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ORIGINAL PAPER

Geochemistry and tectonic setting of mafic rocksfrom the Othris Ophiolite, Greece

Matthias G. Barth Æ Tatjana M. Gluhak

Received: 26 July 2007 / Accepted: 3 June 2008 / Published online: 17 June 2008

� Springer-Verlag 2008

Abstract We present new geochemical analyses of

minerals and whole rocks for a suite of mafic rocks from

the crustal section of the Othris Ophiolite in central Greece.

The mafic rocks form three chemically distinct groups.

Group 1 is characterized by N-MORB-type basalt and

basaltic andesite with Na- and Ti-rich clinopyroxenes.

These rocks show mild LREE depletion and no HFSE

anomalies, consistent with moderate degrees (*15%) of

anhydrous partial melting of depleted mantle followed by

30–50% crystal fractionation. Group 2 is represented by

E-MORB-type basalt with clinopyroxenes with higher Ti

contents than Group 1 basalts. Group 2 basalts also have

higher concentrations of incompatible trace elements with

slightly lower HREE contents than Group 1 basalts. These

chemical features can be explained by *10% partial

melting of an enriched mantle source. Group 3 includes

high MgO cumulates with Na- and Ti-poor clinopyroxene,

forsteritic olivine, and Cr-rich spinel. The cumulates show

strong depletion of HFSE, low HREE contents, and LREE

enrichments. These rocks may have formed by olivine

accumulation from boninitic magmas. The petrogenesis of

the N-MORB-type basalts and basaltic andesites is in

excellent agreement with the melting conditions inferred

from the MOR-type peridotites in Othris. The occurrence

of both N- and E-MORB-type lavas suggests that the

mantle generating the lavas of the Othris Ophiolite must

have been heterogeneous on a comparatively fine scale.

Furthermore, the inferred parental magmas of the SSZ-type

cumulates are broadly complementary to the SSZ-type

peridotites found in Othris. These results suggest that the

crustal section may be genetically related to the mantle

section. In the Othris Ophiolite mafic rocks recording

magmatic processes characteristic both of mid-ocean rid-

ges and subduction zones occur within close spatial

association. These observations are consistent with the

formation of the Othris Ophiolite in the upper plate of a

newly created intra-oceanic subduction zone.

Keywords Ophiolite � Subduction � Basalt �Magma genesis � Hellenides � Jurassic

Introduction

Ophiolites are products of complex tectonic and magmatic

processes that operated during the initial rifting through

seafloor spreading to subduction-facilitated emplacement

stages of ancient oceanic lithosphere in various tectonic

settings (e.g., Coleman 1977; Nicolas 1989; Dilek et al.

2005). Tethyan-type ophiolites (cf. Moores 1982) experi-

enced active margin tectonics during their incorporation

into continental margins through collisional processes and

in many cases also during their magmatic evolution (Pearce

et al. 1984). Ophiolites representing different stages of

evolution of an ocean basin may display different structural

and petrological features. The crustal section of Tethyan-

type ophiolites may be genetically related to the mantle

section if both formed in an oceanic environment, either by

seafloor spreading or in the upper plate of an intra-oceanic

subduction zone (the Mediterranean-type ophiolites of

Communicated by T.L. Grove.

Electronic supplementary material The online version of thisarticle (doi:10.1007/s00410-008-0318-9) contains supplementarymaterial, which is available to authorized users.

M. G. Barth (&) � T. M. Gluhak

Institut fur Geowissenschaften, Universitat Mainz,

Becherweg 21, 55099 Mainz, Germany

e-mail: [email protected]; [email protected]

123

Contrib Mineral Petrol (2009) 157:23–40

DOI 10.1007/s00410-008-0318-9

http://dx.doi.org/10.1007/s00410-008-0318-9

Dilek 2003). However, in some ophiolites spatially asso-

ciated depleted peridotites and basaltic rocks are not linked

by a genetic melt and residua relationship (the Ligurian-

type of Dilek 2003). These ophiolites may have formed

during the early stages of opening of an ocean basin, fol-

lowing continental rifting and breakup (Rampone and

Piccardo 2000). Therefore, the evidence (or lack) of a co-

genetic relationship between mantle peridotites and

associated magmatic rocks can help to constrain the tec-

tonic setting and specific mode of generation of an

ophiolite.

In this paper, we present new geochemical analyses of

minerals and whole rocks for a suite of mafic rocks from

the crustal section of the Othris Ophiolite, Greece, and

discuss the melting conditions and mantle sources of these

mafic rocks. The main goals of this contribution are to test

if the crustal section is genetically linked to the mantle

peridotites and to further constrain the evolution and tec-

tonic setting of the Othris Ophiolite. We also integrate

microprobe data we previously collected on ultramafic

rocks of the mantle section that were never fully published

before. Results of our petrographic and microprobe survey

of the Othris mantle rocks are presented in the electronic

supplementary material (eAppendix 1 and eTables 1, 2, 3,

4, 5).

Geological setting

The Othris Ophiolite in central Greece is located in the

Mirdita-Subpelagonian zone and is part of a NNW-trending

belt, which includes the western Hellenic ophiolites in

Greece, the Mirdita ophiolites in Albania, and the Dinaric

ophiolites in Serbia and Croatia. These Jurassic ophiolites

are bounded on the east and the west by the Korabi-Pela-

gonian and the Apulian microcontinents, respectively (see

review by Robertson 2002). The Axios–Vardar Zone east

of the Korabi-Pelagonian microcontinent forms a second

ophiolite belt that is subparallel to the Mirdita-Subpela-

gonian zone. The ophiolites within both belts are

interpreted to be remnants of the Neotethys Ocean, a net-

work of small ocean basins and microcontinents, which

existed between Eurasia and Gondwana during the Meso-

zoic–Early Tertiary (e.g., Robertson et al. 1991; Stampfli

and Borel 2002). The paleogeographic origin of the Othris

Ophiolite within the Neotethys is debated (e.g., Smith

1993; Robertson and Shallo 2000). One view expressed by

many authors is that the Othris Ophiolite formed in the

Mirdita-Pindos ocean, a relatively narrow ocean basin

lying west of the Korabi-Pelagonian microcontinent (e.g.,

Robertson and Karamata 1994). An alternative view is that

the Othris Ophiolite originated in the Meliata-Vardar

ocean, lying east of the Korabi-Pelagonian microcontinent

(e.g., Smith and Spray 1984). In this paper we assume an

origin for the Othris Ophiolite within the Mirdita-Pindos

ocean. According to Robertson and Shallo (2000), the

Mirdita-Pindos oceanic basin, located between the Apulian

continent in the west and the Korabi-Pelagonian micro-

continent in the east, opened during the Late Triassic–Early

Jurassic. During the Middle Jurassic (160–170 Ma) a

south-westward-dipping intra-oceanic subduction zone

became active within the Mirdita-Pindos oceanic basin.

Subsequently, the ophiolites were emplaced onto the con-

tinental margin during the Late Jurassic–Early Cretaceous.

In the Albanian sector of the Mirdita-Subpelagonian

zone, individual ophiolite massifs in the west near the

Apulian microcontinent are mainly composed of mid-

ocean ridge (MOR)-type basalts and lherzolites, whereas

those in the east near the Korabi-Pelagonian microconti-

nent contain dominantly supra-subduction zone (SSZ)-type

basalts and harzburgites (Robertson and Shallo 2000;

Shallo and Dilek 2003; Saccani et al. 2004). In the Greek

sector, a sharp geographical distinction cannot be made.

The Vourinos ophiolite in the east of the Mirdita-Subpel-

agonian zone, which is characterized by mafic and

ultramafic sequences with island arc and boninitic affinities

(Bizimis et al. 2000; Saccani et al. 2004), is a SSZ ophio-

lite, comparable to the eastern belt in Albania. By con-

trast, pure MOR-type ophiolites, similar to the western

belt in Albania, are subordinate. The crustal sections of

the Pindos and Othris ophiolite complexes are thought to

have formed in more than one tectonic setting as their

basalts exhibit both mid-ocean ridge (MOR) and island

arc affinities (Pearce et al. 1984; Jones and Robertson

1991; Photiades et al. 2003; Saccani and Photiades 2004).

