46 | CHAPTER 3 3.pdf · Recycling of oceanic crust and sediment back into the mantle occurs at subduction zones. Fluids and melts expelled from the subducting slab are an important

46 | CHAPTER 3

EASTERN MEDITERRANEAN SEDIMENTS | 47

48 | CHAPTER 3

ABSTRACT

The Eastern Mediterranean Sea (EMS) is the last remnant of the Tethys Ocean that has been

subducted to the north since the Jurassic. Subduction has led to the formation of multiple

island arcs in the EMS region where the Aeolian and Aegean arcs are currently active. The

EMS is surrounded by continents and receives a large sediment input, part of which is

transported down with the subducting slab into the mantle and potentially contributes a

major flux to the arc volcanism. An along-arc gradient in the composition of subducting

sediment has been evoked to explain the distinct geochemical signature of the easternmost

volcanic centre of the Aegean arc, but direct evidence for this proposal is lacking. We present

a detailed study of the mineralogical, major element, trace element and Sr-Nd-Hf-Pb isotope

composition of 45 Neogene EMS sediment samples obtained from Deep Sea Drilling Project

(DSDP) and Ocean Drilling Program (ODP) drill sites and box cores to characterise their

geochemical composition, distinguish provenance components and investigate the temporal

and spatial variation in provenance to evaluate the potential changing contribution of

subducted EMS sediment to Aegean and Aeolian arc volcanism.

Based on trace element characteristics of EMS sediments, we can distinguish four

provenance components. Nile sediment and Sahara dust are the main components, but

contributions from the Tethyan ophiolite belt and arc volcanic rocks in the north are also

recognised. Pliocene and Quaternary EMS sediment records a strong geochemical gradient

where Nile River sediment entering the EMS in the east is progressively diluted by Sahara

Desert dust towards the west. Pre-Messinian samples, however, have a remarkably

homogeneous composition close to Nile sediment. We relate this rapid increase in Sahara

dust contribution to a late Miocene climate shift leading to decreased Nile runoff and

aridification of the Sahara region. EMS sediment has a restricted range in Pb isotopes

compared to Aegean volcanic rocks and therefore cannot account for the total variation in

Aegean arc lavas. The bulk addition of 0.5-5% of EMS sediment satisfactorily explains most

of the variation in volcanic rocks from the western and central centres, but the low 207Pb/204Pb

and 206Pb/204Pb seen in the eastern volcanic centre Nisyros requires a distinct mantle source.

In the Aeolian arc, a subducted EMS sediment input can only be resolved in Stromboli volcanic

rocks.

Previous pages: Southwesterly winds sweep Sahara dust across the Eastern Mediterranean Sea. Satellite

image acquired on February 1, 2015, with the Moderate Resolution Imaging Spectroradiometer (MODIS).

Image courtesy of NASA's Earth Observatory.


1. INTRODUCTION

1.1. Sediment subduction

Recycling of oceanic crust and sediment back into the mantle occurs at subduction zones.

Fluids and melts expelled from the subducting slab are an important geochemical component

of subduction zone magmas (e.g., Elliott, 2003). The influence of the amount and type of

subducted sediment has been demonstrated in multiple studies wherein along-arc variation

in subducted sediment composition is reflected in geochemical trends in the arc volcanic rocks

(e.g., White and Dupré, 1986; Plank and Langmuir, 1993; Vroon et al., 1995; Carpentier et

al., 2008). In this paper, we report the geochemical composition of sediment in the Eastern

Mediterranean Sea (EMS) in order to investigate its influence on Aeolian and Aegean arc

volcanism. The EMS receives a large influx of aeolian and riverine sediment derived from

sources with contrasting geochemical fingerprints. Hence, generic estimates of subducting

sediment composition such as GLOSS (GLObal Subducting Sediment; Plank and Langmuir,

1998; Plank, 2013) might not be applicable to the EMS. Previous work on Quaternary EMS

sediment has indicated a strong geochemical east-west gradient that is caused by mixing

between contrasting provenance areas: “young and mafic” Nile sediment is diluted towards

the west by “old and felsic” Sahara dust (Venkatarathnam and Ryan, 1971; Krom et al.,

1999b; Weldeab et al., 2002a). This sedimentary gradient has been evoked to explain along-

arc trends in the volcanic rocks of the Aegean arc and most notably the distinct isotope

composition of the easternmost volcano Nisyros (Francalanci et al., 2005; Elburg et al., 2014).

However, the existing EMS sediment dataset is limited to major element and Sr-Nd isotope

data of Quaternary pelagic sediments and therefore lacks the temporal and geochemical

resolution to adequately resolve these questions. In particular, there is no evidence whether

the E-W sedimentary gradient is also present in pre-Pliocene sediment, which is significant

because the age of the sediment currently underneath the Aegean arc is at least 5 Ma. The

aim of this study is: (1) to complement the existing dataset with trace element and Pb-Hf

isotope data and (2) to expand the record back in time in order to investigate the spatial and

temporal variations in EMS sediment geochemical composition. We selected 45 sediment

samples from seven DSDP/ODP sites and six box cores distributed along the Aegean and

Aeolian subduction zones that range from Middle Miocene to Quaternary in age. Combined

mineralogical and geochemical analyses of these samples are used to distinguish provenance

areas that can be used to constrain elemental fluxes into the Aeolian and Aegean subduction

zones.

1.2. Geological setting

The Eastern Mediterranean Sea (EMS) is the last remnant of the Tethys Ocean that has largely

been consumed by the convergence of the African and Eurasian plates. Continuous

northward subduction of Tethyan oceanic crust since the Jurassic has led to the accretion of

several terranes to the Eurasian continent, which together constitute the stacked nappes of

50 | CHAPTER 3

the Hellenide Complex in Greece (van Hinsbergen et al., 2005; Ring et al., 2010). Incipient

continental collision caused a southward shift of the active deformation and uplift in the

forearc (Peloponnesus, Crete) region around 15 Ma (Le Pichon et al., 2002; ten Veen and

Kleinspehn, 2003). Sediments of the EMS are currently being accreted to the Hellenide

Complex wherein the Messinian evaporites act as the detachment level. Pliocene and

Quaternary sediments are incorporated in an accretionary prism: the Mediterranean Ridge (Le

Pichon et al., 2002; Kopf et al., 2003; Figure 1). The Hellenic Trench is thus no longer the

active deformation front, but represents a sediment-starved section of the fore-arc that is still

present as a topographical feature. As a result of the near-complete subduction of oceanic

crust, only three isolated patches of abyssal plain remain in the EMS: the Messina and Sirte

abyssal plains in the Ionian Basin and the Herodotus Abyssal Plain NW of the Nile fan.

The continuous northward subduction of Tethyan oceanic crust has led to the

development of island arc volcanism along the southern margin of Eurasia since the

Paleogene. Tyrrhenian and Aegean back-arc extension due to slab rollback in the Neogene

has caused the rotation and separation of the Tethyan subduction zone into two branches:

the Aeolian and Hellenic subduction zones (Robertson and Grasso, 1995; Figure 1). Volcanism

in the currently active Aeolian and Aegean arcs started in the Pliocene. Notable active

volcanoes in the Eastern Mediterranean region include Stromboli in the Aeolian arc and

Santorini in the Aegean arc. Etna on Sicily is situated in a complex geodynamic setting but is

also potentially influenced by EMS subduction (e.g., Viccaro and Cristofolini, 2008).

1.3. Sedimentation in the EMS

Due to the proximity of surrounding continents, the EMS receives a large amount of

terrigenous sediment from aeolian and riverine sources. The two main components in EMS

sediment are aeolian dust from the Sahara region and sediments delivered by the river Nile

(Venkatarathnam and Ryan, 1971; Krom et al., 1999b; Weldeab et al., 2002a). The Nile is

one of the largest rivers on Earth and its catchment covers a large portion of East Africa. Three

main branches dominate the total discharge of the Nile: the White Nile drains Lake Victoria

whereas the Blue Nile and Atbara River originate in the Ethiopian highlands. The latter two

cover 56 % of the annual discharge and >95 % of the sediment load of the Nile (Foucault

and Stanley, 1989; Padoan et al., 2011). Hence, Nile sediment is dominated by detritus of the

Ethiopian flood basalt province that was emplaced at ca. 30 Ma (Pik et al., 1999) and has a

“young and mafic” geochemical signature. The Nile sediment flux into the EMS prior the

construction of the Aswan Dam in 1964 is estimated at 120 Mt/yr (Milliman and Meade,

1983). When sediment from the Nile enters the EMS, the coarser fraction is deposited in the

Nile fan and the finer fraction is transported in a counter clockwise gyre in the Levantine basin

(Venkatarathnam and Ryan, 1971; Krom et al., 1999b; Weldeab et al., 2002a). To the west

of the Nile fan, Nile sediment is diluted by aeolian dust from the Sahara region. An estimated

70-100 Mt/yr (Ganor and Foner, 1996) of Sahara dust is transported to the Mediterranean in

discrete dust storms. Dust supplied to the EMS is mainly derived from the Western Sahara,


Figure 1. Map of the Eastern Mediterranean Sea with the most important geological features and the

location of the DSDP/ODP drill sites used in this study. Volcano abbreviations: St – Stromboli, V – Vulcano,

E – Etna, SG – Saronic Gulf cluster (Methana, Aegina, Poros), M – Milos, Sa – Santorini, N – Nisyros.

Hoggar Massif and Libyan Desert from where it is transported by south-westerly winds along

the northern margin of the Sahara (Goudie and Middleton, 2001). These areas consist

predominantly of the Proterozoic Saharan Shields and Phanerozoic limestone (Stein et al.,

2007), and hence their geochemical composition can be summarised as “old and felsic”.

