly related to prevailing environmental conditions in the source area, such as cover vegetation or aridity. Furthermore, while the Nd isotope signature is unaffected by grain size variations (cf. Goldstein et al., 1984; Grousset et al., 1992), Sr isotopes have shown a relationship with grain size, i.e. Sr isotope ratios increase with decreasing grain size, as a result of the high Rb/Sr in finer sediments (cf. Biscaye and Dasch, 1971; Wehausen and Brumsack, 2000). Additionally, the Pb isotope composition provides information on source provenances since natural Pb can be transported in the detrital material or in eolian dust for long distances (Grousset et al., 1994; Grousset and Biscaye, 2005; Kylander et al., 2005). However, anthropogenic Pb contamination can overprint the natural Pb isotope signal (Grousset et al., 1994; Kylander et al., 2005). A particular advantage of the Th–U–Pb system is that binary mixtures form straight linear arrays in Pb–Pb isotope space, and deviations from such arrays imply mixtures involving more than two components. Therefore, linear Pb isotope arrays are consistent with binary mixing and imply the existence of multiple and different contributions of Pb sources. Further information can be obtained from the Nd model ages, TDM, which provide an isotopic fingerprint of the crustal source, in terms of timing and processes of crust formation (Arndt and Goldstein, 1987). Since the investigated sediments are a mixture of terrigenous material derived from different source areas, i.e. North Africa and South Europe, we expect to obtain “mixed” Nd model ages, which lie between the model ages of the surrounding potential source materials (Fig. 1a). 1.2. Present-day climate over the westernmost Mediterranean Modern climate conditions over the western Mediterranean and northwestern Africa areas are governed by the Azores high-pressure system linked to the North Atlantic climate variability and by African monsoonal dynamics. Summertime climates are usually dry and hot due to the influence of the atmospheric subtropical high-pressure belt (Sumner et al., 2001). During winter the subtropical high shifts to the south, allowing mid-latitude storms to enter the region from the open Atlantic and bringing enhanced amounts of rainfall to the western Mediterranean. This humidity regime in the western Mediterranean is mainly modulated by the North Atlantic Oscillation (NAO) (Hurrell, 1995; Trigo et al., 2002). A high (positive) NAO index causes a more northerly position of the North Atlantic depression, stronger than usual westerlies and warm and wet winters over Northern Europe and dry and cold winters over southern Europe, the Mediterranean and northern Africa (Wanner et al., 2001; Trigo et al., 2002). Conversely, a low (negative) mode of the NAO leads to opposite conditions. In addition, local climatology of Northern Africa is seasonally modulated by the latitudinal shift of the ITCZ and hence the African monsoon front. The ITCZ directly controls the location of precipitation over North Africa; changes in its position affect how much rainfall occurs and therefore the degree of river runoff. The review by Nicholson (2009) suggests a complex ITCZ pattern, especially for the boreal summer situation. In winter, the equator-ward displacement of the ITCZ (10°N) (Fig. 1a) causes a southward shift of dry subtropical air masses and is associated with the development of strong easterly Saharan Air Layer (SAL) winds, with a northern branch (NSAL). Consequently, a dust plume, usually generated by low-pressure systems, is transported between 15° and 25°N along an E–W axis over the tropical Atlantic Ocean (Holz, 2004) and over the Mediterranean (Moulin et al., 1997).

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239

a) France 87

Switzerland

Rhône

TDM

Nd

Nd

Ebro

40ºN

Nd

293G

DM

t Morocco on M as Atl

s ain

Tunisia

87

204 204

208

204

208

206

204

207

204

208

204

86

Libya Algeria

207

206

Nd

Egypt

Nd

DM 206

40ºN

Languedoc TDM

Guadiana Spain Guadalquivir Guadalhorce Guadalfeo

30ºN

(0)

Têt

86

87

30ºN

86

TDM

Western Sahara

Nd 87

Algeria-Nigeria TDM

86

Nd

20ºN

87

Mali

Mauritania

86

20ºN

Niger

Nd 87

Nd

86

TDM

87

86

Chad

ly er-Ju

summ

Sudan

Senegal Guinea

10ºN

10ºN

ary

winter-Janu 20ºW

8

10

12

14 Aller d

YD

B lling

6

IACP

4

EMHT

2

7.4ka

0

8.2ka

b)

0º

10ºW

10ºE

16

18

20ºE

20 LGM

H1

26 24

Zr/Al

22 20 18 16 14 1a 1b

Holocene

0

2

4

6

GS-1

8

10

12

1c

1d 1e

14

2b

2a

GI-1

GS-2

16

18

20

Fig. 1. a) Map of the studied area showing the location of the marine sediment record 293G retrieved in the westernmost Mediterranean. Potential terrigenous source areas from North Africa, southern Iberia and southwestern Europe with radiogenic signatures and TDM compiled from the literature are included. Dashed red and blue lines indicate the summer (July) and winter (January) positions of the ITCZ, respectively; b) Zr/Al profile from Rodrigo-Gámiz et al. (2011) as proxy for eolian input was used to select 37 samples for radiogenic measurements covering the last 20 ka. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

During boreal summer, the dry subtropical air shifts northward and the ITCZ is located around 20°N (Fig. 1a), representing the onset of the rainy season (summer monsoon) in North Africa with heavy rainfall and changes in atmospheric circulation (Peyrillé and Lafore, 2007). Complex parameters and mechanisms actually modulate precipitation over West Africa during summer (Nicholson, 2009). For instance, a warming in the Mediterranean Sea basin could favour the northward expansion of the monsoon in summer (Rowell, 2003; Hall and Peyrillé, 2006).

2. Material and analytical techniques 2.1. Marine setting, sediment record description and sample selection The Alboran Sea, located in the westernmost Mediterranean Sea, is a semi-closed basin surrounded by the Iberian Peninsula in the North and the North Africa margin in the South. A gravity core (293G) was taken in the East Alboran basin (Fig. 1a) (36°10.414′N, 2°45.280′W, depth 1840 m, length 402 cm) during the oceanographic cruise Training

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Through Research-12 (R/V Professor Logachev) (Comas and Ivanov, 2003). The selected sediment record contains homogeneous green-brownish hemipelagic mud-clays with some nannofossils. High-resolution geochemical, mineralogical and sedimentological analyses at about 75 a/ 1.5 cm (n = 267 samples) from this sediment record have been reported in Rodrigo-Gámiz et al. (2011). Conventional X-Ray Fluorescence (XRF) was used to obtain the elemental geochemistry of major elements and Zirconium (Zr). A previously developed age model based on linear interpolation between ten calibrated AMS 14 C-dates (see Rodrigo-Gámiz et al., 2011, for a description) and two later dates (Rodrigo-Gámiz et al., 2014b) during critical intervals indicates that this marine core spans the last 20 ka with a mean sedimentation rate of ca. 20 cm/ka. A total of 37 samples of potential eolian material were selected covering the last 20 ka on the basis of eolian proxies such as Zr/Al ratio (Fig. 1b), quartz and palygorskite contents and Si/Al and Ti/Al ratios (Supplementary Fig. S1; Rodrigo-Gámiz et al., 2011). Because changes in grain size can have a substantial influence in Sr-isotope compositions used as provenance indicators and in order to avoid any hemipelagic contribution, we consistently worked with the same grain size fraction. Based on the fact that most of the potential Saharan dust particles are fine silts and clay-size minerals (Grousset et al., 1998; Grousset and Biscaye, 2005; Stuut et al., 2009) and the grain-size distribution of this material showed mean sizes below 6 μm (Fig. 3d) (Rodrigo-Gámiz et al., 2011), in this study we investigated the fine silt fraction, i.e. b 37 μm.