In addition, the mantle section of the Othris Ophiolite

shows evidence for both comparatively fertile MOR-type

lherzolites and depleted SSZ-type harzburgites (Barth

et al. 2008).

The Othris Mountains of central Greece extend from

east of Almiros to the Pindos Mountains west of Lamia

(Fig. 1). The Subpelagonian zone in the Othris Moun-

tains consists of a sequence of thrust sheets that had

been emplaced onto the Triassic-Jurassic carbonate

platform overlying the late Carboniferous Hercynian

basement (Hynes et al. 1972; Smith et al. 1975; Ferriere

1985). These thrust sheets formed an ordered progression

from a submarine fan into structurally higher and pene-

contemporaneous pelagic basin sections and finally

ophiolites at the highest stratigraphic level (Hynes 1974;

Menzies and Allen 1974; Smith et al. 1975; Smith

1993). The ophiolitic units of Othris are stacked in

reverse stratigraphic order: cherts and pillow lavas at the

base have been overthrust by sheeted dikes and gabbro

cumulates and then by ultramafic cumulates, near-Moho

mantle rocks, a harzburgite thrust sheet, and finally a

lherzolite-plagioclase lherzolite thrust sheet on top

24 Contrib Mineral Petrol (2009) 157:23–40

123

(Rassios and Konstantopoulou 1993; Rassios and Smith

2000).

The interpretation of the origin and emplacement of the

Othris Ophiolite has changed over time and remains con-

troversial (Smith and Rassios 2003). The first models that

placed the Othris Ophiolite in a plate tectonic setting

interpreted it as being formed near a continental margin at

the inception of rifting (Hynes 1974; Menzies and Allen

1974; Menzies 1976) and subsequently being emplaced by

convergent margin tectonics at a subduction zone (Hynes

et al. 1972). Later studies argued for a relatively slow-

spreading mid-ocean ridge environment (Rassios and

Konstantopoulou 1993; Dijkstra et al. 2001, 2003) or an

island-arc environment (Bizimis et al. 2000; Rassios and

Smith 2000). Recently, Barth et al. (2008) proposed that

the Othris Ophiolite originated during the initial stages of

subduction at or near a mid-ocean ridge, when oceanic

extension rapidly changed to convergence.

Sampling

For the present study sampling was focused on mafic

volcanic rocks in order to investigate the geochemical

variability of the crustal section. In the Othris complex

there are two separate magmatic series (Smith et al.

1975): an earlier suite formed in a continental margin, the

Agrilia Formation, which was probably related to the

opening of the ocean basin, and a later suite, the Mirna

Group, in which gabbroic cumulates and ultramafics were

identified representing spreading in that basin. In general,

outcrop is poor in the crustal section of the Othris

Ophiolite. Therefore, the majority of samples were col-

lected from roadcuts on the road from Lamia to Domokos.

A total of eleven samples were collected (Fig. 1;

Table 1): five from the Agrilia Formation (samples A1,

A9, A10, A11, and A32), two from the Mirna group

(samples M3 and M30), and four from the Sipetorrema

Pillow Lava unit (S5, S6, S7, and S8). The Triassic

Agrilia Formation is part of the Othris group (Smith et al.

1975). The Sipetorrema Pillow Lava unit is part of the

Jurassic Mirna Group, which contains the ophiolite com-

plex and which has been thrust over the Othris Group

Katáchloron

Eretria

FournosKaïtsa

10 km

Almiros

Domokos

Lamia

Stillis

Gulf of MaliakosN

Legendultramafic rocks

mafic rockspillow lavas

Vourinos

Othris

22.5°E

39°N

The Othris Ophiolite22°E

Pindos

Kédros

shearzone FK

M30

A1A9A10A11

M3

S5

S6

S7S8

A32

Fig.1 Simplified geological

map of the Othris Ophiolite

showing the location of the

study area. Modified from

Rassios and Konstantopoulou

(1993). Inset: location map

showing the Othris, Pindos, and

Vourinos ophiolites

Table 1 Locations of the sample collection sites of the mafic rocks

from the Othris Ophiolite, Greece

Sample Type Latitude Longitude

Agrilia Formation

A1 Pillow basalt 38�56039.200N 22�24024.800E





Mirna Group

M3 Dolerite 39�0040.800N 22�22037.900E

M30 Dolerite 39�0057.300N 22�1609.900E

Sipetorrema Pillow Lava Unit

S5 Pillow basalt 38�59021.200N 22�22025.500E

S6 Pillow basalt 38�58043.600N 22�22046.900E

S7 Boninitic cumulate 38�5808.600N 22�23012.800E

S8 Boninitic cumulate 38�5808.600N 22�23012.800E

Contrib Mineral Petrol (2009) 157:23–40 25

123

(Smith et al. 1975). Samples labeled as Mirna Group are

mafic rocks collected from the Mirna Group excluding the

Sipetorrema Pillow Lava unit.

Analytical methods

Major elements

Whole rock major elements and the trace elements Sc, V,

Cr, Ni, Co, Cu, and Zn were determined by X-ray fluo-

rescence (XRF) on fused glass beads and pressed powder

pellets at the University of Mainz.

Mineral major element compositions of the mafic rocks

were determined using the Jeol JXA-8900 RL wavelength-

dispersive electron microprobe (EMP) at the University of

Mainz. Olivine, orthopyroxene, clinopyroxene, and spinel

were analyzed using an accelerating potential of 20 kV, a

beam current of 12 nA, and a spot size of 2 lm. An

accelerating potential of 15 kV, a beam current of 8 nA,

and a spot size of 5 lm were used to analyze plagioclase.

Trace elements

To get high-precision whole rock data for rare earth

elements and other trace elements, we have applied a

recently developed laser ablation-inductively coupled

plasma-mass spectrometry (LA–ICP-MS) technique using

an automated iridium strip heater (Nehring et al. 2008).

About 40 mg of the rock powder was placed on an irid-

ium strip without any flux agent. The melting of the

samples to glass beads took place in a closed box under

an argon atmosphere to suppress oxidation and to limit

volatilization of elements with low boiling points (e.g.,

Cs, Pb). Melting conditions for basaltic samples were

1,200�C and 10 s. The fused glass beads were analyzed

by LA–ICP-MS at the University of Mainz. Ablation was

achieved with a NewWave Research UP-213 Nd:YAG

laser ablation system, using a pulse repetition rate of

10 Hz, pulse energies of *0.3 mJ, and 100 lm crater

diameters. Analyses were performed on an Agilent

7500ce inductively coupled plasma-mass spectrometer in

pulse counting mode (one point per peak and 10 ms dwell

time). Data reduction was carried out using the software

‘‘Glitter’’. The amount of material ablated in laser sam-

pling is different for each spot analysis. Consequently, the

detection limits are different for each spot and are cal-

culated for each individual acquisition. Detection limits

generally range between 0.001 and 0.5 ppm (lg/g). 44Ca

was used as internal standard. Analyses were calibrated

against the silicate glass reference material NIST 612

using the values of Pearce et al. (1997), and the US

Geological Survey (USGS) glass standard BCR-2G was

measured to monitor accuracy.

Results

Petrography

All samples except S7 and S8 have aphanitic textures

typical for mafic volcanic rocks.

The five samples from the Agrilia formation display

fine-grained microporphyritic textures with euhedral to

subhedral elongate plagioclase microphenocrysts and

granular clinopyroxene microphenocrysts embedded in a

groundmass consisting of lath-shaped plagioclase micro-

lites and granular pyroxenes.

The two samples from the Mirna Group have very fine-

grained microporphyritic to intergranular textures. In mi-

croporphyritic varieties, phenocrysts are represented by

euhedral to subhedral plagioclase, clinopyroxene, and rare

spinel. Clinopyroxene phenocrysts are often zoned.

Samples S5 and S6 from the Sipetorrema Pillow Lava

unit have the finest-grained groundmass among the samples

studies. The microphenocrysts are euhedral to subhedral

elongated plagioclase and granular clinopyroxene.

Samples S7 and S8 from the Sipetorrema Pillow Lava

unit are highly phyric basalts containing abundant euhedral

to subhedral olivine and pyroxene phenocrysts 2–3 mm in

size, suggestive of olivine and pyroxene accumulation.

Additional phenocrysts are represented by orthopyroxene

and minor spinel.