Another potential major source of Sahara dust is the area of Lake Chad, but the Harmattan

winds cause the majority of dust from this area to be transported to the west and not towards

the Mediterranean (Prospero et al., 2002).

The relative contributions of Nile detritus and Sahara dust to EMS sediment during the

Quaternary are largely controlled by climate oscillations (e.g., Revel et al., 2010; Box et al.,

2011). Orbitally-induced variations in insolation in subtropical Africa have a pronounced effect

on sedimentation patterns in the EMS. During precession minima, intensification of the

African monsoon leads to increased Nile runoff (Rossignol-Strick, 1985) while enhanced

humidity and vegetation in the Sahara allow for less dust production (Larrasoaña et al., 2003),

and vice versa during precession maxima. Increased Nile runoff during pluvial periods is

thought to be the main driving force behind deposition of organic-rich sapropels in the

Mediterranean (e.g., Freydier et al., 2001; Weldeab et al., 2002b; Scrivner et al., 2004;

Ducassou et al., 2008). It has been suggested that these wet-dry oscillations already occurred

in the Pliocene (Wehausen and Brumsack, 1999; Larrasoaña et al., 2003; Trauth et al., 2009).

In addition to Sahara dust and Nile sediment, rivers in the northern borderlands (around the

Aegean Sea, Turkey) supply an unknown amount of sediment to the EMS. The hinterland of

A

egean Arc

SG

M

SaN

St

V

E

Aeolian Arc

Nile Fan

M

d

edit

a

e

err

nan Ri ge

Messina

Sirte

H

ero

d

otus

Outer deformation front

Inner deformation front

Hellenic Trench

Abyssal plain

Volcanic arc with active volcanoes

DSDP & ODP sampling site

Surface sediment sampling site

Greece

Cyprus

Libya

Turkey

Egypt

Levant

200 km

374

125126

128

Libya

971B

378

130

Nile

375-376

Nile

52 | CHAPTER 3

these rivers consists predominantly of Cretaceous to Cenozoic limestones and Cenozoic

ultramafic rocks of the Tethyan ophiolite belt (Weldeab et al., 2002a). Due to the generally

low trace element concentrations in limestones, the latter are expected to govern the trace

element and isotope budget in sediment derived from the northern borderlands. In the

western part of the EMS, the Ionian Basin probably receives an influx of sediment from the

Adriatic that is derived from the Alps.

2. ANALYTICAL TECHNIQUES

All sample preparation and analyses were carried out at the Vrije Universiteit Amsterdam.

Sediment samples were dried at 50 °C for at least 3 days and subsequently powdered with

agate ball mills. Thermogravimetric analysis (TGA) was carried out with a LECO TGA-601.

Major element concentrations were measured by X-ray fluorescence spectroscopy (XRF) on a

Panalytical MagiX Pro on fused glass beads of Li2B4O7/LiBO2 with a 1:4 dilution. Interference-

corrected spectra intensities were converted to oxide-concentrations against a calibration

curve consisting of 30 international standards. Results were corrected for loss on ignition (LOI)

using the TGA data; iron concentrations are expressed as total ferrous iron (FeO*). Incomplete

dissociation of carbonates, sulphates and clay minerals during ignition led to low and variable

totals. However, replicate analysis (2-3 per sample) normalised to 100% agreed within 3%.

For isotope and trace element analysis, approximately 80 mg of whole rock powder was

digested in a HF-HNO3 mixture in PTFE Parr-bombs for four days at 200 °C and subsequently

transferred to clean PFA vials. Quantitative recovery was achieved by fluxing the bombs with

HCl-HF for 1 day at 200 °C and adding this mixture to the sample solution. Trace element

concentrations were determined by a ThermoFisher X-series-II inductively coupled plasma

mass spectrometer (ICP-MS) with USGS reference material BHVO-2 as calibration and

instrumental drift correction standard, using a method modified after Eggins et al. (1997).

Repeated analyses of USGS reference material BCR-2 indicate accuracy and precision of better

than 10% (2 RSD) for all reported trace elements. Strontium and neodymium isotopes were

separated using conventional ion-exchange techniques and were measured in static

acquisition mode on a Finnigan MAT 262 (Sr) and TritonPlus (Nd) thermal ionisation mass

spectrometers (TIMS). Mass fractionation was corrected by normalising to 86Sr/88Sr = 0.1194

and 146Nd/144Nd = 0.7219; the in-house SIGO Nd standard reagent yielded 0.511334 ± 18

(2σ, n=27), equivalent to a La Jolla reagent standard value of 0.511844 (Griselin et al., 2001).

Hafnium was separated from the matrix according to the method outlined in Münker et al.

(2001) and measured on a Neptune multi-collector ICP-MS following Morel et al. (2008). Data

were corrected offline for isobaric interference and mass fractionation (Vervoort et al., 2004;

Cecil et al., 2011). The JMC-475 Hf standard reagent yielded 176Hf/177Hf = 0.282158 ±

0.000022 (n=112). Lead isotopes were separated by double-pass HBr-anion exchange

chemistry and measured on a TritonPlus TIMS using a 207Pb-204Pb double spike to correct for

mass fractionation (Thirlwall, 2000). Standard reagent NBS 981 (n=12) gave 16.9408 ±


0.0009 (206Pb/204Pb), 15.4982 ± 0.0010 (207Pb/204Pb) and 36.7213 ± 0.0031 (208Pb/204Pb). Total

procedural blanks are <120 pg for Sr, <5 pg for Nd and <30 pg for Pb and Hf and therefore

negligible. Results for external standards are listed in the online supplementary material.

3. SITE DESCRIPTIONS

A representative suite of Eastern Mediterranean sediment samples distributed along the

Aegean subduction zone was obtained from Deep Sea Drilling Project (DSDP) and Ocean

Drilling Program (ODP) drill cores (legs 13, 42A and 160) and box core samples from De Rijk

et al. (1999) (Figure 1). Due to the high sedimentation rates in the EMS, oceanic basement

was not reached in any of the sites and the maximum age of recovered sediment is ca. 15 Ma

(Langhian). The age of the EMS sediment samples is obtained by interpolation of the

sedimentation rate between biostratigraphical markers as reported in the DSDP and ODP site

reports (see sections 3.1 to 3.7 for references). Care was taken during sample selection to

avoid sections of core where hiatuses or tectonics complicated the determination of

sedimentation rates. The age of the samples is at least correct to the stage level (e.g. Upper

Pliocene) on the basis of biostratigraphy, with errors probably less than 20% on ages quoted

in Table 1 as obtained from the interpolated sedimentation rates. A summary of the samples

is given in Table 1; full sample information is provided in the supplementary material.

3.1. DSDP site 374

DSDP site 374 is located in the Messina abyssal plain in the Ionian Basin, south of Calabria,

and contains a continuous sedimentary record from the Quaternary to the Messinian

evaporites. The Pliocene to Quaternary sequence consists of hemipelagic nannofossil marls

and intercalated sapropels. An upward increase in the abundance of sand/silt layers and

sedimentation rate and a decrease in carbonate content in this sequence suggest an increase

in terrigeneous input since the Early Pliocene. The Miocene evaporites are overlain by a ~25

m thick sequence of dolomitic mudstone of Lago Mare facies, likely deposited in an alkali lake

or sea environment (Hsü et al., 1978c).

3.2. DSDP sites 125 and 126

DSDP sites 125 and 126 on the Mediterranean Ridge together contain the most complete

record of EMS sediment. However, a ~1.5 Myr hiatus is present in the lower Pliocene of site

125 (Ryan et al., 1973c, d). The Pliocene and Quaternary sediments at site 125 consist of

uniform carbonate-rich pelagic sediment with sapropels. The Upper Miocene comprises

evaporites overlain by Lago Mare dolomitic mudstones that are similar to those encountered

at site 374. Site 126 is situated in a tectonic cleft in the Mediterranean Ridge roughly 100 km

NE of site 125. The cleft cuts through the Messinian evaporite series, providing access to pre-

evaporite grey-green shales of Serravalian age.

54 | CHAPTER 3

Table 1. EMS sediment sample information. The age of the sample is interpolated based on

sedimentation rates reported in the DSDP/ODP site descriptions; mineralogy is derived from

thermogravimetric analysis. Abbreviations: carb. = carbonate (calcite + dolomite); gyps. = gypsum; ill./chl.

= illite and/or chlorite; smec. = smectite; kaol. = kaolinite.

site sample ID

374 AMS-034

AMS-035

AMS-036

AMS-037

AMS-038

AMS-039

125 AMS-030

AMS-001

AMS-002

AMS-031

AMS-032

126 AMS-033

128 AMS-003

AMS-004

AMS-005

AMS-006

AMS-007

Libya AMS-024

AMS-025

AMS-026

971B AMS-019

AMS-020

AMS-021

AMS-022

AMS-023

378 AMS-045

AMS-046

AMS-047

AMS-048

130 AMS-008

AMS-009

AMS-010

AMS-011

AMS-012

Nile fan AMS-027

AMS-028

AMS-029

376 AMS-016

AMS-043

AMS-044

AMS-017

AMS-018

375 AMS-013

AMS-014

AMS-015

depth

m b.s.f.