2.2. Sediment digestion procedure, element separation and isotope analysis Chemical extractions for Sr, Nd and Pb isotopes were carried out at the Department of Earth Sciences of the University of Geneva (Switzerland). Each sample was treated to leach the carbonate fraction with 1 N acetic acid solution followed by a triple rinse with water. The N37 μm fraction was removed by dry sieving. Around 100–150 mg of the dried alumino-silicate residual fraction was mineralized in pure acids using Teflon bombs heated on an electric hotplate (140 °C) in a two-step procedure: at first adding a mixture of 1 ml HNO3 15 M with 4 ml HF at 140 °C for 7 days, and secondly after evaporation adding 3 ml HNO3 15 M at 140 °C for 2 days. After evaporation, the samples were finally diluted in 1.9 ml HNO3 1 M for chemical separation. Sr, Nd and Pb separations from the prepared solutions were carried out using cascade columns with Sr-Spec, TRU-Spec and Ln-Spec resins following a modified method after Pin et al. (1994). Pb was further purified by anion exchange chromatography using an AG-MP1-M clean resin in hydrobromic medium and small volume columns (0.08 ml). Pb was loaded on Re filaments using the silica gel technique (Gerstenberger and Haase, 1997) and Pb isotope ratios of all samples and standards were measured in static mode on Faraday cups on a multicollector Thermo TRITON mass spectrometer at a pyrometer controlled temperature of 1220 °C. Pb isotope ratios were corrected for instrumental fractionation by a factor of 0.1% per amu based on more than 100 measurements of the SRM981 standard and using the standard values of Todt et al. (1996). Procedural blanks were b200 pg. External reproducibilities (2σ) of the standard ratios are 0.05% for 206 Pb/204Pb, 0.08% for 207Pb/204Pb, 0.10% for 208Pb/204Pb, 0.006% for 206 Pb/207Pb, 0.007% for 208Pb/207Pb and 0.008% for 208Pb/206Pb. Sr was loaded on single Re filaments with a Ta oxide solution and measured on the multicollector Thermo TRITON mass spectrometer at a pyrometer controlled temperature of 1480 °C in static mode using the virtual amplifier design to cancel out biases in gain calibration among amplifiers. 87Sr/ 86 Sr values were internally corrected for fractionation using a 88Sr/86Sr value of 8.375209. Raw values were further corrected for external fractionation by a value of +0.03‰, determined by repeated measurements of the SRM987 standard (87Sr/86Sr = 0.710250;

McArthur et al., 2001). External reproducibility (1σ) of the SRM987 standard is b 7 ppm. Nd was loaded on double Re filaments with 1 M HNO 3 and measured on a 7-collector Finnigan MAT 262 thermal ionization mass spectrometer with extended geometry and stigmatic focusing. 143 Nd/ 144 Nd ratios were measured in dynamic mode (quadruple collector) and internally corrected for fractionation using a 146Nd/144Nd value of 0.721903 and corrected for external fractionation using the JNdi1 standard ( 143 Nd/ 144 Nd = 0.512115 ± 7; Tanaka et al., 2000). Our mean of 19 replicated analyses of this standard in dynamic mode was 0.512097 ± 3 × 10 − 6 (2σ) during the period of analysis. The 143 Nd/ 144 Nd ratio is commonly expressed as ε Nd (0) = [(143 Nd/ 144 Nd) sample ∕ ( 143 Nd/ 144 Nd) CHUR − 1] × 10 4 with the present day Chondritic Uniform Reservoir (CHUR) being 0.512638 (Wasserburg et al., 1981). 2.3. Nd model ages Additionally, another portion (around 100–150 mg) of 35 out of 37 dried alumino-silicate residual sediments was used to calculate the distribution of Nd model ages relative to depleted mantle model ages, TDM , according to DePaolo (1981) (147 Sm/ 144 Nd = 0.21378 and 143Nd/144Nd = 0.513155). The alumino-silicate residual samples were acid leached using a digestion of HNO3 + HF in PFA-lined pressure vessels in a microwave field following the method by Montero and Bea (1998). The 147Sm/144Nd ratio was determined by inductively coupled-plasma mass spectrometry (ICP-MS) in the Centre for Scientific Instrumentation (CIC) at the University of Granada (Spain). The ICP-MS was a quadrupole-based Perkin-Elmer Sciex ELAN-5000a equipped with a conventional cross-flow nebulizer and spray-chamber. Measurements were taken in triplicate with a precision better than 1.2% (2σ) and using Rb, Sr, Sm, Nd and Rh standard solutions of various concentrations as internal standards. 3. Results 3.1. Sr, Nd and Pb isotope composition The Sr and Nd isotope values, Sm and Nd concentrations, 147Sm/144Nd ratio and TDM of the carbonate-free b37 μm sediment fraction are shown in Table 1, and the Pb isotope values are presented in Table 2. Nd model ages in the marine sediment yield a nearly Gaussian distribution with the most frequent TDM value centred at 1.4 Ga with a SD of 0.5 (Fig. 2). The temporal evolution of Nd model ages for the last 20 ka in the westernmost Mediterranean shows the geological age variations of the source regions (Fig. 3a). Maximum TDM values of 1.48 and 1.60 Ga are reached at 16.2 and 10.2 ka cal. BP respectively, followed by 1.46 Ga recorded at 8.9 and 8.7 ka cal. BP. Minimum TDM values ca. 1.34–1.36 Ga are observed at 14.3, 14.1, 13.3, 11.1, 9.4 and 3.4 ka cal. BP (Fig. 3a). The Nd isotope signature ranges between εNd(0) −11.7 and −10.5 (Fig. 3b) with more negative εNd(0) values, i.e. less radiogenic 143 Nd/144Nd ratio, at 17.6, 16.2–16.0, 14.6, 8.9–8.7, 5.6 and 2.2 ka cal. BP (Fig. 3b). Some of these shifts in the εNd(0) signature are also evidenced in the 87Sr/86Sr ratio, showing several periods with less radiogenic values recorded at 18.7, 17.1–16.7, 14.3–14.1, 13.5, 13.1, 10.2, 8.9, 5.6 and 1.1 ka cal. BP (Fig. 3c). The Pb isotope composition, in particular 206Pb/204Pb, 207Pb/204Pb, and 208Pb/204Pb ratios, shows almost the same variability pattern, ranging from 18.8275 to 18.6929, 15.7547 to 15.5947, and 39.0807 to 38.6815, respectively (Fig. 4a–c). Higher Pb isotope ratio values during the last 20 ka are observed at 17.6, 13.5, 11.1 and 4.3–3.4 ka cal. BP (Fig. 4a–c). Conversely, lower values of Pb isotope ratios are recorded at 18.7, 16.7, 12.8, 11.5, 8.9 and 2.2 ka cal. BP (Fig. 4a–c). These latest periods are also represented by less negative values of εNd(0), i.e. more radiogenic 143 Nd/144Nd ratio, and lower values of 87Sr/86Sr (Fig. 3b–c). Finally, the

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Table 1 Sr concentrations, 87Sr/86Sr and 143Nd/144Nd ratios, εNd(0) values, Sm and Nd concentrations, 147Sm/144Nd ratio and TDM on the carbonate-free fraction below 37 μm from the marine core 293G. 147Sm/144Nd = 0.21378 and 143Nd/144Nd = 0.513155 value ratios used to calculate TDM according to DePaolo (1981). Samples