All samples show extensive to high degrees of hydro-

thermal ocean-floor alteration. This alteration is manifested

by the presence of carbonate + iron oxide ± chlorite veins

and crusts replacing the primary lithology. Plagioclase is

partly to completely albitized.

Whole rock chemistry

The major element contents of whole rocks from the Othris

Ophiolite are presented in Table 2 and plotted in element

oxide versus Mg# [100 9 Mg/(Mg + Fe)] abundance plots

(Fig. 2).

As stated in the previous section, all samples are

extensively to highly altered. This observation is reflected

by the high measured LOI values (3.05–11.4%, Table 2).

The chemical effects of hydrothermal alteration, that is,

variable mobilization of CaO, MgO, alkali elements (Na,

K) and large ion lithophile elements (LILE) such as Rb, Ba,

and Sr (Staudigel 2003), can be observed in the samples.

For example, CaO contents in the samples from the Agrilia

Formation are highly variable, ranging from 5.08 to

14.9 wt%, reflecting both loss of CaO in sample A10 and

carbonate addition in sample A9 (Fig. 2c).

The covariation of some selected major and trace ele-

ments against Zr (used as an indicator of differentiation) is

shown in Fig. 3. No systematic elemental variation with


123

Table 2 Whole rock major and trace element composition of mafic rocks from Othris

Agrilia Formation Mirna Group Sipetorrema Pillow Lava

Sample A1 A9 A10 A11 A32 M3 M30 S5 S6 S7 S8Type E-M E-M N-M E-M E-M E-M N-M N-M N-M SSZ SSZ

SiO2 (wt%) 46.0 44.2 49.2 47.9 43.6 47.7 48.8 50.0 45.8 40.5 41.7

TiO2 1.09 0.69 1.60 0.78 1.02 1.05 2.22 1.20 1.25 0.26 0.33

Al2O3 14.3 11.1 14.7 13.0 13.5 15.3 11.8 14.6 15.1 4.43 4.97

FeO 8.60 8.30 11.6 7.41 8.83 8.82 10.9 11.3 10.7 9.59 9.95

MnO 0.11 0.15 0.17 0.14 0.12 0.18 0.18 0.19 0.17 0.14 0.15

MgO 6.51 6.73 6.95 5.35 6.27 7.47 3.75 7.33 7.71 30.8 26.7

CaO 10.3 14.9 5.08 13.4 12.5 7.46 9.39 11.6 11.5 3.43 3.84

Na2O 4.24 3.94 4.14 5.07 3.24 4.01 3.94 3.37 2.84 0.05 0.21

K2O 1.23 0.53 2.04 0.10 1.43 1.60 0.05 0.19 0.11 0.02 0.03

P2O5 0.22 0.08 0.14 0.10 0.20 0.12 0.19 0.08 0.10 0.05 0.08

Cr2O3 0.08 0.16 0.01 0.06 0.10 0.02 0.01 0.05 0.05 0.34 0.36

NiO 0.03 0.03 b.d. 0.01 0.02 0.01 0.01 0.01 0.01 0.18 0.18

LOI 7.18 9.06 4.25 6.71 9.02 6.17 8.61 0.05 4.66 10.05 11.40

Total 99.91 99.92 99.88 99.95 99.90 99.90 99.92 99.88 99.90 99.91 99.89

Mg# 57.4 59.1 51.6 56.3 55.9 60.1 38.0 53.5 56.3 85.1 82.7

Li (ppm) 30.6 5.98 18.5 4.39 28.5 30.0 2.5 30.9 24.6 29.0 73.6

Sc 39.7 35.8 36.5 36.9 39.3 44.6 55.0 44.1 44.6 22.6 24.1

V 288 251 328 262 320 321 555 271 257 120 173

Cr 522 472 36.2 299 698 146 113 332 326 1825 2039

Co 47 47 36 31 43 33 48 46 45 97 103

Ni 191 227 15.6 73.0 193 72.7 60.7 57.6 59.3 1413 1662

Cu 54 51 140 35 51 27 42 103 96 22 47

Zn 78 69 88 63 75 49 97 70 69 56 62

Ga 14 11 18 11 12 15 17 14 19 6 7

Rb 14.6 5.99 15.8 0.93 19.7 28.5 1.02 2.02 0.62 0.75 0.65

Sr 172 151 121 102 203 190 115 221 80.6 43.0 101

Y 19.8 15.6 34.4 17.2 18.5 20.6 51.5 24.6 24.4 5.95 8.18

Zr 67.7 41.2 101 45.8 63.6 75.2 172 58.6 78.0 15.2 22.7

Nb 22.0 3.98 2.48 4.81 20.6 9.28 3.74 1.28 2.43 0.33 0.50

Cs 0.48 0.36 0.10 0.03 0.97 0.43 0.07 0.03 0.02 0.15 0.16

Ba 62.8 74.1 65.2 8.92 69.6 475 7.20 7.48 6.87 1.43 6.40

La 13.9 4.15 4.67 3.95 12.9 7.77 5.80 2.13 3.24 2.78 4.08

Ce 27.7 9.01 13.8 9.67 25.9 17.6 19.2 7.03 9.45 7.25 11.2

Pr 3.25 1.29 2.31 1.39 3.04 2.35 3.24 1.25 1.50 0.95 1.42

Nd 13.5 6.38 12.4 7.00 12.7 10.3 17.7 7.20 8.21 4.02 6.00

Sm 2.95 1.84 3.88 2.07 2.69 2.82 5.99 2.47 2.83 0.78 1.12

Eu 1.00 0.66 1.43 0.77 0.91 1.04 2.18 1.02 1.08 0.26 0.38

Gd 3.28 2.26 5.31 2.64 2.86 3.21 7.57 3.46 3.55 0.94 1.22

Dy 3.49 2.79 6.28 3.03 3.18 3.79 9.33 4.36 4.39 1.10 1.50

Er 2.06 1.68 3.99 1.89 1.94 2.17 5.67 2.63 2.69 0.66 0.84

Yb 2.09 1.74 4.02 1.91 1.97 2.17 5.61 2.62 2.75 0.70 0.94

Lu 0.33 0.26 0.61 0.27 0.30 0.34 0.82 0.38 0.41 0.10 0.13

Hf 1.80 1.17 2.81 1.30 1.69 2.08 4.56 1.66 1.96 0.49 0.70

Ta 1.39 0.25 0.17 0.28 1.28 0.64 0.27 0.10 0.16 0.03 0.03

Pb 0.90 0.56 0.20 0.40 0.64 0.33 1.32 0.23 0.32 0.24 0.37

Th 1.53 0.40 0.32 0.40 1.41 1.15 0.25 0.09 0.15 0.13 0.18

U 0.29 0.19 0.21 0.31 0.35 0.22 0.12 0.04 0.05 0.05 0.13

Major elements, Ga, Co, Cu, and Zn determined by XRF. All other trace elements determined by ICP-MS

b.d. below detection limit, Sample types: E-M E-MORB, N-M N-MORB, SSZ supra-subduction zone


123

magmatic series or geographic location can be observed.

By contrast, the good correlation of most of the elements

with Zr suggests that common processes controlled the

compositions of the different magmatic series. For this

reason, the geochemical features of the analyzed samples

are presented according to the three distinct groups of

mafic rocks recognized on chemical bases: (a) normal mid-

ocean ridge basalt (N-MORB)-type; (b) enriched mid-

ocean ridge (E-MORB)-type; and (c) supra-subduction

zone (SSZ)-type. The major geochemical differences

between these groups lie in the different concentrations of

immobile trace elements such as rare earth elements (REE)

and high field strength elements (HFSE).

Normal mid-ocean ridge-type basalts and basaltic

andesites

Four of the eleven samples (A10, M30, S5, and S6) have

compositions comparable to normal mid-ocean ridge

basalts (N-MORB) and basaltic andesites. The MgO con-

tents (3.75–7.71 wt%) and Mg# [100 9 Mg/(Mg + Fe)]

(38.0–56.3) of these samples indicate a slightly differenti-

ated to differentiated nature. The TiO2 contents range from

1.20 to 2.22 wt% (Fig. 2b), comparable to the high-Ti

basalts and basaltic andesites found in the Agoriani Mel-

ange in the northwestern part of the Othris Ophiolite

(Photiades et al. 2003). The Ti/V ratios range from 26 to 31

(Fig. 4) and plot in the field for MORB (Shervais 1982).