159

255

335

362

380

393

0

23

55

72

90

129

55

82

148

250

472

box core

box core

box core

8

28

70

109

170

143

220

50

302

15

53

256

412

416

box core

box core

box core

22

41

43

78

155

193

466

676

age

Ma

1.0

1.7

2.9

3.7

4.9

5.4

0.4

1.4

2.7

3.8

5.5

12.5

0.2

0.3

0.6

1.0

1.8

0.0

0.0

0.0

0.5

>7

>7

>7

>7

1.9

2.6

0.8

4.0

0.2

0.8

1.5

1.9

1.9

0.0

0.0

0.0

1.0

3.0

4.5

5.5

6.5

7.0

10.0

14.5

Al O2 3

wt %

16.41

14.38

6.82

4.03

4.76

13.43

9.30

5.90

5.84

5.17

14.42

19.43

10.91

3.61

10.08

9.26

8.16

3.92

4.33

6.40

12.38

13.10

13.38

12.66

13.38

9.37

8.27

7.29

7.52

18.27

5.32

18.19

17.52

5.22

5.24

6.39

7.02

8.77

7.95

6.95

10.21

0.22

1.78

10.20

6.83

CaCO3

wt %

7.3

10.8

47.3

60.3

28.6

11.0

34.2

53.9

52.0

61.0

12.2

6.2

25.7

54.5

32.2

32.8

33.3

65.9

62.5

49.9

16.9

9.9

5.1

4.7

6.3

31.3

29.6

39.8

42.9

6.2

52.1

5.8

4.4

58.8

48.7

43.3

41.0

32.1

40.6

47.8

23.4

2.2

3.8

28.3

61.6

carb.

-

+/-

+

+

++

+/-

++

+++

+++

++

+

-

++

+++

++

++

++

++

++

++

+

+/-

-

-

-

++

++

++

++

-

+++

-

-

+++

++

++

++

++

++

++

++

-

-

++

+++

gyps.

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

+++

+++

-

-

ill./chl.

+/-

-

+

++

-

-

-

-

-

+

+/-

+/-

-

+/-

-

-

-

++

++

+

-

-

-

-

-

-

-

-

-

+/-

+/-

+

+/-

+/-

+

+/-

-

-

-

-

-

-

-

+/-

+

smec.

+

+

+/-

+/-

+

+

+/-

-

-

-

+/-

+/-

+/-

-

+/-

+/-

+/-

-

-

-

+

+

+/-

+/-

+/-

+/-

+/-

+/-

+/-

++

-

++

+

-

+

+

+

+/-

+/-

+/-

+

-

+/-

-

-

kaol.

++

++

+

+/-

+

++

+

+

+

+

++

++

+

-

+

+

+

+/-

+

+

++

++

++

++

++

+

+

+

+

+

+/-

+

++

+/-

+/-

+

+

+

+

+

+

-

-

+

-


3.3. DSDP site 128

DSDP site 128 is located in the Hellenic Trench between Crete and the Peloponnesus. The

sands, silts and oozes in this site represent trench fill sediments and are intercalated with

carbonate turbidites. Due to the high sedimentation rate, only Quaternary sediments were

recovered from site 128. The occurrence of glaucophane crystals suggests that at least part

of the sediment is derived from the Aegean HP/LT metamorphic belt (Ryan et al., 1973a).

3.4. ODP site 971B

The Olimpi mud volcano field is located south of Crete on the Mediterranean Ridge and ODP

site 971B was drilled in the flank of the Napoli mud volcano within this field (Emeis et al.,

1996). The deposits of the Napoli mud volcano consist of clast-bearing muddy debris flows

that are overlain by Quaternary pelagic sediment. Although these mud flows were deposited

in the Pleistocene, they represent remobilised pre-evaporite sediment. Compression in the

Mediterranean Ridge led to the build-up of fluid pressure beneath the evaporite seal near the

backstop of the accretionary prism. Faulting of the evaporite series allowed fluidized mud to

escape to the surface and form mud volcanoes (Robertson, 1996; Huguen et al., 2004).

Microfossil assemblages indicate that the mud matrix of these flows is dominated by Eocene

to Middle Miocene sediment (Emeis et al., 1996). Hence, these mud flows form an average

of pre-Messinian sediment in the EMS.

3.5. DSDP site 130

DSDP site 130 was drilled in the Mediterranean ridge northwest of the Nile fan and close to

the Herodotus abyssal plain. Quaternary sediments recovered from this site are characterised

by a strong bimodal distribution: pelagic marl oozes alternating with carbonate-poor

terrigenous clays to sands. Sedimentary textures indicate that the latter are predominantly

deposited by turbidity currents and their mineralogy points to a Nile provenance (Ryan et al.,

1973b).

3.6. DSDP sites 375 and 376

DSDP sites 375 and 376 on the Florence Rise west of Cyprus both comprise a discontinuous

record from the Middle Miocene to Quaternary, but complement each other and together

yield an almost complete sedimentary record. The sediments are predominantly nannofossil

marls interspersed with sapropels and volcanic ash layers, interrupted by evaporites and Lago

Mare dolomitic marlstone during the Messinian. The Pliocene sequence is extremely

condensed (5.5 m) due to low sedimentation rates and multiple hiatuses (Hsü et al., 1978a).

3.7. DSDP site 378

DSDP site 378 is situated in the Sea of Crete, north of the Hellenic trench and Crete. It contains

a continuous sequence of nannofossil marls and oozes with intercalated sapropels overlying

56 | CHAPTER 3

the Messinian evaporites. In sharp contrast with upper Miocene sections in sites 374, 125 and

376, no dolomitic Lago Mare facies is present (Hsü et al., 1978b).

3.8. Libya and Nile surface sediments transects

The box core samples are obtained from a previous study (De Rijk et al., 1999) and derived

from the continental shelf and rise of the African plate. The samples are located on transects

from 300 to 3000 m water depth.

4. RESULTS

Results of the thermogravimetric and geochemical analyses of 45 EMS sediment samples are

summarised in Tables 1 and 2. The full dataset is available in the online supplementary

material. Sample location, carbonate content and mineralogy are shown in Table 1. Key trace

element ratios and radiogenic isotope compositions are given in Table 2.

4.1. TGA analysis

Stepwise heating and weighing of sample powder from 25 to 1000 °C allowed the

construction of LOI versus temperature diagrams. By combining the LOI curves with first

derivative curves (the speed of weight loss), the mineralogy of the samples could be

determined. However, it proved to be difficult to distinguish between different clay mineral

groups (kaolinite, smectite and illite/chlorite) in carbonate-rich samples and therefore the TGA

mineralogy only provides a rough estimate that is not comparable to X-ray diffraction

analyses. Nevertheless, the distribution of clay minerals in the EMS as derived by TGA is in

good agreement with results from previous studies (Venkatarathnam and Ryan, 1971;

Foucault and Mélières, 2000; Ehrmann et al., 2007; Hamann et al., 2009). Clay-rich sediments

from site 130 (NW of Nile fan) are rich in smectite while the highest kaolinite-proportions are

found on the Mediterranean Ridge. Illite- and/or chlorite-rich sediments are predominantly

found in the Ionian Basin (site 374) and on the shelf of the African plate in Libya and Nile

surface sediment transects.

Carbonate content was determined as the weight loss above 660 °C. Hence, breakdown

of clay minerals in this temperature range (illite and/or chlorite) resulted in carbonate contents

that are erroneously high by a few wt. %. It was not possible to distinguish between calcite,

dolomite and aragonite. Carbonate contents for the EMS sediment suite range from 5 to 60

wt. % and the highest concentrations are encountered in samples from the Mediterranean

Ridge. Pre-Messinian samples are characterised by lower carbonate contents compared to

Pliocene and Quaternary samples. Total organic carbon (TOC) could not be measured

accurately as the breakdown of smectite-group clay minerals and gypsum interfered with the

weight loss due to organic carbon dissociation at temperatures <500 °C.


Figure 2. Whole rock K/Al vs. Mg/Al plot for the EMS sediment samples. Evaporite samples (AMS-013

and 018) and dolomite-rich sample AMS-038 (Mg/Al = 3.2) are excluded from the figure. The EMS

sediment samples overlap with previously published results (Krom et al., 1999a, b; Revel et al., 2010; Box

et al., 2011) and are a mixture of different clay mineral groups. Ideal compositions of chlorite, kaolinite,

smectite and illite are shown for reference (Deer et al., 1992).

4.2. Major and trace elements

The major element composition of the EMS sediment samples is highly variable and SiO2,

Al2O3 and CaO concentrations vary by an order of magnitude. While CaO and SiO2 contents

are proportional to the fraction of biogenic carbonate and quartz, K/Al and Mg/Al are

controlled by the mineralogy of the detrital clay fraction, although elevated Mg/Al can also

indicate the presence of dolomite. The EMS sediment samples have K/Al and Mg/Al

overlapping with previously published results (Figure 2), indicative of mixing between the four

principal clay mineral groups that were also recognized in the TGA results. Clay-rich samples

from sites 130 (NW of Nile fan) and 971B (mud volcano) plot on a mixing line between

smectite and illite at low Mg/Al. Higher Mg/Al in carbonate rich samples from the

Mediterranean ridge suggest a larger proportion of chlorite in these samples. The three

samples with the highest Mg/Al (excluding dolomite-rich sample AMS-038) are also

characterised by the most pronounced illite/chlorite dissociation peak in the TGA data.