Depth (cm)

Age (ka cal. BP)

Sr (ppm)

87

1σ ∗ 10−6

143

293G 1 16.5–18 293G 1 34.5–36 293G 1 55.5–57 293G 2 12–13.5 293G 2 33–34.5 293G 2 54–55.5 293G 3 16.5–18 293G 3 28.5–30 293G 3 33–34.5 293G 3 42–43.5 293G 4 6–7.5 293G 4 28.5–30 293G 4 42–43.5 293G 5 0–1.5 293G 5 12–13.5 293G 5 22.5–24 293G 5 30–31.5 293G 5 39–40.5 293G 5 46.5–48 293G 5 54–55.5 293G 6 1.5–3 293G 6 7.5–9 293G 6 12–13.5 293G 6 19.5–21 293G 6 28.5–30 293G 6 34.5–36 293G 6 39–40.5 293G 6 48–49.5 293G 6 51–52.5 293G 6 52.5–54 293G 7 1.5–3 293G 7 7.5–9 293G 7 15–16.5 293G 7 18–19.5 293G 7 24–25.5 293G 7 33–34.5 293G 7 40.5–42

17.25 36.75 56.25 70.75 91.75 112.75 133.75 145.75 150.25 159.25 181.25 203.75 217.25 232.75 244.75 255.25 262.75 271.75 279.25 286.75 292.25 298.25 302.75 310.25 319.25 325.25 329.75 338.75 341.75 343.25 351.25 357.25 364.75 367.75 373.75 382.75 390.25

1.1 2.2 3.4 4.3 5.6 6.8 8.0 8.7 8.9 9.4 10.2 10.8 11.1 11.5 12.0 12.7 12.8 13.1 13.3 13.5 13.6 13.8 13.9 14.1 14.3 14.6 15.2 16.0 16.1 16.2 16.7 17.1 17.6 17.8 18.2 18.7 19.2

449.325 565.668 610.58 557.265 538.196 581.792 628.359 652.7198 642.647 651.014 645.088 581.759 565.881 553.416 556.243 582.626 596.618 581.398 561.848 588.682 579.871 543.292 526.542 503.957 445.874 429.905 403.880 385.348 425.231 456.444 447.815 423.367 375.639 414.314 539.771 641.713 603.560

0.716771 0.717964 0.717332 0.717776 0.716461 0.717616 0.717816 0.717795 0.714397 0.717152 0.716134 0.717591 0.717651 0.717110 0.717639 0.716945 0.716280 0.715134 0.717187 0.716166 0.717347 0.717056 0.716427 0.715899 0.716152 0.717414 0.717045 0.716053 0.716376 0.715867 0.713110 0.715785 0.717949 0.717095 0.718042 0.714939 0.717840

2 2 2 2 2 2 5 2 2 3 1 2 2 3 2 3 2 3 3 1 3 3 2 2 2 2 4 2 3 2 3 2 2 3 2 2 2

0.512088 0.512068 0.512081 0.512096 0.512066 0.512087 0.512077 0.512037 0.512041 0.512099 0.512095 0.512071 0.512078 0.512059 0.512057 0.512068 0.512062 0.512068 0.512079 0.512064 0.512052 0.512064 0.512053 0.512062 0.512076 0.512060 0.512080 0.512047 0.512062 0.512044 0.512067 0.512057 0.512041 0.512060 0.512057 0.512074 0.512060

Sr/86Sr

207

Pb/206Pb values range between 0.83788 and 0.83286, the Pb/206Pb ratio shows variations between 2.07794 and 2.06768, and the 208Pb/207Pb ratio between 2.48654 and 2.47854 (Fig. 4d–f), displaying more radiogenic signatures at 18.7, 16.0, 13.5, 11.1, 9.4 and 4.3–2.2 ka cal. BP. 208

4. Discussion 4.1. Radiogenic isotope variations in the westernmost Mediterranean during the last 20 ka Radiogenic isotopes have provided an additional tool for reconstructing palaeoenvironmental conditions in the westernmost Mediterranean. For the last 20 ka, the 87 Sr/ 86 Sr ratio showed low values during the LGM, at 18.7 ka cal. BP, and the last Heinrich event (H1), ca. 16.7–16.2 ka cal. BP (Fig. 3c). This could suggest the release of radiogenic Sr during chemical weathering, leaving a low radiogenic signature in the source rock (Blum and Erel, 1997). However, during the Greenland Stadial 2 (GS-2, from 20 to 14.7 ka cal. BP) generally arid conditions and semi-desert vegetation cover prevailed over the continental areas surrounding the western Mediterranean basin (e.g., Fletcher and Sánchez Goñi, 2008; Combourieu Nebout et al., 2009), and no prolonged weathering processes have been registered. In addition, contemporaneous Nd and Pb isotopic fluctuations (Figs. 3b, 4) suggest that the Sr isotope signature indeed reflects variations in the terrigenous provenance rather than a weathering signal. The Sr isotope signature and εNd(0) values for these period are comparable with the signal of Sahara material (Jeandel et al., 2007; Abouchami et al., 2013). Although TDM model ages do not show major

Nd/144Nd

1σ ∗ 10−6

εNd(0)

Sm (ppm)

Nd (ppm)

147

2 3 2 2 4 8 2 8 6 16 23 14 2 7 6 1 7 1 7 4 18 1 6 7 2 25 2 5 1 3 1 8 9 2 26 3 14

−10.7 −11.1 −10.9 −10.6 −11.2 −10.7 −10.9 −11.7 −11.6 −10.5 −10.6 −11.1 −10.9 −11.3 −11.3 −11.1 −11.2 −11.1 −10.9 −11.2 −11.4 −11.2 −11.4 −11.2 −11.0 −11.3 −10.9 −11.5 −11.2 −11.6 −11.1 −11.3 −11.6 −11.3 −11.3 −11.0 −11.3

4.99 3.92 4.85 4.71 4.77 4.60 4.97 4.66 4.27 4.41 3.92 3.07 3.90 4.39 4.26 3.91 4.25 4.22 2.79 4.03 4.14 – 4.42 4.03 4.39 4.04 3.73 − 4.49 4.33 4.07 4.44 3.82 4.02 2.71 3.19 4.13

27.82 22.26 27.7 26.46 26.66 26.06 27.39 26.10 23.96 25.04 21.48 15.43 22.65 25.35 24.16 22.65 23.57 23.59 16.38 22.56 23.60 – 25 23.68 25.81 22.65 20.82 − 24.92 23.79 23.73 25.11 22.19 23.29 15.40 18.06 24

0.108 0.106 0.104 0.108 0.108 0.107 0.110 0.108 0.108 0.106 0.110 0.120 0.104 0.105 0.107 0.104 0.109 0.108 0.103 0.108 0.106 – 0.107 0.103 0.103 0.108 0.108 − 0.109 0.110 0.104 0.107 0.104 0.104 0.106 0.107 0.104