The basalts display low to moderate Cr (36–332 ppm) and

Ni (16–61 ppm) contents (Fig. 2e, f). REE patterns are

consistent with N-MORB compositions (Fig. 5a), as they

have moderate depletions in light REE (LREE; (La/

Sm)N = 0.54–0.76) and flat heavy REE (HREE) patterns

(YbN = 12.5–26.8). The samples show an overall deple-

tion in incompatible trace elements, do not show any HFSE

anomalies, and show both positive and negative Sr ano-

malies (Fig. 6a), indicative of plagioclase accumulation

and fractionation, respectively.

SiO [wt%]2

Mg#

3040506070809044

46

48

50

52

54

56

N-MORB-typeE-MORB-typeSSZ-typeMid-Atlantic Ridge

TiO [wt%]2

Mg#

304050607080900

1

2

3

CaO [wt%]

Mg#

30405060708090

4

8

12

16 FeO* [wt%]

Mg#

304050607080906

8

10

12

14

16

18

Cr [ppm]

Mg#

304050607080900

500

1000

1500

2000

2500Ni [ppm]

Mg#

3040506070809010

100

1000

(a) (b)

(c) (d)

(f)(e)

Fig. 2 Variation diagram of

selected major and trace

elements versus Mg#

[100 9 Mg/(Mg + Fe)] for

whole rock compositions of

mafic rocks from Othris,

Greece, compared to mid-ocean

ridge basalts (MORB) from the

Mid-Atlantic ridge (55�S–52�N;

data source: PetDB database).

Major element concentrations

are recalculated to 100% on

LOI-free bases


123

Enriched mid-ocean ridge-type basalts

Five samples (A1, A9, A11, A32, and M3) have composi-

tions comparable to enriched mid-ocean ridge basalts

(E-MORB). The MgO contents (5.35–7.47 wt%) and Mg#

(55.9–60.1) of these samples indicate a slightly differenti-

ated nature. The TiO2 contents range from 0.69 to 1.09 wt%

(Fig. 2b), comparable to the high-Ti basalts and basaltic

andesites of Photiades et al. (2003). The Ti/V ratios range

from 18 to 24 (Fig. 4) and plot at the border between the

fields for convergent margin basalts and MORB (Shervais

1982). The basalts display moderate to high Cr (146–

698 ppm) and Ni (73–227 ppm) contents (Fig. 2e, f). REE

patterns are consistent with E-MORB compositions

(Fig. 5b), as they have enrichments in LREE [(La/

Sm)N = 1.2–3.0] and flat HREE patterns (YbN = 8.3–

10.4). The samples show an overall enrichment in incom-

patible trace elements, do not show any negative HFSE

anomalies, and show no or only small Sr anomalies (Fig. 6b).

MgO [wt%]

Zr [ppm]

0 50 100 150 200 2500

5

10

15

FeO* [wt%]

Zr [ppm]

0 50 100 150 200 2506

7

8

9

10

11

12

13

14

Y [ppm]

Zr [ppm]

0 50 100 150 200 2500

10

20

30

40

50

60

TiO [wt%]2

Zr [ppm]

0 50 100 150 200 2500

1

2

3

Cr [ppm]

Zr [ppm]

0 50 100 150 200 2500

500

1000

1500

2000

2500

La [ppm]

Zr [ppm]

0 50 100 150 200 2500

5

10

15

20

25

SSZ-typeMid-Atlantic Ridge

E-MORB-typeN-MORB-type

enrichedsource

depletedsource

(a) (b)

(c) (d)

(e) (f)

Fig. 3 Variation diagram of

selected major and trace

elements versus Zr (in ppm) for

whole rock compositions of

mafic rocks from Othris,

Greece, compared to mid-ocean

ridge basalts (MORB) from the

Mid-Atlantic ridge (55�S–52�N;

data source: PetDB database).

Major element concentrations

are recalculated to 100% on

LOI-free bases

bulk rock

Ti/1000 [ppm]

V [p

pm]

00 2 4 6 8 10 12 14 16

100

200

300

400

500

600

N-MORB-typeE-MORB-typeSSZ-typeMid-Atlantic Ridge

Ti/V = 10

20

50

Fig. 4 Vanadium versus titanium discrimination diagram for mafic

rocks from Othris compared to MORB from the Mid-Atlantic

ridge(55�S–52�N; data source: PetDB database). Modified after

Shervais (1982). Ti/V \ 20: SSZ basalts; 20 \ Ti/V \ 50: MORB;

Ti/V [ 50: within-plate basalts


123

Supra-subduction zone-type cumulates

The two highly phyric samples (S7 and S8) have much

higher MgO (26.7–30.8 wt%) and lower SiO2 (40.5–

41.7 wt%) contents than the MORB-type samples

(Fig. 2a). The very high MgO, Cr (1,825–2,039 ppm),

and Ni (1,413–1,662 ppm) contents (Fig. 2e, f) and high

Mg# (82.7–85.1) of these samples are consistent with an

accumulation of olivine and pyroxene phenocrysts. The

TiO2 concentrations are very low (0.26–0.33 wt%,

Fig. 2b), comparable to the very low-Ti basaltic andesites

and andesites of Photiades et al. (2003). The Ti/V ratios

range from 13 to 15 (Fig. 4) and plot in the field for

convergent margin basalts (Shervais 1982). The cumu-

lates display flat HREE patterns with low HREE

concentrations (YbN = 3.4–4.5) and enrichments in

LREE [(La/Sm)N = 2.2–2.3] and other incompatible

trace elements (Figs. 5c, 6c). The samples show strong

negative Nb and Ta anomalies (Nb/La = 0.12). These

characteristics are typical for magmas generated in SSZ

settings.

These rocks have compositions similar to the komatiitic

rocks from the Agrilia Formation described by Cameron

and Nisbet (1982), Paraskevopoulos and Economou

(1986), Economou-Eliopoulos and Paraskevopoulos

(1989), and Paraskevopoulos and Economou-Eliopoulos

(1997).

N-MORB-typeanhydrous near-fractional melting

sam

ple

/N-M

OR

B

1

10

A10M30S5S6

E-MORB-typeanhydrous near-fractional melting

sam

ple

/N-M

OR

B

1

10

A1A9A11A32M3

SSZ-typefluid-induced hydrous melting

La Ce Pr Sr Nd Sm Zr Eu Ti Gd Dy Y YbEr

sam

ple

/N-M

OR

B

0.1

1

10

S7S8

16% accumulated melt(4% garnet, 12% spinel)

7% accumulated melt(4% garnet, 3% spinel)no crystal fractionation

30% olivine addition

(a)

(b)

(c)

0% crystal fractionation

50%

20% accumulated melt(8% garnet, 12% spinel)no crystal fractionation

30%

10% accumulated melt(spinel stability field)

Fig. 5 N-MORB-normalized whole rock trace element diagram for

mafic rocks from Othris. a Normal mid-ocean ridge-type basalts and

basaltic andesites, b enriched mid-ocean ridge-type basalts, c supra-

subduction zone-type cumulates. Also shown are the results obtained

for different melting models (see text for details). Element abun-

dances are normalized to the N-MORB values of Sun and

McDonough (1989)

N-MORB-type

sam

ple

/prim

itive

man

tle

10

100

A10M30S5S6N-MORB

E-MORB-type

sam

ple

/prim

itive

man

tle1

10

100

A1A9A11A32M3E-MORB

SSZ-type

Th U Nb Ta La Ce Pr Sr NdSm Zr Hf Eu Ti Gd Tb Dy Ho Er TmY Yb Lu

sam

ple

/prim

itive

man

tle

1

10

S7S8

(a)

(b)

(c)

Sr Ti

Ti

Ti

Fig. 6 Mantle-normalized whole rock trace element diagrams for

mafic rocks from Othris. a Normal mid-ocean ridge-type basalts and

basaltic andesites, b enriched mid-ocean ridge-type basalts, c supra-

subduction zone-type cumulates. Also shown are the N-MORB and

E-MORB values of Sun and McDonough (1989). Element abun-

dances are normalized to the primitive mantle values of McDonough

and Sun (1995)


123

Mineral chemistry

The major element compositions of plagioclase, clinopy-

roxene, orthopyroxene, spinel, and olivine are presented in

Tables 3, 4, 5, 6, 7.