The EMS sediment samples have REE patterns similar to average continental crust (Figure

3a) with chondrite normalised (La/Lu)N ratios between 7 and 11.5 and mild negative Eu-

anomalies (0.7-0.8). Absolute Nd concentrations vary between 14 and 47 ppm for all samples

except two evaporite samples that have significantly lower REE concentrations. In multi-

element variation diagrams (Figure 3b), the samples are characterised by a generally negative

smectite

kaolinite

chlorite

illite

0.1 0.2 0.3 0.4

0.2

0.4

0.6

0.8

1.0

1.2

K/Al

Mg

/Al

374

128

Libya

971B

378

130

Nile

375/6

125/6

literature

58 | CHAPTER 3

site

374

125

126

128

Libya

971B

378

130

Nilefan

376

375

sample

AMS-034

AMS-035

AMS-036

AMS-037

AMS-038

AMS-039

AMS-030

AMS-001

AMS-002

AMS-031

AMS-032

AMS-033

AMS-003

AMS-004

AMS-005

AMS-006

AMS-007

AMS-024

AMS-025

AMS-026

AMS-019

AMS-020

AMS-021

AMS-022

AMS-023

AMS-045

AMS-046

AMS-047

AMS-048

AMS-008

AMS-009

AMS-010

AMS-011

AMS-012

AMS-027

AMS-028

AMS-029

AMS-016

AMS-043

AMS-044

AMS-017

AMS-018

AMS-013

AMS-014

AMS-015

Th/Nb

0.53

0.59

0.48

0.52

0.51

0.67

0.60

0.49

0.51

0.46

0.66

0.25

0.65

0.60

0.61

0.59

0.58

0.55

0.62

0.54

0.30

0.27

0.27

0.27

0.27

0.80

0.93

0.70

0.64

0.28

0.47

0.24

0.27

0.35

0.40

0.42

0.47

0.41

0.41

0.48

0.55

0.54

0.30

0.53

0.35

La/Nb

1.77

1.88

1.65

2.15

1.65

1.98

1.88

1.71

1.85

1.75

2.01

1.20

1.95

2.18

1.92

1.86

1.89

2.28

1.78

1.87

1.20

1.16

1.16

1.14

1.15

2.22

2.40

2.17

2.01

1.19

1.71

1.10

1.25

1.51

1.77

1.56

1.71

1.50

1.55

1.67

1.76

2.02

1.07

1.64

1.44

Sc/Nb

0.72

0.78

1.43

1.29

0.62

0.88

1.07

0.80

0.90

1.01

0.90

0.65

1.21

1.10

1.40

1.30

1.43

0.95

0.66

1.05

0.54

0.61

0.68

0.69

0.66

1.42

1.28

1.33

1.28

0.62

1.16

0.57

0.74

0.66

1.18

0.84

0.88

1.30

0.88

0.98

1.65

2.22

0.61

1.59

0.85

Ni/Nb

1.81

2.41

2.50

3.80

2.77

3.49

4.68

4.30

5.22

3.41

3.08

2.81

6.43

6.34

6.52

8.37

8.08

3.06

0.52

3.86

2.08

3.50

3.72

3.41

3.67

12.94

12.82

23.24

17.23

2.98

5.69

1.81

2.21

3.25

2.79

2.66

3.54

7.16

10.52

7.15

15.78

41.53

5.20

20.98

8.27

Rb/Nb

4.94

5.60

4.47

5.32

4.75

7.41

5.56

4.31

4.72

4.20

6.88

1.62

6.50

3.75

6.06

6.18

6.10

3.88

2.99

4.02

2.18

1.99

1.90

1.85

1.74

7.66

7.70

6.44

6.54

2.22

3.78

1.67

2.09

2.65

2.50

2.95

3.46

4.05

3.70

3.98

5.53

5.09

2.31

5.60

3.09

87 86Sr/ Sr

0.710929

0.710439

0.709357

0.709167

0.710616

0.714608

0.710059

0.709397

0.709341

0.709292

0.713781

0.709405

0.709878

0.709261

0.709623

0.709428

0.709357

0.709396

0.709483

0.709550

0.709073

0.708842

0.708863

0.708834

0.708691

0.709300

0.709212

0.709233

0.709188

0.710234

0.709348

0.707970

0.708950

0.709110

0.709163

0.709350

0.709471

0.709165

0.709186

0.709138

0.708989

0.708819

0.708719

0.709093

0.708887

0.512101

0.512101

0.512125

0.512137

0.512066

0.512104

0.512134

0.512112

0.512180

0.512195

0.512107

0.512552

0.512146

0.512208

0.512151

0.512181

0.512200

0.512146

0.512119

0.512154

0.512390

0.512486

0.512485

0.512485

0.512508

0.512279

0.512293

0.512292

0.512299

0.512391

0.512211

0.512527

0.512510

0.512306

0.512343

0.512302

0.512218

0.512338

0.512331

0.512287

0.512349

n.d.

0.512460

0.512334

0.512425

143 144Nd/ Nd

0.282548

0.282584

0.282453

n.d.

0.282395

0.282572

0.282412

0.282409

0.282478

0.282472

0.282586

0.282891

0.282536

0.282266

0.282542

0.282614

0.282604

0.282282

0.282292

0.282415

0.282768

0.282789

0.282750

0.282763

0.282724

0.282661

0.282656

0.282687

0.282599

0.282846

0.282534

0.282872

0.282877

0.282682

0.282460

0.282420

0.282364

0.282704

0.282676

0.282645

0.282787

n.d.

0.282734

0.282723

0.282855

176 177Hf/ Hf

18.8556

18.8483

18.8563

18.8402

18.8259

18.8075

18.8811

18.8586

18.8892

18.8534

18.8041

18.9505

19.0009

19.0341

18.9889

18.9526

18.9301

18.6969

18.7124

18.7545

18.8577

18.9276

19.0492

18.8919

18.9057

18.8839

18.8605

18.9306

18.9250

19.1280

18.8944

19.0422

18.9999

18.9596

18.4380

18.4944

18.3769

18.9404

18.9505

18.9234

18.9091

n.d.

18.9383

18.9329

18.8718

206 204Pb/ Pb

15.6933

15.6933

15.6884

15.6861

15.6854

15.6874

15.6920

15.6893

15.6896

15.6836

15.6873

15.6656

15.7139

15.7079

15.7123

15.7015

15.6980

15.6755

15.6758

15.6807

15.6764

15.6740

15.6850

15.6706

15.6692

15.6905

15.6903

15.6953

15.6891

15.7242

15.6897

15.6883

15.6784

15.6926

15.6449

15.6559

15.6450

15.6904

15.6954

15.6918

15.6948

n.d.

15.6818

15.6955

15.6841

207 204Pb/ Pb

39.0081

38.9916

38.9853

38.9767

38.9520

38.9492

39.0024

39.0292

39.0073

38.9726

38.9427

38.9260

39.0904

39.0923

39.0798

39.0366

39.0265

38.7842

38.8122

38.8621

38.9203

38.9253

38.9233

38.9079

38.9435

38.9648

38.9829

39.0207

38.9894

39.1790

39.0113

38.9971

38.9381

39.0270

38.4389

38.4918

38.3409

38.9875

39.0196

39.0151

38.9953

n.d.

38.9351

39.0311

38.9593

208 204Pb/ Pb


Table 2. Selected trace element ratios and Sr–Nd–Hf–Pb data of the EMS sediment samples. Internal

errors (2 SE) for the isotope ratios are provided in the online supplementary material but are typically

±0.000009 for Sr, ±0.000008 for Nd, ±0.000012 for Hf, ±0.0009 for 206Pb/204Pb and 207Pb/204Pb, ±0.0020

for 208Pb/204Pb.

Ba-anomaly (Ba-content between 100 and 250ppm) and positive lead anomaly. Sample AMS-

027 from the Nile transect is extremely enriched in Ba (11,800 ppm) coupled with a positive

U-anomaly, probably due to barite precipitation in organic-rich sediment (Plank and Langmuir,

1998). Elevated uranium concentrations are recorded in five other samples, but without the

Ba-enrichment. Strontium content is highly variable (100-4,000 ppm), as are Zr-Hf

concentrations. All samples have a negative Nb-Ta anomaly, but the magnitude of the

anomaly varies by a factor 2.5.

Figure 3. Whole rock REE (a) and multi-element variation (b) diagrams showing the range of the EMS

sediment samples. The two evaporite samples have significantly lower trace element concentrations and

have therefore been plotted separately. GLObal Subducting Sediment II (GLOSS-II) is shown for reference

(Plank, 2013). Note the large variation in Ba, U, Sr and Zr-Hf concentrations for the EMS sediment

samples. Normalisation values for C1 chondrite from McDonough and Sun (1995) and primitive mantle

(PM) from Sun and McDonough (1989).

0.1

1

10

100

1000

sa

mp

le/C

1 c

ho

nd

rite

La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

evaporites

GLOSS-II

0.1

0.01

1

10

100

1000

10,000

sa

mp

le/P

M

CsRb Ba Th U Nb Ta La Ce Pb Sr NdSm Zr Hf Gd Y Yb

GLOSS-II

evaporites

a

b

60 | CHAPTER 3

4.3. Radiogenic isotopes

The 143Nd/144Nd isotope composition of the EMS samples varies between 0.5120 and 0.5126,

and 87Sr/86Sr ranges from 0.7079 to 0.7146, with 90% of the samples between 0.7087 and

0.7098. The range in 143Nd/144Nd reported here overlaps with previous studies in the EMS

(Freydier et al., 2001; Weldeab et al., 2002a) and is expanded to more radiogenic values. The

strontium isotope data cluster around Neogene seawater at 0.709 (e.g., McArthur et al.,

2012), in contrast with previous studies where samples were leached to remove the biogenic

carbonate component (Freydier et al., 2001; Weldeab et al., 2002a; Revel et al., 2010). The 176Hf/177Hf isotope composition of the sediments varies from 0.2822 to 0.2829 and is

positively correlated with 143Nd/144Nd but plots systematically to the high-εHf side of the

terrestrial array of Vervoort et al. (1999). The surface sediments of the African passive margin

are an exception as these samples have unradiogenic Hf compositions relative to their εNd.

The lead isotope composition of the EMS sediment samples is relatively restricted with the

majority of the samples at 206Pb/204Pb = 18.8-19.1, 207Pb/204Pb = 15.66-15.72 and 206Pb/204Pb

= 38.9-39.15. The surface sediments are again the exception as they have a much less

radiogenic Pb-isotope composition (see Figure 8).