Sm/144Nd

TDM (Ga) 1.38 1.38 1.34 1.38 1.43 1.38 1.43 1.46 1.46 1.36 1.41 1.60 1.36 1.39 1.42 1.37 1.44 1.44 1.34 1.43 1.42 – 1.43 1.36 1.34 1.43 1.41 − 1.44 1.48 1.37 1.42 1.41 1.38 1.42 1.40 1.38

shifts until 16.2 ka cal. BP (with a small peak value of ca. 1.48 Ga; Fig. 3a), the source variation in the terrigenous material is further supported by the increased values of eolian proxies such as Zr/Al, Si/Al, and Ti/Al ratios, major quartz content and the presence of palygorskite (Supplementary Fig. S1; Rodrigo-Gámiz et al., 2011), which is in agreement with cold and arid conditions during both time periods (Bout-Roumazeilles et al., 2007; Jullien et al., 2007). Model results and palaeoclimate records have documented southward shifts of the ITCZ during boreal summer with an increase in eolian dust input during these cold periods (e.g. Steager et al., 2011; Arbuszewski et al., 2013; McGee et al., 2013, 2014). Thereafter, the Bølling–Allerød (B–A) period (from ca. 14.7 to 12.9 ka cal. BP) showed rapid shifts to likely weathered conditions with preferential leaching of radiogenic Sr at 14.3–14.1, 13.6, and 13.1 ka cal. BP (Fig. 3c). Assuming the onset of the African Humid Period (AHP) in North Africa around 14.5 ka cal. BP (deMenocal et al., 2000), the progressive increase of tropical vegetation and major humid conditions over northern African and southern European areas would have favoured major weathering and the release of radiogenic Sr from the terrigenous source. Typical εNd(0) values from fluvial sediments derived by the Rhône River in the western Mediterranean range between −10.8 and − 9.7 (Henry et al., 1994), whereas the typical average Saharan– Sahelian isotopic εNd(0) signature ranges from − 14 to − 11 (see synthesis in Scheuvens et al., 2013), and the average European atmospheric input has εNd(0) values between −12.2 and −10.2 (Henry et al., 1994). The εNd(0) values at 14.3 and 13.3 ka cal. BP (−11.0 and −10.9; Fig. 3b) are close to both the riverine input signature and the European atmospheric input (Henry et al., 1994), suggesting that the B–A period was characterized by a mixture of terrigenous material.

242

Table 2 Pb isotopic data measured on the carbonate-free fraction below 37 μm from the marine core 293G. 206Pb/204Pb, 207Pb/204Pb, 208Pb/204Pb, 207Pb/206Pb, 208Pb/206Pb, and 208Pb/207Pb ratios and their respective errors. Depth (cm)

Age (ka cal. BP)

206

Pb/204Pb

293G 1 16.5–18 293G 1 34.5–36 293G 1 55.5–57 293G 2 12–13.5 293G 2 33–34.5 293G 2 54–55.5 293G 3 16.5–18 293G 3 28.5–30 293G 3 33–34.5 293G 3 42–43.5 293G 4 6–7.5 293G 4 28.5–30 293G 4 42–43.5 293G 5 0–1.5 293G 5 12–13.5 293G 5 22.5–24 293G 5 30–31.5 293G 5 39–40.5 293G 5 46.5–48 293G 5 54–55.5 293G 6 1.5–3 293G 6 7.5–9 293G 6 12–13.5 293G 6 19.5–21 293G 6 28.5–30 293G 6 34.5–36 293G 6 39–40.5 293G 6 48–49.5 293G 6 51–52.5 293G 6 52.5–54 293G 7 1.5–3 293G 7 7.5–9 293G 7 15–16.5 293G 7 18–19.5 293G 7 24–25.5 293G 7 33–34.5 293G 7 40.5–42

17.25 36.75 56.25 70.75 91.75 112.75 133.75 145.75 150.25 159.25 181.25 203.75 217.25 232.75 244.75 255.25 262.75 271.75 279.25 286.75 292.25 298.25 302.75 310.25 319.25 325.25 329.75 338.75 341.75 343.25 351.25 357.25 364.75 367.75 373.75 382.75 390.25

1.1 2.2 3.4 4.3 5.6 6.8 8.0 8.7 8.9 9.4 10.2 10.8 11.1 11.5 12.0 12.7 12.8 13.1 13.3 13.5 13.6 13.8 13.9 14.1 14.3 14.6 15.2 16.0 16.1 16.2 16.7 17.1 17.6 17.8 18.2 18.7 19.2

18.7330 18.6947 18.8039 18.8030 18.7776 18.7737 – 18.7632 18.7429 18.7564 18.7703 18.7879 18.8230 18.7524 18.7799 18.7570 18.7095 – 18.7358 18.7814 18.7699 18.7990 18.7685 18.7699 – 18.7525 18.7880 18.7803 18.7570 18.7557 18.6920 18.7852 18.8275 18.8065 18.7850 18.7384 18.7875

1σ ∗ 10−4

207

– 1 74 176 4 104 – 3 1 71 3 18 124 20 3 69 117 – 2 34 3 21 2 52 – 3.1 19 20 41 2 151 1.7 79 13 77 104 2.1

15.6880 15.6640 15.7231 15.7128 15.6694 15.6795 – 15.6685 15.6673 15.6838 15.6828 15.6778 15.7547 15.6492 15.6753 15.6632 15.6120 – 15.6675 15.7232 15.6681 15.6878 15.6666 15.6763 – 15.6703 15.6774 15.6723 15.6891 15.6603 15.5947 15.6557 15.6872 15.6631 15.6610 15.6411 15.6618

Pb/204Pb

1σ ∗ 10−4

208

– 2 62 147 5 87 – 3 1 58 3 15 107 16 3 57 94 – 2 31 3 19 2 43 – 3.1 16 17 32 2 124 1.8 66 10 65 89 2.4

38.9090 38.8239 39.0572 39.0695 38.8879 38.9016 – 38.8763 38.8649 38.9278 38.9293 38.9328 39.0807 38.8343 38.9149 38.8914 38.7490 – 38.8474 39.0222 38.8735 38.9378 38.8717 38.9050 – 38.8706 38.8933 38.8763 38.9069 38.8435 38.6815 38.8553 38.9532 38.8853 38.8742 38.7936 38.8583

Pb/204Pb

1σ ∗ 10−4

207

– 5 158 356 14 214 – 10 3 144 7 37 264 41 10 143 237 – 4 83 10 47 5 105 – 9.3 39 39 82 5 308 5.8 164 27 161 218 7.7

0.83754 0.83788 0.83607 0.83572 0.83447 0.83512 – 0.83507 0.83591 0.83620 0.83551 0.83444 0.83701 0.83457 0.83468 0.83506 0.83442 – 0.83623 0.83714 0.83475 0.83454 0.83473 0.83535 – 0.83564 0.83444 0.83448 0.83646 0.83496 0.83433 0.83340 0.83323 0.83286 0.83375 0.83464 0.83363

Pb/206Pb

1σ ∗ 10−5

208

– 0 4 8 1 4 – 1 0 3 0 1 6 1 1 3 4 – 0 3 1 1 0 3 – 0 1 1 2 0 6 0 4 1 3 4 0

2.07703 2.07673 2.07706 2.07794 2.07094 2.07190 – 2.07193 2.07359 2.07553 2.07395 2.07216 2.07673 2.07092 2.07214 2.07334 2.07104 – 2.07342 2.07786 2.07105 2.07161 2.07109 2.07293 – 2.07285 2.07012 2.06999 2.07435 2.07102 2.06950 2.06839 2.06894 2.06768 2.06953 2.07029 2.06830

Pb/206Pb

1σ ∗ 10−5

208

– 1 13 14 3 10 – 3 1 6 1 3 22 2 3 6 7 – 1 12 3 7 1 4 – 2 2 3 8 1 10 1 6 2 6 7 2

2.48018 2.47854 2.48395 2.48654 2.48174 2.48079 – 2.48117 2.48064 2.48211 2.48228 2.48328 2.48085 2.48140 2.48255 2.48275 2.48191 – 2.47947 2.48212 2.48105 2.48229 2.48116 2.48151 – 2.48054 2.48080 2.48060 2.47984 2.48039 2.48048 2.48187 2.48311 2.48261 2.48228 2.48050 2.48108