Plagioclase

Plagioclase has anorthite contents of 20–52% in the

N-MORB-type basalt S5 and of up to 68% in the SSZ-type

cumulate S8. Plagioclase in the E-MORB-type basalts was

completely albitized. Considering that in almost all sam-

ples plagioclase is heavily altered, the measured anorthite

contents may not reflect the original igneous compositions.

Pyroxene

The clinopyroxenes in the mafic rocks have compositions

ranging from aluminian diopsides to Mg-rich augites.

Clinopyroxenes are heterogeneous within samples and are

variable between samples, varying between Wo43En51Fs6

and Wo42En33Fs25. Cr2O3 concentrations are positively and

Al2O3 and TiO2 concentrations are negatively correlated

with Mg#. Na2O does not show a clear correlation with

Mg#. Clinopyroxenes in the three groups of mafic rocks

mirror the whole rock compositions, i.e., the MORB-type

and SSZ-type samples plot in the MORB field and con-

vergent margin field of Beccaluva et al. (1989), respectively

(Fig. 7). At a given Mg# clinopyroxenes in N-MORB-type

rocks have comparatively high Na2O and Cr2O3 concen-

trations coupled with moderate Al2O3 and TiO2

concentrations. Clinopyroxenes in E-MORB-type rocks

have higher Al2O3 and TiO2 concentrations coupled with

moderate Na2O and Cr2O3 concentrations. Clinopyroxenes

in the SSZ-type rocks have cores with very high Mg#, high

Cr2O3, and very low TiO2 and Al2O3 concentrations. Rims

have more evolved compositions similar to clinopyroxenes

in N-MORB-type rocks.

Orthopyroxene was found only in the SSZ-type sample

S8. The orthopyroxenes analyzed are enstatites (Mg# = 79)

with low Cr2O3 concentrations (0.03–0.05 wt%).

Spinel

Spinel was observed in three E-MORB-type samples (A1,

A9, and A11) and in the two SSZ-type samples (S7 and

S8). Spinel composition is slightly heterogeneous within

samples and is variable between samples (Fig. 8). Spinels

in E-MORB-type rocks have Cr# [100*Cr/(Cr + Al)]

ranging from 54 to 77 and Mg# from 59 to 69. These

compositions plot at the Cr-rich end of the MORB field and

in the island arc tholeiite field of Barnes and Roeder

(2001). Spinels from the SSZ-type samples have higher

Cr# (78–89) and lower Mg# (38–60) than spinels from the

MORB-type rocks. These very Cr-rich spinels are compa-

rable to spinels in boninites (Barnes and Roeder 2001). All

samples have low TiO2 contents (0.17–0.66 wt%).

Olivine

Primary olivine has been preserved only in the SSZ-type

samples S7 and S8 (Fig. 9). Olivine in the studied rocks

have high forsterite contents ranging from 88.9 to 94.0,

relatively high NiO contents from 0.24 to 0.47 wt%, and

Cr2O3 from 0.03 to 0.18 wt%.

Discussion

The new mineral major element and whole rock major and

trace element data obtained for the mafic rocks from Othris

suggest that these rocks were generated from different

mantle sources and formed in distinct tectonic settings. The

main objectives of this discussion are to quantify the

variations in the degree of partial melting, to evaluate the

effects of crystal fractionation, to determine the different

mantle sources of the crustal section of the Othris Ophio-

lite, and to constrain the tectonic setting of the Othris

ophiolite.

Trace element modeling

Trace element abundances were used to evaluate if the

compositional spectrum of the MORB-type rocks can be

explained by anhydrous partial melting both in the spinel

and garnet stability field and subsequent crystal

Table 3 Major element composition of plagioclase determined by

electron microprobe

Sample S5 S8

Type N-MORB SSZ

Spots n = 10 1r n = 13 1r

SiO2 62.43 3.65 52.49 1.96

Al2O3 23.16 2.06 29.70 1.48

Fe2O3 0.69 0.25 1.15 0.22

MgO 0.07 0.07 0.43 0.54

CaO 3.92 2.71 12.48 1.81

BaO 0.03 0.02 0.04 0.03

Na2O 9.37 1.64 4.30 0.88

K2O 0.20 0.20 0.07 0.04

Total 99.86 100.65

An# 18.9 13.3 61.5 8.5

Sample locations: S Sipetorrema Pillow Lava unit


123

fractionation events. In addition, a fluid-induced referti-

lization-hydrous melting model was assessed for the SSZ-

type rocks. An incremental, non-modal batch melting

model was used, in which melt was extracted at 0.1%

increments. Initial composition, source mineralogy, melt-

ing phase proportions, and partition coefficients are given

in Table 8. Details of these models have been described in

Barth et al. (2003). Rayleigh fractionation of a solid

assemblage of 60% olivine + 38% clinopyroxene + 2%

spinel was used to model crystal fractionation. This sim-

plified fractionation model does not include plagioclase,

which is observed in the Othris samples as phenocrysts.

However, since plagioclase has low partition coefficients

for Cr and Y (Table 8) its inclusion in the fractionation

assemblage would not alter the results significantly (see

below). Moreover, this simple melting and fractionation

model correctly reproduces the compositional trends of

MORB from the mid-Atlantic ridge (Fig. 10).

The extent of partial melting and crystal fractionation

was calculated based on the concentration of Cr, Y and

HREE (Figs. 5, 10); the mantle source was evaluated based

on the concentrations of Th, Nb, Ta, and LREE as these

elements are judged to be relatively immobile during

hydrothermal alteration (Staudigel 2003).

Table 4 Major element composition of clinopyroxene determined by electron microprobe

Sample A1 A9 A10 A11 A32

Type E-MORB E-MORB N-MORB E-MORB E-MORB

Spots n = 21 1r n = 28 1r n = 24 1r n = 43 1r n = 15 1r

SiO2 50.09 0.62 49.82 1.07 49.95 0.94 50.87 1.07 49.37 0.66

TiO2 1.06 0.20 0.73 0.23 1.37 0.28 0.49 0.31 0.99 0.15

Al2O3 5.92 0.44 4.84 1.10 4.01 0.84 3.49 1.18 5.77 0.77

Cr2O3 0.47 0.32 0.55 0.31 0.07 0.03 0.71 0.35 0.51 0.28

FeO 5.45 0.62 5.44 0.48 10.03 0.49 5.05 1.36 5.38 0.61

MnO 0.14 0.03 0.14 0.03 0.25 0.04 0.14 0.04 0.15 0.03

MgO 16.23 0.43 16.00 0.89 13.21 0.46 16.44 1.05 16.25 0.76

CaO 19.95 0.58 20.88 0.61 20.64 0.46 21.22 0.74 19.97 0.67

NiO 0.03 0.02 0.03 0.02 b.d. 0.02 0.02 0.03 0.02

V2O3 0.09 0.03 0.06 0.02 0.10 0.03 0.05 0.03 0.10 0.03

Na2O 0.22 0.04 0.19 0.03 0.40 0.04 0.20 0.06 0.20 0.03

Total 99.69 98.72 100.07 98.72 98.75

Mg# 84.2 1.7 84.0 1.5 70.1 1.2 85.3 4.2 84.3 1.6

Sample M3 S5 S6 S7 S8

Type E-MORB N-MORB N-MORB SSZ SSZ


SiO2 50.46 0.91 50.19 1.23 50.86 1.13 50.61 3.16 52.49 0.91

TiO2 0.84 0.23 0.74 0.29 0.82 0.38 0.63 0.48 0.23 0.11

Al2O3 4.72 0.88 4.17 1.08 3.36 1.06 4.29 2.74 1.92 0.65

Cr2O3 0.29 0.20 0.56 0.46 0.33 0.16 0.42 0.41 0.72 0.26

FeO 5.67 0.74 6.78 1.68 7.18 2.15 6.55 2.53 4.63 0.87

MnO 0.15 0.03 0.18 0.05 0.18 0.06 0.16 0.05 0.13 0.03

MgO 16.09 0.58 16.52 1.18 15.34 1.32 15.77 2.21 17.62 0.70

CaO 21.02 0.69 19.03 1.10 20.88 0.49 21.03 0.68 21.03 0.36

NiO 0.03 0.02 0.02 0.01 0.01 0.01 0.03 0.02 0.04 0.02

V2O3 0.07 0.03 0.07 0.03 0.06 0.03 0.03 0.03 0.02 0.02

Na2O 0.20 0.04 0.28 0.06 0.33 0.09 0.26 0.05 0.23 0.03

Total 99.57 98.58 99.39 99.82 99.09

Mg# 83.5 2.0 81.3 4.3 79.2 6.3 81.1 8.1 87.1 2.5

CoO, ZnO, and K2O were measured but below the detection limit in all samples

b.d. below detection limit. Sample locations: A Agrilia Formation, M Mirna Group, S Sipetorrema Pillow Lava Unit