Figure 4. Triangular diagram showing the relative concentrations of Al2O3, SiO2 and CaCO3 in the EMS

sediment samples. The majority of the samples plot on a dilution trend away from the detrital clay fraction

by biogenic carbonate. Only the turbidite sample from site 128 (AMS-004) shows significant dilution with

silica. The evaporite samples from site 375/376 are dominated by gypsum and contain little Al2O3, SiO2

and carbonate.

carbonate dilution

detritalclay

evaporites

Al O × 52 3

SiO2 CaCO3

374

128

Libya

971B

378

130

Nile

375/6

125/6

W

E

silic

a d

ilutio

n

turbidite


5. DISCUSSION

5.1. Mineralogical composition of EMS sediment

Deep sea sediment is typically a mixture of several lithological components: continental

detrital phases, hydrothermal or hydrogenous Mn-Fe deposits and biogenic carbonate or opal

(e.g., Dymond, 1981; Plank and Langmuir, 1998; Plank, 2013). Major element composition

and thermogravimetric dissociation curves can be used to infer the lithological and

mineralogical composition of the EMS sediment samples. Figure 4 shows a triangular diagram

of the relative proportions of Al2O3, SiO2 and CaCO3 in which detrital clay, biogenic or detrital

silica and biogenic carbonate form distinct end members. EMS sediment samples are

predominantly two-component mixtures of terrigenous detrital clay and carbonate.

Mudstone samples such as the black clays from site 130 (NW of Nile fan) and mud debris

flow deposits from site 971B (mud volcano) are dominated by the detrital clay fraction, while

pelagic marls and oozes from the Mediterranean Ridge (e.g. sites 125/126 and 374) are

diluted with up to 60% biogenic carbonate. Dilution with silica, either as biogenic opal or

detrital quartz grains, is not observed except for a turbidite sample from site 128 (Hellenic

Trench). Thermogravimetric dissociation curves of the Messinian evaporites from site 375/376

(Cyprus) indicate that these samples consist of >90 % gypsum with low carbonate and clay

mineral content.

5.2. Controls on trace element variation

5.2.1. First-order control: carbonate dilution

Trace elements with a short residence time in seawater adhere to clay minerals and hence

their concentrations are governed by the proportion of clay minerals in marine sediments

(e.g., Plank and Langmuir, 1998; Plank, 2013). As the major element variation of EMS

sediment indicates dilution of a terrigenous clay fraction by biogenic carbonate, these detrital

trace elements are expected to correlate with the Al2O3 content. Three trace element groups

can be distinguished based on their correlation with aluminium content. Representative

examples of each group (Nb, Th, La, Rb, Zr, Sr) are shown in Figure 5. Group I contains trace

elements that show a strong (R2 > 0.80) positive correlation with Al2O3 and are therefore

detrital in origin. This group includes Li, Be, Sc, Zn, Ga, Y, Nb, In, Sn, REE, Hf, Ta and Th.

Several trace elements that are typically detrital (Plank and Langmuir, 1998; Carpentier et al.,

2009) show increased scatter when plotted against Al2O3, suggesting a second-order control

on their distribution in EMS sediment. These trace elements constitute group II and include

Cr, Co, Ni, Cu, Rb, Zr, Cs and Pb. Possible causes of the anomalous behaviour of group II

elements are discussed below. Group III trace elements do not vary systematically with the

proportion of clay minerals or show a negative correlation. Strontium concentration is strongly

correlated with the CaCO3 content of the EMS samples and ranges from 100 ppm in clay-rich

samples to 2,000 ppm in the most carbonate-rich samples. Due to the solubility of Sr in

seawater and its incorporation in carbonate, the detrital Sr isotope signal in these sediments

62 | CHAPTER 3

Figure 5. Whole rock Al2O3 versus representative group I, II and III trace elements that illustrate the

distinct behaviour of each group. The concentration of group I trace elements (a, b and c) is controlled

by dilution of the detrital clay fraction with biogenic carbonate. Group II elements (d and e) show

increased scatter due to second-order controls on the trace element distribution. The group III trace

element strontium (f) is enriched through biogenic carbonate addition. Note the break in scale and the

very high Sr content of sample AMS-027. Dolomite-rich samples have significantly lower Sr-content at a

given Al2O3. See text for further discussion. GLOSS-II (Plank, 2013) is shown for reference.

a d

b e

c f

GLOSS-II

dolomite-rich

high Sr+Ba

carbonate addition

374 128 Libya 971B 378 130 Nile

W E

125/6 375/6

high-Zr group

dilutio

n of t

he detri

tal

clay f

racti

on

dilutio

n of t

he detri

tal

clay f

racti

on

dilutio

n of t

he detri

tal

clay f

racti

on

dilution of the detrital clayfraction: increased scatter

dilution of the detrital clayfraction: increased scatter

Group I element

Group I element

Group I element

Group II element

Group II element

Group III element

500

1500

10

20

30

40

0

4

8

12

0

0

20

40

0

40

80

120

0

100

200

300

0

1000

2000

4000

10

30

50

0 20105 15 0 20105 15

Al O (wt %)2 3 Al O (wt %)2 3

Nb

(p

pm

)T

h (

pp

m)

La

(p

pm

)

Rb

(pp

m)

Zr (p

pm

)S

r (pp

m)


has suffered a large overprint: 87Sr/86Sr does not reflect the source of the sediments but cluster

around the Neogene seawater value at 0.709 (e.g., McArthur et al., 2012). One anomalous

sample has a Sr-content of almost 4,000 ppm, much higher than the average 2,000 ppm in

calcite. This positive Sr-anomaly is coupled with a very high Ba-concentration (11,800 ppm),

suggesting that Sr and Ba may be precipitated in barite through high organic productivity

(Plank and Langmuir, 1998). Other elements in group III are Rh, which correlates well with Sr

and CaCO3, and U that is controlled by the redox state (Klinkhammer and Palmer, 1991).

Black clay samples from site 130 (AMS-008 and 012) have the highest U-concentrations.

5.2.2. Second-order control: zircon effect

Zircon is the main carrier of Zr and Hf in igneous rocks and as it is highly resistant to chemical

and physical weathering, zircon is also likely to dominate the Zr and Hf budget in sedimentary

rocks. However, this heavy accessory mineral tends to be concentrated in coarse sediments

deposited near the continental margin. Pelagic sediments are therefore generally depleted in

Zr and Hf compared to the continental crust (Patchett et al., 1984). Figure 5c shows the

relationship between Zr and Al2O3 for the EMS sediment samples. The majority of the samples

plot on a typical carbonate dilution trend as seen for Th, Nb and La. In contrast, a group of

samples comprising the surface sediment samples from the African passive margin and a

carbonate turbidite from site 128 (Hellenic trench) has a Zr content that is roughly two times

higher than expected from their Al2O3. The higher Zr content suggests that these coarser

sediments are enriched in zircon compared to the pelagic samples.

The absence or presence of zircon in sedimentary rocks will also influence their Hf

isotope systematics. Due to its high Hf content and very low Lu/Hf ratio, zircon has an

unradiogenic Hf isotope composition compared to its host rock (e.g., Kinny and Maas, 2003).

The presence of detrital zircon in coarser sediments therefore results in a negative shift in εHf

away from the terrestrial Nd/Hf isotope array (Vervoort et al., 1999). This is the well-known

“zircon effect” in mature sediments (e.g., Patchett et al., 1984; Vervoort et al., 1999; Chauvel

et al., 2008; Carpentier et al., 2009). Pelagic clay is complementary to zircon-rich sediments:

a generally lower absolute Hf-content but higher Lu/Hf ratio lead to a more radiogenic Hf

isotope composition (e.g., Chauvel et al., 2008; Vervoort et al., 2011). As a result, most

pelagic marine sediments have εHf higher than the terrestrial array at a given εNd (e.g.,

Vlastélic et al., 2005; Bayon et al., 2009; Carpentier et al., 2009), which is also the case for

most of the EMS sediment samples (Figure 6a). The influence of zircon on 176Hf/177Hf is

illustrated by comparing the deviation of εHf from the terrestrial array (ΔεHf; Carpentier et

al., 2009) with Zr/Nb (Figure 6b). The main group of samples has significantly lower Zr/Nb (7-

11) than upper continental crust (UCC, 15.8; McLennan, 2001) and generally positive ΔεHf

(-3 to +6), in line with the zircon deficiency expected for pelagic sediments. A negative

correlation is present between ΔεHf and Zr/Nb for the main group samples, suggesting that

the abundance of zircon is responsible for the variation in ΔεHf. The high-Zr samples have

high Zr/Nb and low ΔεHf (up to -10) on a continuation of the curved trend defined by the

64 | CHAPTER 3

Figure 6. The zircon effect in the EMS sediment samples: a) 143Nd/144Nd vs. 176Hf/177Hf diagram. The

majority of the samples plot towards the high 176Hf/177Hf side of the terrestrial array (Vervoort et al.,

1999), with the exception of the high-Zr samples from the African continental margin; b) ΔεHf (deviation

from the terrestrial array; Carpentier et al., 2009) vs. Zr/Nb. The high-Zr samples have high Zr/Nb and low

ΔεHf, consistent with the addition of zircon. See text for further discussion.

main group, which is consistent with the presence of detrital zircon in these coarse samples.

Surprisingly, the zircon effect trend crosses ΔεHf = 0 at a lower Zr/Nb (~9) than upper

continental crust (15.8; McLennan, 2001) and GLOSS-II (13.7; Plank, 2013). This suggests a

systematically higher Nb content for the EMS sediment samples compared to average crust

or sediment. The zircon effect can account for the scattered trends against Al2O3 for elements

such as Zr and Hf and to a lesser extent for elements present in other heavy accessory phases

(e.g., Nb in ilmenite). However, the coarse high-Zr samples are not the cause of the scatter in

many of the trace element vs. Al2O3 diagrams (Figure 5a&c) and another factor appears to be

of influence. We argue that the excess scatter might be controlled by variation in the

provenance of the EMS sediments.