Pb/207Pb

1σ ∗ 10−5 – 1 10 18 2 14 – 2 1 9 1 4 14 3 2 7 10 – 1 8 2 5 1 6 – 1 3 3 7 1 15 1 9 2 8 9 1

M. Rodrigo-Gámiz et al. / Chemical Geology 410 (2015) 237–250

Samples

M. Rodrigo-Gámiz et al. / Chemical Geology 410 (2015) 237–250

14

16 B lling

Aller d

YD

18

20 LGM

H1

1.70 1.65 1.60 1.55 1.50 1.45 1.40 1.35 1.30

Nd

(0)

a)

b)

-12.5 -12.0 -11.5 -11.0 -10.5 -10.0 -9.5

c)

6

0.719 0.718 0.717 0.716 0.715 0.714 0.713 0.712

d)

TDM (Ga)

12

86

10

IACP

8 EMHT

6 7.4ka

4

8.2ka

2

0

The Younger Dryas (YD) cold period (from ca. 12.9–11.7 ka cal. BP) has been associated with a sharp increase in regional aridity that interrupted the AHP in subtropical African records (e.g., deMenocal et al., 2000; Gasse, 2000). In contrast, the Sr, Nd, and Pb isotope compositions present almost constant values (Figs. 3b–c, 4). Previous geochemical and mineralogical data from this record (see Supplementary Fig. S1 and Rodrigo-Gámiz et al., 2011) showed a palaeoenvironmental scenario involving a combination of two phases, a first dry phase followed by a second with more humid conditions over poorly vegetated continental areas. Palynological records from the western Mediterranean region also highlight an early period with very dry conditions and an increase in semi-desert taxa, followed by a more humid period with a slight increase in forest vegetation (e.g., Combourieu Nebout et al., 2009). Similar dry–humid conditions have been reported for this period based on marine records from offshore NW Africa, suggesting a shift of the northern limit of the African rain belt and associated wind systems (Meyer et al., 2011). However, these two phases are not reflected in the 87 Sr/86Sr ratio, probably because it was dominated by sporadic runoffs without sufficient time to have a prolonged weathering effect. Therefore, the absence of substantial differences in the isotope composition and

87

Fig. 2. Distribution of Nd model ages, TDM, in marine core 293G, calculated according to DePaolo (1981).

243

5 4 3 2 1 0 1a 1b

Holocene

0

2

4

6

GS-1

8

10

12

1c

1d 1e

14

2b

2a

GI-1

GS-2

16

18

20

Fig. 3. 87Sr/86Sr ratio (a), εNd(0) (b) and TDM (c) plotted against age. Error bars are smaller than symbol size if they are not shown. Data from African (Scheuvens et al., 2013) and European sources (Henry et al., 1994) are showed. Grey vertical bars show periods with less radiogenic 87Sr/86Sr values. Upper panel: light red vertical bars indicate the Allerød (A), and Bølling (B) warm time intervals. Light blue vertical bars indicate main cold periods, such as the last Heinrich event (H1), the Older Dryas and Younger Dryas (YD) time intervals, and the 8.2 ka cold event. Short dashed red vertical bars show the early–middle Holocene transition (EMHT) at ca. 8.5–8.4 ka cal. BP and the demise of the African Humid Period (AHP) at 7.4 ka cal. BP. Lower panel: Greenland Stadials (GS-1, GS-2) and Interstadials (GI-1) and Holocene time intervals according to the event stratigraphy timing proposed by the INTIMATE group (Lowe et al., 2008). Black squares indicate twelve 14C-AMS dates with two-sigma probability interval using Calib 6.0.2 software (Stuiver and Reimer, 1993) and the Marine09 calibration curve (Reimer et al., 2009). (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

M. Rodrigo-Gámiz et al. / Chemical Geology 410 (2015) 237–250

comparable signatures and Nd model ages to the B–A period (Fig. 3a–c) likely suggests similar terrigenous source areas. During the Holocene, major Sr–Nd–Pb isotope excursions are dated at ca. 11.5, 10.2, 8.9–8.7, 5.6, 2.2 and 1.1 ka cal. BP (Figs. 3b–c, 4). High and low TDM values of 1.60 Ga at 10.8 ka cal. BP and of 1.34 Ga at 3.4 ka cal. BP (Fig. 3a), respectively, suggest variations in the source rock contributions, from pre-Pan-African and Pan-African and Iberian rocks. However, the Sr, Nd, or Pb isotope compositions did not show any substantial excursion at older TDM values around 10.8 ka cal. BP. In contrast, previous mineralogical data and typically fluvial-derived geochemical ratios from this record have shown an increase in eolian input signalled by higher palygorskite and quartz contents over fluvial-derived material at the early Holocene (Supplementary Fig. S1; Rodrigo-Gámiz et al., 2011). The end of the AHP over northwestern Africa has been documented by terrigenous records and Sr isotopes between 6.0 and 4.0 ka cal. BP (Gasse and Van Campo, 1994; Swezey, 2001; Gasse, 2002; Kuhlmann et al., 2004; Cole et al., 2009). However, the Holocene environmental

18.85

10

12

14 Aller d

YD

16 B lling

8 EMHT

6 7.4ka

4

8.2ka

2

0

conditions in the western vs. eastern Africa have shown to be different and complex. For instance, Sr isotope ratios from the Arabian Sea reflected a first aridification step at 8.5 ka cal. BP followed by a second phase at 6.0 ka cal. BP that ceased at 3.8 ka cal. BP when the modern-day dry climate over North Africa was established (Jung et al., 2004). Moreover, a palaeolake in northern Kenya recorded a dry trend just before the 8.2 ka cal. BP cooling event (Junginger et al., 2014 and references therein). In contrast, in western and northern Africa water-level variations were described around 10.5 to 8.5–8.0 and 7.5–4.5 ka cal. BP (Gasse and Van Campo, 1994; Swezey, 2001; Gasse, 2002; Lézine et al., 2011). In the Iberian Mediterranean region, pollen sequences reflected a significant environmental variability during the Holocene (Carrión, 2002; Combourieu Nebout et al., 2009; Carrión et al., 2010; Fletcher et al., 2010). The major transition from a mesophytic optimum (warm and humid conditions) with relatively high lake levels to xerophytic conditions and high fire activity is traced to the period from 7.5 to 5.0 ka cal. BP (Carrión et al., 2001a, 2001b). Major Saharan dust inputs in a southern Iberian alpine lake have been documented to increase

IACP

244

18

20 LGM

H1

a)

204

18.80 18.75

206

18.70 18.65

b) 15.80 15.70

204

15.75

f)

2.490 2.488 2.486 2.484 2.482 2.480 2.478 2.476

e)

1a 1b

Holocene

0

2

4

6

GS-1

8

10

12

1c

1d 1e

14

2b

2a

GI-1

206

0.839 0.838 0.837 0.836 0.835 0.834 0.833 0.832

207

d)

207

2.080 2.078 2.076 2.074 2.072 2.070 2.068 2.066

15.55

208

208

206

208

204

15.60

c)

39.1 39.0 38.9 38.8 38.7 38.6

207

15.65

GS-2

16

18

20

Fig. 4. 206Pb/204Pb (a), 207Pb/204Pb (b), 208Pb/204Pb (c), 207Pb/206Pb (d), 208Pb/206Pb (e), and 208Pb/207Pb (f) ratios in marine core 293G (b37 μm fraction) plotted against age. Error bars are smaller than symbol size if they are not shown. Grey bars show periods with more unradiogenic Pb ratios.