123

Low-pressure fractional crystallization of the MORB-type

rocks

The different mantle sources and different degrees of

melting of the N-MORB-type and E-MORB-type rocks are

reflected by the lack of correlation between Zr and highly

incompatible trace elements such as Nb and La (Fig. 3f). In

contrast, the good correlation between Zr and many major

elements, compatible and moderately incompatible trace

elements (Fig. 3) points to fractional crystallization as the

main process that controlled the evolution of the MORB-

type rocks. Petrographical observations demonstrate that

these rocks have experienced fractional crystallization of

Cr-spinel, clinopyroxene, and plagioclase, most likely

preceded by olivine fractionation. This crystallization

sequence is also confirmed by the decrease of MgO, Ni,

and Cr coupled with the increase of FeO with increasing Zr

(Fig. 3; Table 2). As typical for tholeiitic basalts, the

MORB-type rocks do not show an increase of SiO2 with

increasing fractionation.

The co-variation of Ni and Cr is plotted in Fig. 11, since

these elements are mainly distributed within the early mafic

minerals (i.e., olivine, Cr-spinel and clinopyroxene). The

general trend observed in this figure argues for the crys-

tallization of olivine, clinopyroxene, and possibly small

amounts of Cr-spinel, as deduced from the sharp decrease

of both Ni and Cr in the more fractionated rocks.

Clinopyroxene and plagioclase probably started crys-

tallizing almost simultaneously, given that the mafic rocks

display no or only small Eu anomalies (Fig. 5). Negative Sr

anomalies coupled with small negative Eu anomalies point

to plagioclase fractionation in several samples (e.g., sample

A10), whereas positive Sr anomalies in samples A9 and S5

Table 5 Major element composition of orthopyroxene determined by

electron microprobe

Sample S8

Type SSZ

Spots n = 2 1r

SiO2 54.35 0.66

TiO2 0.43 0.06

Al2O3 1.64 0.05

Cr2O3 0.04 0.02

FeO 13.28 0.11

MnO 0.42 0.02

MgO 28.51 0.12

CaO 1.64 0.20

NiO 0.05 0.02

V2O3 0.02 0.01

Na2O 0.05 0.01

Total 100.45

Mg# 79.3 0.2

CoO, ZnO, and K2O were measured but below the detection limit in

all samples

Sample locations: S Sipetorrema Pillow Lava Unit

Table 6 Major element composition of spinel determined by electron microprobe

Sample A1 A9 A11 S7 S8

Type E-MORB E-MORB E-MORB SSZ SSZ


SiO2 0.07 0.04 0.07 0.03 0.15 0.07 0.09 0.07 0.05 0.03

TiO2 0.60 0.07 0.52 0.11 0.49 0.09 0.35 0.14 8.35 16.10

Al2O3 22.85 0.64 14.78 2.11 15.52 1.78 6.84 1.16 6.15 3.54

Cr2O3 42.26 1.17 49.41 2.99 48.53 1.59 55.76 3.53 37.23 25.74

Fe2O3 7.07 0.44 9.00 0.98 9.22 0.38 10.52 1.98 20.16 14.86

FeO 12.37 0.10 13.80 1.02 13.69 1.66 17.82 1.20 20.70 5.71

MnO 0.17 0.01 0.18 0.03 0.22 0.02 0.27 0.02 0.28 0.04

MgO 14.99 0.01 13.22 0.84 13.40 1.29 9.77 0.84 8.02 3.77

ZnO 0.07 0.04 0.06 0.03 0.08 0.04 0.11 0.03 0.08 0.05

NiO 0.14 0.01 0.12 0.04 0.11 0.06 0.10 0.03 0.15 0.09

V2O3 0.23 0.02 0.15 0.03 0.13 0.02 0.08 0.04 0.34 0.44

CoO 0.04 0.01 0.04 0.02 0.03 0.02 0.05 0.02 0.06 0.02

CaO 0.08 0.01 0.10 0.05 0.04 0.03 b.d. b.d.

Total 100.97 101.46 101.62 101.78 101.60

Mg# 68.4 0.2 63.0 3.2 63.5 5.0 49.4 3.8 40.4 18.2

Cr# 55.4 1.4 69.2 4.3 67.8 3.2 84.5 3.0 70.6 18.1

Fe3+ was calculated from the cation sums. Na2O and K2O were measured but below the detection limit in all samples

b.d. below detection limit. Sample locations: A Agrilia Formation, S Sipetorrema Pillow Lava Unit


123

suggest plagioclase accumulation (Fig. 6). The crystalli-

zation of Fe–Ti oxides was most likely not significant,

because the Ti content increases continuously in with

increasing Zr (Fig. 3b).

N-MORB-type basalts and basaltic andesites

Barth et al. (2003) showed that the clinopyroxene com-

position of MOR-type peridotites from the Fournos Kaıtsa

sub-massif of the Othris peridotite massif can be explained

by moderate degrees (*15%) of anhydrous partial melting

of depleted mantle. These authors proposed a multistage

melting model of *4% melting in the garnet stability field

followed by *12% melting in the spinel stability field.

The superchondritic MREE/HREE ratios of the N-

MORB-type basalts point to an episode of partial melting

in the garnet stability field. MREE to HREE patterns are

steeper than the REE patterns produced by single-stage

melting in the spinel stability field. The multistage melting

model of Barth et al. (2003) followed by 30–50% crystal

fractionation reproduces the concentrations of the moder-

ately incompatible trace elements in the basaltic samples

(Fig. 5a). For samples S5 and S6 the calculated and the

measured Cr contents agree reasonably well. However,

sample A10 has much lower Cr contents than the model.

The basaltic andesite M30 requires more than 50% frac-

tionation, which is beyond the scope of the model. The

calculated LREE abundances are lower than the measured

abundances, suggesting a mantle source that is slightly less

depleted than the MORB source of Barth et al. (2003).

E-MORB-type basalts

The E-MORB-type basalts have, on average, higher Mg#,

Cr, Ni, and LREE and lower HREE than the N-MORB-type

basalts, implying a lower degree of crystal fractionation,

lower degrees of partial melting, a more enriched source,

and/or a higher proportion of melting in the garnet stability

field. Samples A1, A32, and M3 can be modeled by *4%

melting in the garnet stability field followed by *3%

Table 7 Major element composition of olivine determined by elec-

tron microprobe

Sample S7 S8

Type SSZ SSZ

Spots n = 14 1r n = 16 1r

SiO2 41.10 0.26 41.02 0.37

Al2O3 0.03 0.01 0.03 0.02

Cr2O3 0.08 0.02 0.08 0.03

FeO 9.29 1.48 8.64 1.63

MnO 0.16 0.03 0.15 0.04

NiO 0.35 0.06 0.38 0.06

MgO 49.78 1.10 50.32 1.25

CaO 0.18 0.06 0.15 0.06

CoO 0.03 0.01 0.03 0.02

Total 101.05 100.84

Fo 90.5 1.5 91.2 1.7

TiO2, V2O3, ZnO, Na2O, and K2O were measured but below the

detection limit in all samples

Sample locations: S Sipetorrema Pillow Lava Unit

MO

RB

IAT

BON

MO

RB

IAT

BON

MORB

IATBON

clinopyroxeneMg# > 80N-MORB-type

S5S6

clinopyroxeneMg# > 80E-MORB-type

Na O20 25 50 75 100

TiO2

0

25

50

75

SiO /1002

25

50

75

100

A1A9A11A32M3

clinopyroxeneMg# > 80SSZ-type

S7S8

Na O20 25 50 75 100

TiO2

0

25

50

75

SiO /1002

25

50

75

100

Na O20 25 50 75 100

TiO2

0

25

50

75

SiO /1002

25

50

75

100

(a)