5.3. Provenance of EMS sediment

5.3.1. Provenance of EMS sediment derived from trace element variations

With the exception of Zr and Hf that are affected by the zircon effect, the concentrations of

group I and II trace elements are controlled by the detrital clay fraction and are therefore

terrigenous in origin. As these trace elements tend not to fractionate during weathering and

sedimentation, trace element ratios reflect the composition of the source region of sediments

(e.g., McLennan et al., 1990). Therefore, ratios of detrital trace elements can be used to

identify provenance areas of deep sea sediments (e.g., Vroon et al., 1995). In addition, the

use of trace element ratios instead of concentrations eliminates the effect of dilution of the

detrital component by biogenic carbonate, evaporite or silica. We use the Th/Nb, La/Nb,


Figure 7. Whole rock trace element discrimination diagrams that allow the distinction of provenance

components in EMS sediment samples. Compiled literature data for potential provenance areas (median

and 50% spread) are shown for reference; see text for further discussion and references. A – Aegean

volcanic rocks, N – Nile sediment, SCB – Sahara dust from Chad Basin region, SHM – Sahara dust from

Hoggar Massif region, O – Tethyan ophiolites. Shaded field indicates mixing array between Nile sediment

(N) and Sahara dust (SH). GLObal Subducting Sediment II (GLOSS-II) from Plank (2013) and Upper

Continental Crust (UCC) from McLennan (2001).

Sc/Nb, Ni/Nb and Rb/Nb ratios of the EMS sediment samples to distinguish provenance areas

in Figure 7. These trace elements were selected from the total dataset because their ratios

over niobium show the largest variation between provenance components. Other group I and

II trace elements either mirror their geochemical twins, such as Rb-Cs, Nb-Ta, Zr-Hf, Ni-Co-

Cr, or do not vary significantly between provenance areas (e.g. Li, Be, Ga). Trace element

compositions for potential source areas of EMS sediment have been compiled from the

literature. Direct analyses of sediments were only available for Sahara dust, the other source

O

O

O

O

A

A

A

A

N

N

NN

SCB

SCB

SCB

SCB

SHM SHM

SHM

SHM

UCC

0 0.25 0.50 0.75 1.0 1.25 0 0.25 0.50 0.75 1.0 1.25 1.5

Th/Nb Th/Nb

0

2

4

6

8

0.1

1

10

100

Ni/N

bR

b/N

b

0

0.5

1.0

1.5

2.0

2.5

0.5

1.0

1.5

2.0

2.5

3.0S

c/N

bL

a/N

ba

b d

c

374

128

Libya

971B

378

130

Nile

375/6

125/6W

E

GLOSS-II

66 | CHAPTER 3

areas have been estimated from bedrock data. Sahara dust samples (Moreno et al., 2006;

Castillo et al., 2008) together with sediment from ODP site 658 off the coast of Mauritania

(Cole et al., 2009) show a bimodal composition and have therefore been separated in two

distinct source areas: dust from the Hoggar Massif and Western Sahara (SHM in Figure 7) and

dust from the Chad Basin (SCB in Figure 7). The sediment discharge of the Nile consists for

~95% of detritus from the Ethiopian flood basalts (Padoan et al., 2011), the composition of

which is taken to represent Nile sediment (N in Figure 7; Stewart and Rogers, 1996; Pik et al.,

1999). Geochemical data of Tethyan ophiolites (O in Figure 7) are taken from Dilek et al.

(2008) and Dilek and Furnes (2009). The composition of Aegean volcanic rocks (A in Figure

7) is compiled from Elburg et al. (2014). As discussed in the introduction, there is general

consensus that EMS sediment is predominantly a mixture between Sahara dust and Nile

sediment (Venkatarathnam and Ryan, 1971; Krom et al., 1999b; Weldeab et al., 2002a). This

is supported by the trace element discrimination diagrams presented in Figure 7 as the

majority of the samples cluster around a mixing line between the compositions of aeolian

dust from the northern part of the Sahara (SHM) and Nile sediment (N) that have been

compiled from the literature. The samples that plot closest to the Nile component at Th/Nb ≈

0.3 are smectite-rich black clays from site 130 (NW Nile fan) and samples from the mud

volcano (site 971B). In contrast, carbonate-rich samples from the Mediterranean Ridge (site

125) with kaolinite and illite as dominant clay minerals overlap perfectly with the northern

Sahara component. However, binary mixing between these components alone is not

sufficient to account for the observed trace element variation. Therefore, we propose that

two additional provenance components can be resolved in our EMS sediment dataset. A

component with low Nb content and higher (Th, La, Sc, Rb, Ni)/Nb ratios than SH dominates

the samples from the Aegean Sea (site 378). The overlap with the field for Aegean volcanic

rocks (A in Figure 7) suggests that this component represents subduction related volcanic

rocks from the Aegean area, either as ash layers or, more likely, detritus of these volcanic

rocks. In addition to arc volcanic rocks, the geology of the Aegean region comprises trace-

element poor Cenozoic limestones and Tethyan ophiolite complexes. The former will not

significantly contribute to the trace element budget of Aegean sediment, but addition of

detritus derived from (ultra)mafics is a plausible origin for the high Ni/Nb ratios in site 378

sediments. Many samples from other EMS sites also display higher Sc/Nb and Ni/Nb ratios

compared to the Sahara-Nile mixing array. As these high ratios of compatible element over

Nb are most pronounced in samples from sites 375 and 376 near Cyprus, we suggest that

this high Ni/Nb and Sc/Nb component represents detritus from the Tethyan ophiolite belt (O

in Figure 7). This Tethyan ophiolite component is not restricted to the Troodos ophiolite on

Cyprus but might also originate from the belt of (ultra)mafic rocks extending from the

Dinarides through Greece and Turkey.

The composition of EMS sediment can thus be explained by mixing between four

components that are summarised in Table 3. Nile sediment and Sahara dust are the main

contributors, but a Tethyan ophiolite component from the northern margin of the EMS can


also be recognised in many samples. The Aegean volcanic rocks component dominates in the

Aegean Sea site 378, but in the EMS only samples from the Hellenic trench (site 128) show

elevated (Th, La, Sc, Rb, Ni)/Nb ratios (see also Figure 8c). A resolvable Aegean volcanic rocks

component in these trench fill sediments is supported by the presence of glaucophane, which

suggests a northerly provenance (Ryan et al., 1973a). Several samples from site 374 in the

Ionian basin have lower Sc/Nb and Ni/Nb compared to the Sahara-Nile mixing array and

slightly more negative εNd-values (Figure 8c). This provides tentative evidence for a fifth,

Adriatic, component that could represent detritus from the Alps as suggested by Weldeab et

al. (2002a). The present dataset and lack of suitable literature data, however, does not allow

a clear distinction between a possible Adriatic component and Sahara dust. Although the

Chad Basin (SCB-component in Figure 7) is one of the largest dust sources in the Sahara

(Prospero et al., 2002) it is not resolved in our trace element dataset, which confirms that

dust is predominantly transported to the W.

5.3.2. Isotope composition of the provenance components

Multiple studies have investigated the Sr-Nd isotopic variation in EMS sediment, but hafnium

and lead isotope data have not been reported previously. Strontium isotope variation in the

EMS sediment samples is controlled by biogenic carbonate and all samples cluster around

Neogene seawater values of ~0.709 (e.g., McArthur et al., 2012). Only four clay- or dolomite-

rich samples have a Sr isotope composition that is clearly distinct from seawater and reflects

a Nile sediment or Sahara dust influence. In contrast, hafnium isotope ratios are controlled by

the zircon effect described above. Hence, the discussion of provenance below will focus on

Nd and Pb isotopes.

Neodymium isotope evidence corroborates the distinction of four provenance

components. The Nile sediment component, which is derived from the Ethiopian flood

basalts, has a relatively radiogenic Nd-composition that is in good agreement with published

data (see Figure 8). Sahara dust has lower 143Nd/144Nd, consistent with derivation from older,

felsic terrains. The majority of the samples plot on a mixing array between these two

components, but the Aegean samples (site 378) are displaced towards higher Th/Nb at a 143Nd/144Nd of ~0.5123 (Figure 8c). The Tethyan ophiolite component cannot be resolved in

the Nd-isotope data, but its 143Nd/144Nd is likely to be close to typical MORB values.

Lead isotope data of the EMS sediment samples are not solely governed by provenance.

The samples cluster around two mixing arrays in a 207Pb/204Pb vs. 206Pb/204Pb diagram (Figure

8b). In particular the samples that have a predominant Sahara dust composition according to

trace element and Nd data define a linear trend in which the passive margin sediments have

the least radiogenic composition. These linear trends in Pb-isotopes have been recognised

before in marine sediments and have been interpreted as mixing with a radiogenic dissolved

Pb component (Vlastélic et al., 2005). Such a dissolved component probably originates from

the preferential release of radiogenic lead from U- and Th-rich accessory minerals during

weathering (Blanckenburg and Nägler, 2001; Harlavan and Erel, 2002). Hence, marine

68 | CHAPTER 3

sediment from a single source area can display large Pb-isotope variations due to mixing

between unradiogenic detrital minerals and radiogenic fluids. We propose that this process is

also responsible for the variation in 206Pb/204Pb of the EMS sediment samples. The lead isotope

variation can thus be explained as three component mixing between a dissolved radiogenic

(DR in Figure 8) component at 206Pb/204Pb = 19.20 and 207Pb/204Pb = 15.72, a high 207Pb/204Pb

component representing Sahara dust and detritus from the Ethiopian flood basalts at lower 207Pb/204Pb. The passive margin samples (Libya and Nile) have lower 206Pb/204Pb than literature

Figure 8. Whole rock isotope data for the EMS sediment samples. The majority of the samples are located

close to mixing curves between Sahara dust (SD) and Nile sediment (EFB), but lead isotopes require the

contribution of a dissolved radiogenic (DR) component; see text for further discussion. 207Pb/204Pb (18.8)

in (d) is the 207Pb/204Pb extrapolated from the DR component at 206Pb/204Pb = 18.8; see inset of (d). Isotope

data for Ethiopian flood basalts (EFB – Nile component) from Pik et al. (1999) and Sahara dust (SD) from

Biscaye et al. (1997), Grousset et al. (1998) and Revel et al. (2010).