M. Rodrigo-Gámiz et al. / Chemical Geology 410 (2015) 237–250

from 7.0 to 6.0 ka cal. BP and increased since then until present (Jiménez-Espejo et al., 2014). Abrupt changes in fluxes preceding the 8.2 ka cold event by 0.2 and 1.0 ka with North African eolian dust inputs over central Europe have also been described in a peat bog record (Le Roux et al., 2012). In particular, a significant period of increased dust deposition described in the ombrotrophic bog during the early–middle Holocene transition (EMHT), at ca. 8.5–8.4 ka cal. BP, was recorded with a decrease of εNd(0) signature up to −12.5, which is comparable with old continental shields from the Sahara (Le Roux et al., 2012). During the same period, enhanced fluxes of Saharan dust deposition have also been found in lake and ice records from Africa (Gasse, 2000; Thompson et al., 2002). Furthermore, changes in fluvial regime were interrupted by a short-term arid event at 8.5–7.3 ka cal. BP in marine records from the eastern Mediterranean (Blanchet et al., 2013). Thus, an arid signal from North Africa seems consistent with the Sr–Nd excursion and εNd(0) values of −11.7 described at 8.9–8.7 ka cal. BP, likely corresponding to the EMHT (Table 1, Fig. 3b–c).

245

During the late Holocene, a strong reduction in tropical trees and Sahelian grassland cover in northern Chad allowed large-scale dust mobilization from 4.3 ka cal. BP (Kröpelin et al., 2008). Around 3.0–2.0 ka cal. BP, a widespread phase of fluvial–lacustrine deposition and eolian stabilization is dated in the Sahara region (Swezey, 2001), while around 2.7 ka cal. BP, a desert ecosystem was established with periods of very severe droughts especially at 2.0–1.2 ka cal. BP (Gasse, 2002; Kröpelin et al., 2008). We have identified major Sr–Nd–Pb shifts at 5.6, 2.2 and 1.1 ka cal. BP (Figs. 3c, 4), but there is not enough correspondence with African source signatures. Therefore, although we have to consider some uncertainties in the chronological control of the different archives, Sr, Nd and Pb isotope variations and Nd model ages obtained in the marine sediment from the western Mediterranean have shown main recognized shifts at 16.7– 16.2 (corresponding with the H1) and 8.9–8.7 ka cal. BP (ca. EMHT). These changes seem related to variations in terrigenous material from different source areas associated with the main phases of eolian dust input in accordance with other nearby palaeorecords.

volcanic end-member

10 5 0

-10

Nd

(0)

-5

Mali -15

Senegal Mauritania

-20 (Mauritania)

-25 -30 -35 0.700

0.705

0.710

0.715

0.720 87

0.725

0.730

0.735

0.740

86

Sr/ Sr

-10.0

-10.5

Nd

(0)

-11.0

European atmospheric input

-11.5

-12.0

Mauritania -12.5 0.713

0.714

0.715

0.716

0.717

0.718

87

Sr/86Sr

Fig. 5. εNd(0) distribution vs. 87Sr/86Sr ratio modified after Grousset et al. (1998). Error bars are smaller than symbol size (Table 1). Data from African (Scheuvens et al., 2013) and European sources (Henry et al., 1994) are showed.

246

M. Rodrigo-Gámiz et al. / Chemical Geology 410 (2015) 237–250 206

Pb/204Pb

18.66 18.68 18.70 18.72 18.74 18.76 18.78 18.80 18.82 18.84 18.86

a) 15.76 15.74 15,72

E

207

Pb/204Pb

15.70 15.68

(EMHT)

N

15.66 15.64 15.62 15.60 15.58

39.2

b) 39.1

E

20

38.9 (EMHT)

208

N

Pb/ 4Pb

39.0

38.8

38.7

38.6 18.66 18.68 18.70 18.72 18.74 18.76 18.78 18.80 18.82 18.84 18.86 206

204

Pb/ Pb

c) 2.078

y=2.2279x+0.2119 R2=0.824

2.076

(EMHT)

2.072

208

206

Pb/ Pb

2.074

2.070

2.068

2.066 0.832

0.833

0.834

0.835 207

0.836

0.837

0.838

206

Pb/ Pb

Fig. 6. (a) 207Pb/204Pb vs. 206Pb/204Pb, (b) 208Pb/204Pb vs. 206Pb/204Pb and (c) 208Pb/206Pb vs. 207Pb/206Pb ratios for 293G core (b37 μm fraction). Error bar values are shown in Table 2. Dashed contour represents Pb isotope composition of granitic K-feldspars from North Africa (N) and European sources (E) (Juteau et al., 1986; Fagel et al., 2004).

M. Rodrigo-Gámiz et al. / Chemical Geology 410 (2015) 237–250

4.2. Terrigenous provenance We have compiled Sr, Nd, and Pb isotopic data as well as TDM from available literature as possible source areas of eolian and riverine derived terrigenous input to the westernmost Mediterranean (Fig. 1a). Based on published literature, the Sr–Nd isotopic composition of the eolian dust from the Saharan region can be explained by a mixture of three main end-members: an almost constant contribution of a volcanic end-member that can be considered as the background, and two crustal end-members, i.e. the Archean (2.3-to-3 Ga) and post-Archean “Birimian” (2.3-to-1.7 Ga) geological formations (Grousset and Biscaye, 2005). Thus, the “North Africa domain” approximately includes a mixture of these three end-member domains. Three sub-provinces can be also distinguished according to the geology of the source areas: Northern sources (Morocco, Algeria, Mauritania and Mali); Eastern/Southern sources (Libya, Chad, Guinea, Senegal) and the Archean Saharan shield, with its main outcrops located in Mauritania (Grousset and Biscaye, 2005). The combination of 87Sr/86Sr and εNd(0) values (Fig. 5) shows a distribution of samples in two sub-clusters, i.e., Fields I and II (Fig. 5). Field I comprises northern African regions such as Morocco, Mauritania, and Mali, as well as more remote regions such as Niger and Guinea, and is the main source cluster over the time period studied. Conversely, Field II corresponds to isotopic signatures from the present Senegal region (Grousset et al., 1998). Specifically, those sediment samples clustering in Field II are dated at 16.7 ka (last H1 event) and 8.9 ka cal. BP, supporting that the shifts in the 87Sr/86Sr ratio at these periods (Figs. 3c, 5) mainly recorded variations in source areas rather than weathering effects. Combining TDM data compiled (Fig. 1a), Field I will be characterized by TDM around 1.0–1.8 Ga, and Field II will have higher values of ca. 2.5 Ga (Bea et al., 2010). However, the two data points included in Field II do not show particularly old TDM values, i.e. 1.37–1.46 Ga (Table 1 and Fig. 3a). These two TDM data may suggest high mobilization and contribution of mixed material from North Africa, corresponding to Field I, and southern European areas (Nägler, 1990), although some older material transported from Senegal (Field II) cannot be completely discarded (Fig. 1a). TDM values are generally lower than 1.46 Ga, except at 10.8 ka cal. BP that is recorded the oldest Nd model age (1.60 Ga, Table 1 and Fig. 3a), pointing out to a main contribution of material from the Anti-Atlas African granitoids from Field I and Iberian Cambro-Ordovician rocks. Low εNd(0) values at 10.2, 9.4, 6.8, 4.3 and 1.1 ka cal. BP (Fig. 3b) may be also related to typical values of material derived from the Rhône and Têt Rivers in the western Mediterranean (Henry et al., 1994; Stille and Schaltegger, 1996), with a relative contribution of southern Iberian rocks from Ossa Morena due to the input of southern Iberian rivers (López-Guijarro et al., 2008) (Fig. 1a). Concerning Pb isotope ratios, obtained values reflect in general a typical eolian Saharan dust signature (Grousset et al., 1994; Kylander et al., 2005; Erel and Torrent, 2010). Cross-plots of 207Pb/204Pb vs. 206 Pb/204Pb (Fig. 6a) and 208Pb/204Pb vs. 206Pb/204Pb (Fig. 6b) show that unradiogenic Pb signatures correspond to a typical composition of granites from North Africa (206Pb/204Pb = 18.74, 207Pb/204Pb = 15.66, and 208Pb/204Pb = 38.80; Juteau et al., 1986), whereas the signatures outside this group, however, could be associated with a European granite source (Juteau et al., 1986). A linear correlation in 207Pb/204Pb vs. 206Pb/204Pb (Fig. 6a) and 208Pb/204Pb vs. 206Pb/204Pb (Fig. 6b) can be drawn, suggesting the presence of two distinct Pb components of similar age, but different time-integrated Th/U ratios. According to potential European Pb sources described by Fagel et al. (2004), the Pb signature could correspond to upper and post-Paleozoic materials related to the Variscan orogeny from Western Europe. Moreover, the cross-plot of 208Pb/206Pb vs. 207Pb/206Pb shows a good linear correlation (R2 = 0.824, Fig. 6c), with some outliers during the late Holocene corresponding to 4.3, 3.4 and 2.2 ka cal. BP. These events are also characterized with high Sr isotopic values over the last 20 ka (Fig. 3c) and minimum TDM values, between 1.34 and 1.38 (Fig. 3a), as result of a contribution of Iberian and Anti-Atlas sediments. More radiogenic Sr