(b)

(c)

Fig. 7 TiO2-Na2O-SiO2/100 (wt%) discrimination diagram for clin-

opyroxenes from mafic rocks from Othris. a Normal mid-ocean ridge-

type basalts and basaltic andesites, b enriched mid-ocean ridge-type

basalts, c supra-subduction zone-type cumulates. For clarity, only

clinopyroxenes with Mg#[80 are shown. Abbreviations: MORB mid-

ocean ridge basalt, IAT island-arc tholeiites, BON boninites. Modified

after Beccaluva et al. (1989)


123

melting in the spinel stability field (Fig. 5b). Samples A9

and A11 have lower REE contents and lower LREE/HREE

ratios than samples A1, A32, and M3, indicating higher

degrees of melting. The trace element contents of samples

A9 and A11 can be reproduced by *8% melting in the

garnet stability field followed by *12% melting in the

spinel stability field. The measured Cr contents suggest

10–50% crystal fractionation. However, the low HREE con-

tents suggest lower degrees of fractionation (\20%). Since

the calculated LREE abundances are considerably lower than

the measured abundances, an enriched mantle source is

required to explain the observed trace element abundances.

SSZ-type cumulates

The very high MgO, Cr, and Ni contents and high Mg# of

samples S7 and S8 suggest a cumulate origin. Bickle (1982)

discussed an approach to the estimate the original melt

composition in the case of lavas that accumulated pheno-

crysts. Based on experimental data of Fe–Mg partitioning

between olivine and melt, Bickle (1982) proposed a distri-

bution coefficient KD = 0.314 for Mg-rich lavas and

forsteritic olivine. The most magnesian olivines (Fo94) in

samples S7 and S8 would be in equilibrium with melts

containing *30 wt% MgO (calculated on a volatile-free

basis). This is a much higher MgO content than inferred for

the melt composition of komatiitic rocks from the Agrilia

Formation with broadly similar bulk composition (12–

17 wt% MgO; Cameron and Nisbet 1982; Economou-Elio-

poulos and Parakevopoulos 1989). Considering the low SiO2

(46–47 wt%) and Al2O3 (6 wt%) contents of the calculated

melt, we do not judge the calculated melt composition to be

trustworthy. A likely source of error is the forsterite content

of olivine, because the maximum forsterite content of olivine

samples S7 and S8 is higher than that in olivines in komatiitic

rocks from the Agrilia Formation (Fo90–92; Cameron and

Nisbet 1982; Economou-Eliopoulos and Parakevopoulos

1989) and in olivines in peridotites from the Othris Ophiolite

(Fo90–92; eTable 5 and Barth et al. 2003). Python et al. (2007)

argued that the forsterite content of olivine can increase

during high temperature hydrothermal alteration. Using the

average olivine composition (Fo91) of samples S7 and S8, we

calculate a melt composition of 20–22 wt% MgO, 6–7 wt%

CaO, 12 wt% FeO, 49 wt% SiO2, and 8–9 wt% Al2O3. The

calculated melt has MgO, CaO, SiO2, and Al2O3 concen-

trations broadly similar to but considerably higher FeO

concentrations than the experimental boninitic liquids of

Falloon and Danyushevsky (2000).

The fluid-induced refertilization-hydrous melting model

for SSZ-type rocks assumes a mantle source that has been

depleted by 9% anhydrous near-fractional melting in the

spinel stability field. The concentrations of moderately

incompatible trace elements of samples S7 and S8 can be

modeled by *10% hydrous melting induced by addition of

a fluid with the maximum trace element content of Bizimis

et al. (2000), followed by 30–50% olivine addition

(Fig. 5c). However, the modeled concentrations of highly

incompatible trace elements are considerably lower than

the measured concentrations. The most likely reasons for

this discrepancy are (a) the mantle source was less depleted

in highly incompatible trace elements than the model

source; (b) the fluid component had higher concentrations

of highly incompatible trace elements than the model

composition of Bizimis et al. (2000); and/or (c) mixing of a

depleted SSZ-type magma and a more enriched melt.

Geological implications

In agreement with the earlier studies of Pearce et al. (1984)

and Photiades et al. (2003) our melt modeling suggests that

spinel

Mg#304050607080

Cr#

0

20

40

60

80

100

A1A9A11S7S8Othris peridotites

E-MORB-type

SSZ-type

}}

Fig. 8 Diagram of Cr# [100 9 Cr/(Cr + Al)] versus Mg# in spinel

from mafic rocks from Othris compared to spinel from Othris

peridotites (eTable 4 and Barth et al. 2003)

olivine

Fo88 89 90 91 92 93 94 95

NiO

[wt%

]

0.20

0.25

0.30

0.35

0.40

0.45

0.50

S7S8Othris peridotites

SSZ-type}

Fig. 9 Plot of forsterite content versus NiO [in wt%] in olivine from

mafic rocks from Othris compared to olivine from Othris peridotites

(eTable 5 and Barth et al. 2003)


123

the crustal section of Othris Ophiolite records two distinct

melting regimes, i.e., anhydrous MOR-type and hydrous

SSZ-type melting. The occurrence of both N- and

E-MORB-type lavas suggests that the mantle generating

the lavas of the Othris Ophiolite must have been hetero-

geneous on a comparatively fine scale. The E-MORB-type

magmas were erupted without significant mixing with N-

MORB magmas. This observation is consistent with an

origin in a slow-spreading system, since these heteroge-

neities may not have survived in a fast-spreading system

that is underlain by a steady-state melt lens (e.g., Sinton

and Detrick 1992). Thus, the N-MORB and E-MORB

magmatism of the Othris Ophiolite might have been

comparable to the N- to E-MORB magmatism observed on

Macquarie Island, which is the result of a waning spreading

system (Varne et al. 2000).

The N-MORB-type basalts from the Agrilia Formation

and the Sipetorrema Pillow Lava unit are consistent with

about *15% anhydrous partial melting of a depleted

MORB source, in excellent agreement with the degrees of

partial melting inferred for the Othris MOR-type perido-

tites (Barth et al. 2003, 2008). The MOR-type peridotites

from the Fournos Kaıtsa sub-massif and the western

Katachloron sub-massif are moderately depleted and simi-

lar to abyssal peridotites.

The E-MORB-type basalts from the Agrilia Formation

and the Mirna Group record between 7 and 20% anhydrous

partial melting of an enriched mantle source. The minimum

degree of melting inferred for the MOR-type peridotites

(13%; Barth et al. 2008) is not as low as the minimum

degree of melting inferred for the E-MORB-type basalts.

Furthermore, Barth et al. (2008) did not deduce an enriched

mantle source for the MOR-type peridotites. This implies

that the complementary residue of the E-MORB-type

basalts is not exposed or has not been sampled in the Othris

peridotite massif, possibly because such comparatively

undepleted peridotites would probably occur at deeper

stratigraphic levels that are not exposed or have not been

emplaced.