DREFB

SD

to EFB

Sahara S

hields?

mixing

18.4 19.218.6 18.8 19.018.418.2 18.6 18.8 19.0 19.2

206 204Pb/ Pb

206 204Pb/ Pb

0.5120

0.5122

0.5124

0.5126

143

144

Nd

/N

d

15.74

15.72

15.70

15.68

15.66

15.64

207

204

Pb

/P

b

0.5120

0.5122

0.5124

0.5126

143

144

Nd

/N

d

0 0.40.2 0.6 0.8 1.0 0 0.40.2 0.6 0.8 1.0

15.75

15.70

15.65

15.60

15.55

207

20

4P

b/

Pb

(18.8

)

EFB

SD

Th/Nb Th/Nb

374 128 Libya 971B 378 130 Nile

W E

125/6 375/6

EFB

SD

a

c

b

d

DR18.8

measured

207 /204

Pb Pb (18.8)

206 /204

Pb Pb

207

/204

Pb

Pb

19.018.0

15.6

15.7

15.5

0.5120

0.5124

0.5128

0.2 0.6


data for Sahara dust (Figure 8b), which suggests that these samples contain atypical amounts

of unradiogenic detrital phases (e.g. feldspar). As the Nile and Sahara components are

themselves mixing arrays, 206Pb/204Pb and 207Pb/204Pb of the EMS sediments cannot be used

directly to distinguish mixing between these components. Instead, we introduce a normalised 207Pb/204Pb (18.8) parameter, which is the extrapolated 207Pb/204Pb of a sample from the dis-

solved radiogenic composition to 206Pb/204Pb = 18.8 (Figure 8d inset), to eliminate the

dissolved radiogenic component. This 207Pb/204Pb (18.8) ratio correlates well with Th/Nb and

confirms predominant mixing between Sahara dust and Nile sediment.

5.4. Spatial and temporal variation in provenance

With the recognition of provenance components on the basis of trace elements and isotopes,

it is possible to track spatial and temporal variations in provenance of the EMS sediment

samples. Figure 9 depicts the variation of two key provenance tracers, Th/Nb and εNd, in

space and time. Pliocene and Quaternary EMS sediment samples are characterised by a large

variation in dominant provenance component and a compositional E-W gradient. Nile

sediment entering the EMS in the east is progressively diluted by Sahara dust towards the

west, consistent with previous studies (Venkatarathnam and Ryan, 1971; Krom et al., 1999b;

Weldeab et al., 2002a). Significant geochemical variation is observed within individual sites.

In particular, sediment at site 130 (NW of Nile fan) is highly heterogeneous and alternates

between Sahara dust dominated, carbonate-rich pelagic marls and smectite-rich turbidites

with a strong Nile provenance. Here, the within-site variation in provenance is thus

predominantly controlled by the depositional mechanism, wherein Nile sediment turbidites

are interspersed with periods of slow, pelagic sedimentation with a higher Sahara dust

contribution. A likely secondary control on Pliocene-Quaternary within-site heterogeneity is

the same orbitally induced variations in monsoon intensity that are responsible for the

Holocene oscillations and sapropel formation (e.g., Revel et al., 2010). The large variability in

Pliocene-Quaternary EMS sediment geochemistry and provenance is absent from the pre-

Messinian samples. Miocene EMS sediment from sites 125/6 and 971B is characterised by a

restricted range in εNd and Th/Nb that is similar to Quaternary clays from the Nile fan. The

debris flows from site 971B mud volcano, which comprise well-mixed sediment of Eocene to

Miocene age, form an average of pre-Messinian EMS sediment that is close to the pure Nile

component in composition. This suggests that the contribution of aeolian dust to EMS

sediment was limited until the late Miocene and that the present-day provenance gradient is

a relatively recent feature (Figure 10). The increase in Sahara dust contribution since the

Messinian salinity crisis coincides with rapid global climate change from a warm and wet

“greenhouse” in the Paleogene to a colder climate characterised by glacial-interglacial cycles

in the Pliocene and Quaternary (e.g., Fauquette et al., 2007; Micheels et al., 2009). The

warmer Miocene climate affected sedimentation in the EMS in two ways. Firstly, the Sahara

region was covered with grassland vegetation up to the Middle Miocene (Jacobs, 2004),

leading to little dust production. The Sahara desert did not emerge until the late Miocene,

70 | CHAPTER 3

with the oldest aeolian deposits dated at 5-7 Ma (Vignaud et al., 2002; Schuster et al., 2006).

Secondly, higher precipitation over North Africa during the Middle Miocene led to increased

fluvial runoff into the EMS (Gladstone et al., 2007). Although the higher fluvial flux was largely

accommodated by the Nile, it has also been postulated that Lake Chad drained to the north

through the Sahabi River (Griffin, 2006; Figure 10). Our trace element discrimination data

(Figure 7) neither support nor rule out a Lake Chad component in the pre-Messinian samples

as literature data roughly overlap with Nile sediment. In addition to an increased Nile sediment

flux and lower dust production, the Miocene paleogeography also favoured the absence of

an east-west provenance gradient. Numerical models suggest that enhanced deep circulation

in the Miocene EMS due to a larger surface area of the Adriatic Sea was effective in

transporting and distributing terrigenous Nile sediment throughout the EMS (Meijer et al.,

2004). To what extend the permanent closure of the Mediterranean-Indian Ocean gateway

at ca. 14 Ma (Rögl, 1999) affected sedimentation patterns in the EMS is uncertain. Even

though this event has had a large impact on oceanic circulation, the amount of sediment

delivered by the Nile has been orders of magnitude larger than any pelagic sediment

Figure 9. Spatial and temporal variation in EMS sediment provenance as derived from variation in Th/Nb

(a) and εNd (b). Pliocene and Quaternary samples display the largest variation and a distinct gradient of

Nile sediment progressively diluted by Sahara dust to the west is evident. In contrast, Miocene samples

below the Messinian salinity crisis (MSC) evaporites have a more restricted geochemical composition that

is dominated by Nile sediment. The samples from site 971B, the Napoli mud volcano, represent well-

mixed debris flows of Eocene-Miocene sediment and hence represent a best average composition of pre-

Messinian EMS sediment.


Figure 10. Sediment fluxes into the EMS during the Quaternary/Pliocene (a) and Upper Miocene (ca. 12

Ma; b). The Quaternary/Pliocene is characterised by dilution of Nile sediment with Sahara dust towards

the west, leading to an E-W provenance-induced geochemical gradient in EMS sediment. During the

Miocene, input of Nile sediment was higher while dust production was much lower compared to the

present-day, leading to a more homogeneous sediment composition.

transported through the gateway from the East. The average Eocene-Miocene sediments

from site 971B indicate that Nile detritus was dominant in that period and no Indian Ocean

component can be distinguished. Pre-Messinian sediment from site 375/6 is more variable

and shows a higher Sahara dust contribution (Figure 9). This is possibly due to the location of

this site on a ridge, which favours a relatively higher contribution of Sahara dust as Nile

sediment is predominantly transported by deep bottom currents.

5.5. EMS sediment: implications for arc volcanism

5.5.1. Aegean arc

The shift of the active deformation front of the Aegean subduction system from the Hellenic

trench to the south in the Miocene resulted in ongoing accretion of the EMS sediment pile to

the Hellenide complex. Despite the formation of the Mediterranean Ridge accretionary prism,

Cyprus

Libya Egypt

Levant

200 km

TurkeyQuaternary/Pliocene

Sahara dust Nile sediment

Libya

Egypt

Levant

200 km

TurkeyU. Miocene

Sahara dust

Nile sediment

Sahabi river? Nile

72 | CHAPTER 3

40-80% of the sediment is subducted rather than accreted (Kopf et al., 2003). Along-arc

variation in subducted sediment composition has been proposed to explain the geochemical

heterogeneity in Aegean arc magmas (Francalanci et al., 2005). In particular, an increased

input of Nile-derived sediment in the eastern section of the arc is evoked to explain the

relatively unradiogenic Pb-isotope composition of Nisyros volcanic rocks (Elburg et al., 2014).

With the new EMS sediment dataset, it is possible to better quantify the potential contribution

of subducted sediment to Aegean arc lavas. A simple bulk mixing model involving a depleted

MORB mantle source (DMM) with EMS sediment is presented in order to draw some general

conclusions (Figure 11). In this model, we use samples AMS-001 and AMS-011 as

representative examples of Sahara dust and Nile component dominated sediment,

respectively, based on their trace element and isotope composition. The sediment-DMM

mixing lines are strongly dependent on the Pb isotope composition of the EMS sediments due

to the low Pb content of DMM (0.018 ppm; Workman and Hart, 2005). Therefore, variation

in trace element contents and isotope composition of the samples and the choice of DMM

composition is of little influence to the conclusions drawn from the model.