247

and Pb isotopic signatures have been described in sediments deposited in the Gulf of Cadiz and the Portuguese margin, implying major contributions of eroded material from Iberia with respect to a North African source (Stumpf et al., 2011). Moreover, differences between Pb signals from the Saharan dust and from ancient Iberian mining and smelting ca. 3.2 ka BP were identified in a northern Iberian peat bog (Kylander et al., 2005). Similarly, some southern Iberian records have shown a significant anthropogenic lead pollution signal by early metallurgy since 3.9 ka cal. BP (García-Alix et al., 2013). However, we do not have clear evidences in our record for this kind of potential lead pollution during these periods. Therefore, natural Pb signatures in the marine sediment record from the westernmost Mediterranean can also be explained in terms of simple binary mixing between North Africa and European sources (Fig. 6c). Variations along the mixing line reflect changes in the relative mixing proportions of the two-natural Pb end-member sources. The more radiogenic Pb source corresponds to a European source, and the less radiogenic end member to a North Africa provenance. Enhanced signatures of unradiogenic Pb occurred during colder periods such as the last H1 event (16.7 ka cal. BP), while more radiogenic Pb signatures dominated during most of the time period, particularly during warmer time intervals such as the Allerød period (13.5 ka cal. BP) and some Holocene intervals (11.1, 9.4, 1.1 ka cal. BP). Thus, the combination of the three isotopic systems (Nd, Sr, Pb) and Nd model ages in the marine record displays broadly concordant variations, providing a first approach of terrigenous input provenance in the westernmost Mediterranean. Even though further Sr, Nd and Pb signatures and TDM data from some potential source areas are needed to provide a more accurate source discrimination, results mainly reflect a mixture of terrigenous material from North African and South European sources linked with prevailing climate conditions rather than a strong influence of a particular area. 4.3. Transport mechanisms Traditional interpretations about transport mechanisms of terrigenous material based on grain size distribution indicated that small grain sizes (b 6 μm) are indicative of riverine runoff, and coarse grain sizes (N 6 μm) are in general associated with eolian input (Sarnthein et al., 1981). However, we cannot project different transports based on the physical sorting since the grain size distribution in the marine record showed always values lower than 6 μm (Fig. 3d; Rodrigo-Gámiz et al., 2011). Interestingly, major sizes (N 3 μm; Fig. 3d) are almost contemporaneous with major shifts in Sr isotopes, i.e. less radiogenic 87 Sr/86Sr ratio (Fig. 3c), which could suggest to be an indicator for high humidity in the source area. Nevertheless, some of these shifts, like at 16.70 ka cal. BP have occurred during less humid periods, and therefore we cannot establish a strong relationship between grain size and Sr composition variations. The terrigenous material deposited in the westernmost Mediterranean has been derived by either atmospheric inputs or river runoff from South European and North African sources. The main southern European rivers, due to the scarcity of rivers along North Africa, are (Fig. 1a): (1) the Rhône River draining detrital sediments from the Alps (Juteau et al., 1986; Nägler, 1990; Stille and Schaltegger, 1996), which were further transported from the Gulf of Lion to the Alboran Sea by the oceanic gyre (Millot, 1999); (2) the Guadalhorce and Guadalfeo Rivers from southern Iberia that discharge material in the Alboran Sea as torrential or sporadic runoff due to their nonpermanent feature; and (3) the Guadiana and Guadalquivir Rivers, located in the southwestern Iberia, which fluvial discharge can transport high amounts of detrital particulate matter from the Guadalquivir basin (Erel and Torrent, 2010) facilitating the sediment distribution into the western Mediterranean by Atlantic inflow waters (Grousset et al., 1998). Regarding atmospheric transport, some pressure configurations have been described to cause Saharan dust input to the Iberian Peninsula:

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(1) low pressures over West and/or Southwest Portugal; (2) high pressure over the East or Southeast Iberian Peninsula; and (3) the combination of both low and high pressure systems (Avila et al., 1998; Rodríguez et al., 2001). Interannual variations in dust transport to the Atlantic Ocean have been correlated with the climatic variability defined by the NAO (Hurrell, 1995; Moulin et al., 1997). Furthermore, terrigenous input is largely controlled by the migrations of the ITCZ, which also determines the position of the subtropical Azores anticyclone. ITCZ migrations are associated with alternating dry and wet periods according to the monsoon activity fluctuations. These are accompanied by changes in the wind regime and intensity, as well as variable coverage of land by vegetation, and variations in rainfall amount and river discharge (Clemens and Prell, 1990; Gasse and Van Campo, 1994; Gasse, 2000). During summer a northward shift in the ITCZ position located around 20–30°N leads to a more northerly position of the monsoon rain belt system, decreasing the aridity over North Africa and causing northward expansion of the Sahelian savannah into the Sahara region (e.g., Moulin et al., 1997; Bergametti et al., 1989a; Rodríguez et al., 2001; Torres-Padrón et al., 2002). This situation, together with low wind intensity, would reduce the dust emission from the southern and southwestern African regions, favouring terrigenous transport from the northern African regions to the western and central parts of the Mediterranean (Moulin et al., 1997). In general, most of the terrigenous sediment cluster in the northern African (Western Sahara, Morocco, Mauritania, Mali, Niger, and Algeria), southern Iberia and southern France areas (Fig. 1a). In contrast, the southernmost position of the ITCZ during winter leads to a more southerly position of the anticyclone centre and the monsoon rain belt system, increasing the aridity in North Africa and decreasing the cover vegetation (e.g., Matthewson et al., 1995; Arbuszewski et al., 2013; McGee et al., 2013, 2014). This configuration accompanied by major wind intensity gives rise to greater eolian dust mobilization from latitudes of 10–20°N and transport toward the westernmost Mediterranean. This situation could explain a second terrigenous source area represented by more unradiogenic terrigenous material like old Archaean rocks from the present day Senegal region recorded at 16.7–16.2 (H1) and 8.9–8.7 ka cal. BP (EMHT) (Fig. 1a). Therefore, a combination of changes in wind intensity, the displacement and distance of the anticyclone centres and ITCZ migrations as well as the availability of the eolian material according to the vegetation cover, are the main factors controlling the different transport mechanisms of terrigenous material deposited in the western Mediterranean during the last 20 ka. 5. Conclusions Radiogenic isotope (Sr, Nd, Pb) signatures and Nd model ages have allowed us to reconstruct source areas and transport mechanisms of terrigenous material in the westernmost Mediterranean since the Last Glacial Maximum. Nd, Sr and Pb isotope compositions from carbonatefree sediments of the b 37 μm size fraction and Nd model ages have been integrated with previous geochemical and mineralogical data for such reconstruction. Lower 87Sr/86Sr signatures reflect either the influence of chemical weathering or variations in terrigenous provenance. Contemporaneous shifts in the 87Sr/86Sr ratio and εNd(0) during the H1 (at 16.7–16.2 ka cal. BP) and the EMHT (at 8.9–8.7 ka cal. BP) suggest variations in the provenance of terrigenous material during these arid and cold periods. The isotopic signatures in the terrigenous component showed an older contribution of likely old Archaean rocks from the present-day Senegal region. In contrast, warmer periods display radiogenic signatures that indicate terrigenous provenance from the northern Africa (Western Sahara, Morocco, Mauritania, Mali, Niger, and Algeria), southern Iberia and southern France areas. Furthermore, although Pb ratios do not discriminate among potential sources in North Africa or South Europe, they reflect the mixture between two-

natural Pb sources, i.e., a radiogenic European source, transported by river runoff or atmospheric input mainly during the Holocene, and a less radiogenic North Africa source during the glacial period. Variations in TDM further support the mixture of pre-Pan-African and Pan-African and Iberian derived-material in the western Mediterranean marine sediments. Migrations of the ITCZ along with changes in the intensity and regime of the wind system and the African monsoon, and the position of the subtropical Azores anticyclone seem to be the main transport mechanisms controlling terrigenous provenance. Supplementary data to this article can be found online at http://dx. doi.org/10.1016/j.chemgeo.2015.06.004.

Acknowledgments This work was supported by the European Regional Development Fund (ERDF)-cofinanced grants CGL2012-32659, CGL2009-07603 (Secretaría de Estado de Investigación, Desarrollo e Innovación, MINECO), Project RNM-5212 and Research Group RNM-179 (Junta de Andalucía). We are also grateful to the oceanographic cruise Training Through Research Programme (UNESCO-Moscow State University). We would like to thank F. Grousset for his helpful comments on an early version of the manuscript and P. Montero and F. Bea for the TDM data provided. We likewise thank M. Senn-Gerber and D. Fontignie for their laboratory assistance. J. L. Sanders post-edited the English style. Comments and suggestions from two anonymous reviewers have substantially improved the final version of the manuscript. References Abouchami, W., Näthe, K., Kumar, A., Galer, S.J.G., Jochum, K.P., Williams, E., Horbe, A.M.C., Rosa, J.W.C., Balsam, W., Adams, D., Mezger, K., Andreae, M.O., 2013. Geochemical and isotopic characterization of the Bodélé Depression dust source and implications for transatlantic dust transport to the Amazon Basin. Earth Planet. Sci. Lett. 380, 112–123. Arbuszewski, J.A., deMenocal, P.B., Cléroux, C., Bradtmiller, L., Mix, A., 2013. Meridional shifts of the Atlantic intertropical convergence zone since the Last Glacial Maximum. Nat. Geosci. 6, 959–962. Arndt, N.T., Goldstein, S.L., 1987. Use and abuse of crust-formation ages. Geology 15, 893–895. Avila, A., Alarcón, M., Queralt, I., 1998. The chemical composition of dust transported in red rains — its contribution to the biochemical cycle of a holm oak forest in Catalonia (Spain). Atmos. Environ. 32, 179–191. Bea, F., Montero, P., Talavera, C., Abu, Anbar M., Scarrow, J.H., Molina, J.F., Moreno, J.A., 2010. The palaeogeographic position of Central Iberia in Gondwana during the Ordovician: evidence from zircon chronology and Nd isotopes. Terra Nova 22, 341–346. Bergametti, G., Dutot, A.L., Buat-Ménard, P., Remoudaki, E., 1989a. Seasonal variability of the elemental composition of atmospheric aerosol particles over the northwestern Mediterranean. Tellus 41B, 553–561. Bergametti, G., Gomes, L., Remoudaki, E., Desbois, M., Martin, D., Buat-Ménard, P., 1989b. Present transport and deposition patterns of African dusts to the north-western Mediterranean. In: Leinen, M., Sarnthein, M. (Eds.), Paleoclimatology and Paleometeorology: Modern and Past Patterns of Global Atmospheric Transport. NATO ASI Series, Mathematical and Physical Sciences vol. 282. Kluwer, Boston, pp. 227–250. Biscaye, P.E., Dasch, E.J., 1971. The rubidium, strontium, strontium-isotope system in deep-sea sediments: Argentine Basin. J. Geophys. Res. 76, 5087–5096. Blanchet, C.L., Tjallingii, R., Frank, M., Lorenzen, J., Reitz, A., Brown, K., Feseker, T., Brückmann, W., 2013. High- and low-latitude forcing of the Nile River regime during the Holocene inferred from laminated sediments of the Nile deep-sea fan. Earth Planet. Sci. Lett. 364, 98–110. Blum, J.D., Erel, Y., 1997. Rb–Sr-isotope systematics of a grantic soil chronosequence: the importance of biotite weathering. Geochim. Cosmochim. Act 61, 3193–3204. Bout-Roumazeilles, V., Combourieu, Nebout N., Peyron, O., Cortijo, E., Landais, A., Masson-Delmotte, V., 2007. Connection between South Mediterranean climate and North African atmospheric circulation during the last 50,000 yr BP North Atlantic cold events. Quat. Sci. Rev. 26, 3197–3215. Box, M.R., Krom, M.D., Cliff, R.A., Bar-Matthews, M., Almogi-Labin, A., Ayalon, A., Paterne, M., 2011. Response of the Nile and its catchment to millennial-scale climatic change since the LGM from Sr isotopes and major elements of East Mediterranean sediments. Quat. Sci. Rev. 30, 431–442. Braun, J.J., Viers, J., Dupre, B., Polve, M., Ndam, J., Muller, J.-P., 1998. Solid/liquid REE fractionation in the lateritic system of Goyoum, East Cameroon: the implication for the present dynamics of the soil covers of the humid tropical regions. Geochim. Cosmochimi. Act 62, 273–299. Caquineau, S., Gaudichet, A., Gomes, L., Legrand, M., 2002. Mineralogy of Saharan dust transported over northwestern tropical Atlantic Ocean in relation with source regions. J. Geophys. Res. 107, 4251.

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