Several lines of reasoning argue against subducted sediment heterogeneity as the

dominant control on the variation in Aegean arc lavas. First, we have argued that the east-

west gradient in EMS sediment provenance and composition is a recent feature that

developed in the latest Miocene due to progressive aridification of the Sahara region and that

pre-Messinian sediment has an on average Nile composition. Given the geometry of the

Aegean subduction zone and a constant subduction rate of 5-6 cm/yr, the minimum age of

sediment currently underneath the volcanic arc is ca. 5 Ma. Hence, an east-west gradient is

probably lacking in the sediments that potentially contribute to the present-day volcanism in

the Aegean arc. Second, the dataset shows that the range in Pb-isotope composition of EMS

sediment, excluding the unradiogenic African passive margin samples, is relatively small

compared to the variation observed in Aegean arc volcanic rocks (Figure 8a, b). Moreover,

the Nile component has higher 206Pb/204Pb than Sahara dust due to preferential leaching of

radiogenic lead (see section 5.3.2.). An increased input of Nile sediment would therefore lead

to a more radiogenic Pb-isotope composition of the volcanic rocks instead of the anomalously

low 206Pb/204Pb seen in Nisyros volcanic rocks.

Figure 11 shows a simple binary mixing model between DMM and Sahara dust and Nile

sediment in comparison with three volcanic centres from the Aegean arc. The Pb-isotope

composition of EMS sediment largely overlaps with that of Methana and Santorini lavas. Bulk

addition of 0.5 to 5% of EMS sediment to a DMM source can explain the Nd- and Pb-isotope

variation in Santorini lavas. It should be noted that the 0.5-5% sediment fraction is likely to

be an overestimate as we have modelled bulk mixing of sediment with DMM. In contrast,

transfer of these elements from the slab to the mantle wedge occurs through fluids or melts

enriched in Nd and Pb (e.g., Kessel et al., 2005). Methana, the western volcanic centre in the

Saronic Gulf, shows a much larger variation in Nd-Pb isotopes compared to Santorini. Part of

the range can be explained by bulk addition of >5% EMS sediment, but contamination of


magmas by arc crust is probably also required to account for the total variation in Methana.

Nisyros lavas have significantly lower 206Pb/204Pb compared to the western and central Aegean

arc and their variation falls well outside the mixing models in Figure 11. Therefore, we propose

that the unradiogenic composition of Nisyros volcanic rocks is not the result of increased Nile

sediment input, but that a distinct mantle source with low 206Pb/204Pb and 207Pb/204Pb is

required.

Figure 11. Bulk mixing model of a depleted mid-ocean ridge basalt mantle source (DMM) with EMS

sediment in comparison with Aegean arc volcanic rocks from Nisyros (eastern), Santorini (central) and

Methana (western Aegean arc). The addition 0.5-5% EMS sediment can account for the isotopic variation

in lavas from Santorini and part of Methana, but Nisyros requires a distinct mantle source. See text for

further discussion. DMM Pb and Nd concentrations are 0.018 and 0.581 ppm respectively (Workman and

Hart, 2005); sample AMS-001 is taken as representative sample for Sahara dust, AMS-011 for Nile

sediment. Data fields of Aegean volcanic rocks from Elburg et al. (2014); symbols for EMS sediment as in

previous figures.

DMM

Nisyros (E)

Santorini (C)

Methana (W)

0.5%

1%2%

DMM

0.5%

1%

2%

5%

18.4 18.6 18.8 19.0 19.20.5120

0.5128

0.5124

0.5132

15.5

15.6

15.7

15.8

206 204Pb/ Pb

143

144

Nd

/N

d207

204

Pb

/P

b

a

b

Aegean Arc

74 | CHAPTER 3

Figure 12. Bulk mixing model of a HIMU-OIB mantle source with EMS sediment compared to Aeolian

arc and Etna volcanic rocks. Of the Aeolian volcanoes, only the composition of Stromboli volcanic rocks

can be explained by the addition of EMS sediment to the mantle wedge. The other volcanic centres have

a more radiogenic Pb-isotope composition, more compatible with crustal assimilation (CA) (Peccerillo et

al., 2013). See text for further discussion. Data from Peccerillo (2005) (Aeolian, Stromboli), Tommasini et

al. (2007) (Stromboli) and Viccaro and Cristofolini (2008) (Etna). Symbols for EMS sediment as in previous

figures.

5.5.2. Aeolian arc

Subduction of the Ionian plate beneath Calabria has led to the formation of the Aeolian arc

north of Sicily (Figure 1) and hence EMS sediment is a potential component of Aeolian volcanic

rocks. The isotopic composition of Aeolian lavas is strikingly different from the Aegean arc

with high 206Pb/204Pb (19.0-19.8; Figure 12). Controversy remains about the source of this

radiogenic Pb-isotope composition in the Aeolian volcanic rocks and both old hydrothermally

altered oceanic crust (e.g., Francalanci et al., 2007; Tommasini et al., 2007) and an Etna-like

HIMU-OIB (high 238U/204Pb ocean island basalt) mantle source (e.g., Ellam et al., 1988;

Peccerillo et al., 2013) have been evoked. Multiple studies suggest an increase in subducted

sediment contribution from the western islands (Alicudi, Filicudi) to Stromboli in the east,

18.8 19.2 19.6 20.0

206 204Pb/ Pb

0.5120

0.5128

0.5124

0.704

0.708

0.712

14

31

44

Nd

/N

d87

86

Sr/

Sr

Etna

Stromboli

Aeolian Arc

HIMU-OIB

HIMU-OIB

a

b

0.5%

1%

2%

5%

0.5%

1%

2%

5%

CA

Aeolian Arc


based on a west-to-east decrease in 206Pb/204Pb, increase in 87Sr/86Sr (e.g., Peccerillo et al.,

2013) and increasing B content and δ11B (Tonarini et al., 2001). Figure 12 shows a simple bulk

mixing model of EMS sediment with an HIMU-OIB, Etna-like (Viccaro and Cristofolini, 2008)

mantle source. Bulk addition of 2-20% of Nile sediment can account for the Sr-Nd-Pb

composition of the least Pb-radiogenic Stromboli volcanic rocks, which is in line with

calculations by Peccerillo et al. (2013). The other Aeolian centres fall outside the mixing model

but plot on sublinear mixing arrays with the Calabrian basement (Peccerillo et al., 2013).

Hence, it is possible that the Sr-Nd-Pb isotope variation in these volcanic rocks is controlled

by crustal processes rather than variable contribution of subducted sediment. If, in contrast,

a mantle wedge with typical MORB composition (206Pb/204Pb 18.5) is assumed, bulk mixing

with EMS sediment fails to reproduce the radiogenic Pb-isotope composition of the Aeolian

volcanic rocks and a distinct high-206Pb/204Pb component is required, possibly altered oceanic

crust (Francalanci et al., 2007; Tommasini et al., 2007). Combining the trace element

arguments presented in Peccerillo et al. (2013) with the EMS sediment mixing model (Figure

12) supports a HIMU-OIB-like Aeolian mantle wedge and increasing sediment input from west

to east. The relatively unradiogenic Pb-isotope composition of Stromboli is the result of

effective overprinting of the high-206Pb/204Pb mantle by EMS sediment. Isotopic variations in

Etna volcanic rocks are also in accordance with the bulk addition of 0.1 % EMS sediment to

explain the large 206Pb/204Pb variation at relatively constant 87Sr/86Sr (Figure 12a).

6. CONCLUSIONS

The new Eastern Mediterranean Sea sediment dataset presented here corroborates the

suggestion that EMS sediment is predominantly a mixture between Sahara dust and Nile

sediment. However, trace element data indicate the involvement of at least two additional

provenance components. An (ultra)mafic component derived from the Tethyan ophiolite belt

in the northern borderlands of the EMS is recognised in most samples. Low-Nb sediment from

the Cenozoic Aegean arc volcanic rocks is a subordinate component that is only evident in

sediment from the Hellenic Trench. Our approach demonstrates the usefulness of a combined

trace element-isotope study in that provenance components with overlapping isotope

composition can be clearly discriminated on the basis of distinct trace element fingerprints.

These trace element fingerprints have a large potential to be useful in palaeoclimatology

studies investigating provenance fluctuations in the EMS due to their ease-of-analysis

compared to isotope data. Pliocene and Quaternary EMS sediment records a strong

provenance-induced geochemical gradient where Nile sediment entering the EMS in the east

is progressively diluted by Sahara dust towards the west. In contrast, pre-Messinian EMS

sediment has a relatively uniform composition dominated by Nile sediment with minor input

of Sahara dust compared to present-day. We relate the increase in Sahara dust production to

a late Miocene shift in climate and the progressive aridification of the Sahara region.

76 | CHAPTER 3

The bulk addition of 0.5-5% EMS sediment to a depleted MORB mantle satisfactorily

explains the Pb-Nd isotope variation in volcanic centres from the western and central Aegean

arc. However, the contribution of EMS sediment cannot account for the anomalously

unradiogenic Pb-composition of Nisyros volcanic rocks in the easternmost section of the Arc.

Increased input of Nile sediment in the east has an opposite effect and would increase the 206Pb/204Pb of the lavas. Hence, we relate the geochemical signature of the eastern volcanic

centres to a distinct mantle source rather than a difference in subducted sediment

composition. Aeolian arc volcanic rocks do not show significant contribution of EMS sediment

to the mantle wedge, with the exception of Stromboli. Bulk addition of 5-20% of EMS

sediment can explain some of the variation observed in Stromboli volcanic rocks, which

appears to make it unique amongst the other Aeolian volcanic centres.

Acknowledgements

EMS sediment samples were kindly provided by the DSDP/ODP core repository in Bremen. We

thank Simon Troelstra for donating six surface sediment samples from the African passive

margin. Roel van Elsas, Sergei Matveev, Bas van der Wagt and Richard Smeets are thanked

for their assistance with sample preparation and analyses. The MC-ICPMS and TIMS facilities

at VU University Amsterdam are funded by NWO through grant no. 175.107.404.01 and

834.10.001 respectively.


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46 | CHAPTER 3 3.pdf · Recycling of oceanic crust and sediment back into the mantle occurs at subduction zones. Fluids and melts expelled from the subducting slab are an important