The SSZ-type cumulates from the Sipetorrema Pillow

Lava unit are products of hydrous melting of a previously

depleted mantle source. The Othris peridotites from the

Metalleio, Eretria, and eastern Katachloron sub-massif,

which record evidence of SSZ-type mantle, i.e., they are

Table 8 Partition coefficients, initial mineralogy, melting proportions, MORB source and fluid composition in ppm (lg/g)

Olivine opx cpx Spinel Garnet Plag MORB Fluid

Source Maximum

Ti 0.007 0.12 0.3 0.07 0.29 0.06 927 436

Cr 0.7 4 4 70 2 0.1 2630

Ni 10 1.1 2.6 5 5 0.1

Y 0.006 0.07 0.46 0.002 2.8 0.03 3.44

Zr 0.003 0.05 0.12 0.04 0.27 0.003 6.195 365

Sr 0.01 0.04 0.06 0.0006 0.0025 1.6 12.93 1930

La 0.000007 0.0005 0.0536 0.0006 0.0016 0.16 0.161 18.6

Ce 0.00001 0.0009 0.09 0.0006 0.005 0.1 0.5376 48

Pr 0.00003 0.004 0.13 0.0007 0.018 0.09 0.108 5.88

Nd 0.00007 0.009 0.17 0.0008 0.052 0.07 0.7375 22.2

Sm 0.0007 0.02 0.29 0.0008 0.25 0.06 0.304 4.9

Eu 0.001 0.03 0.35 0.0009 0.4 0.06 0.118 2.85

Gd 0.0012 0.04 0.4 0.0009 0.8 0.05 0.417 3.46

Dy 0.004 0.07 0.45 0.0015 2.2 0.03 0.559 0.5

Er 0.01 0.07 0.47 0.003 3.6 0.02 0.381 0.3

Yb 0.005 0.08 0.49 0.0015 6.6 0.01 0.392 0.3

Xa 0.55 0.25 0.18 0.02

Pa dry -0.20 0.43 0.72 0.05

Pa hydrous -0.10 0.52 0.56 0.02

Xa 0.57 0.21 0.16 0.06

Pa dry 0.08 -0.19 0.81 0.30

Fractionation 0.60 0.38 0.02

Partition coefficients from Kelemen et al. (1993), Johnson (1998), and from the GERM database (http://www.earthref.org/GERM/). Xa min-

eralogy of the MORB source (Johnson et al. 1990). Pa dry melting mode for dry melting. Spinel melting modes from Baker and Stolper (1994);

garnet peridotite melt modes from Walter (1998). Pa hydrous melting mode for hydrous melting in an island arc environment from Bizimis et al.

(2000). Composition of the fluid component from Bizimis et al. (2000)


123

http://www.earthref.org/GERM/

highly depleted and are similar to peridotites from the Izu-

Bonin-Mariana forearc (Barth et al. 2008), may be the

complementary residua to the SSZ-type cumulates.

In order to test if the crustal section of the Othris

Ophiolite is genetically linked to the mantle section we

need to compare the melting conditions of the mafic rocks

of the Mirna Group, including, the Sipetorrema Pillow

Lava unit, to the melting conditions of the Othris peridotite

massif as the Othris peridotite massif is part of the Mirna

Group (Smith et al. 1975). The excellent agreement

between the melting conditions inferred from the

N-MORB-type basalts and the MOR-type peridotites as

well as the broadly complementary nature of the boninitic

cumulates and the SSZ-type peridotites suggest that the


section. Hence, our modeling results support the interpre-

tation that both the crustal section and the mantle section of

the Othris Ophiolite formed in an oceanic environment,

probably in an infant arc or in a forearc (see below).

Accordingly, the Othris Ophiolite can be classified as a

Mediterranean-type ophiolite (Dilek 2003).

Furthermore, our melt modeling implies that the mantle

sources and melting regimes of the crustal section of the

Othris complex evolved from anhydrous MOR-type melting

during the Middle Triassic–Middle Jurassic opening of the

Mirdita-Pindos oceanic basin, as recorded by the Triassic

Agrilia Formation, to both MOR- and fluid-induced SSZ-

type melting during the closure of the Mirdita-Pindos oce-

anic basin by intra-oceanic subduction during the Middle–

Late Jurassic, as recorded by the Jurassic Mirna Group

(including the Sipetorrema Pillow Lava unit). An important

trait of the Sipetorrema Pillow Lava unit is the close spatial

association of MORB-type and SSZ-type rocks. The close

association of different lava suites has also been reported

for other ophiolites in the Mirdita-Subpelagonian Zone (see

review by Robertson 2002). At several localities including

the Pindos ophiolite in the NW of Othris and in the Alba-

nian ophiolites further north lavas ranging from MORB to

island-arc tholeiites (IAT) and boninites are spatially jux-

taposed or interlayered (Capedri et al. 1980; Pearce et al.

1984; Jones and Robertson 1991; Bebien et al. 2000; Bor-

tolotti et al. 2002; Hoeck et al. 2002; Saccani and Photiades

2004). Taken together, these observations indicate that two

or more distinct melting regimes coexisted spatially and/or

temporally in a relatively restricted sector across an intra-

oceanic subduction setting.

The geological and petrological characteristics shown by

the Othris Ophiolite are compatible with the intra-oceanic

thrusting model for the Hellenic-Dinaric ophiolites pro-

posed by Spray et al. (1984), Bebien et al. (2000),

Insergueix-Filippi et al. (2000), and Barth et al. (2008).

According to this model, the Albanide-Hellenide ophiolites

formed by ridge collapse (Fig. 12), i.e., forced initiation of

subduction at or near a mid-ocean ridge (e.g., Boudier and

Coleman 1981; Boudier et al. 1988). The kinematic

numerical models of Insergueix-Filippi et al. (2000)

Y [ppm]

Cr

[ppm

]

101 10 100

100

1000

10000

MORBsource

residue

fractional

crystallization

accumulatedpartial melt

Othris peridotitesMid-Atlantic Ridge

SSZ-type

N-MORB-typeE-MORB-type

bulk rock

4%8%12%

4%8%

20%20%

40%30%

M30

A10

anhydrous near-fractional melting

Fig 10 Cr versus Y diagram for mafic rocks from Othris compared to

Othris peridotites and MORB from the Mid-Atlantic ridge. Also

shown are the results obtained for a melting model including 4% of

near-fractional melting (incremental batch melting at 0.1% incre-

ments) in the garnet stability field followed by near-fractional melting

in the spinel stability field (solid lines). After 4% of partial melting in

the garnet stability field, a new source modal mineralogy was

calculated to account for the subsolidus garnet to spinel phase

transition. MORB source composition, melting and residue paths for

near-fractional melting, and peridotite compositions are from Barth

et al. (2003, 2008). Numbers along the solid lines are percent melting.

Dashed line shows the Rayleigh fractionation trend of a solid

assemblage of 60% olivine + 38% clinopyroxene + 2% spinel (see

text for details). Numbers along the dashed line are percent crystal

fractionation

bulk rock

Cr [ppm]0 500 1000 1500 2000 2500

Ni [

ppm

]

0

500

1000

1500

2000

SSZ-typeMid-Atlantic Ridge

N-MORB-typeE-MORB-type

primarymelts

olivineaddition

olivine

cpx

opxplagfractionation

spinel

Fig. 11 Ni versus Cr diagram for mafic rocks from Othris compared

to MORB from the Mid-Atlantic ridge. Fractional crystallization

trends for olivine, clinopyroxene (cpx), orthopyroxene (opx), spinel,

and plagioclase (plag) based on the partition coefficients in Table 8

are shown. The arrow indicates the effect of olivine addition to a

primary melt using the average olivine composition of samples S7

and S8. Primary melt composition is taken from Liang and Elthon

(1990)


123

demonstrate that preservation of high temperatures in the

mantle wedge explains the occurrence of boninitic mag-

matism at the earliest stages of subduction initiation, which

is partly contemporaneous with the decline of MORB

magmatism and the initiation of arc magmatism. In the case

of localized fluid flux from the slab, hydrous partial melting

of previously depleted peridotites may be contained to

limited areas, explaining the close proximity of SSZ-type

and MOR-type rocks in the Sipetorrema Pillow Lava unit.

As an alternative, Dilek and Flower (2003), Flower and

Dilek (2003), Saccani et al. (2004), and Beccaluva et al.

(2005) recently proposed a tectonic model of western

Pacific-style arc-trench rollback. The main difference

between these two models is an episode of slab rollback

and associated inter-arc or back-arc opening in the arc-

trench rollback model, causing extension in the upper plate,

whereas in the intra-oceanic thrusting model the upper

plate is under compression. As discussed by Barth et al.

(2008), a western Pacific-style tectonic model is unlikely to

be applicable to the Mirdita-Pindos oceanic basin, because

the geodynamic boundary conditions such as the density

differences between subducting and overriding plates are

very different for subduction initiation in the western

Pacific in the Eocene and in the Mirdita-Pindos oceanic

basin, a comparatively short-lived and narrow marginal

basin.

On the other hand, Benoit et al. (1999), Python and

Ceuleneer (2003), and Nonnotte et al. (2005) suggested

that N-MORB-type and depleted hydrous magmas may

coexist in a single tectonic setting. The depleted hydrous

magmas may be generated at a mid-ocean ridge by shallow

re-melting of partly hydrated peridotites residual after

MORB extraction.

Irrespective of the exact tectonic model, the well-pre-

served ophiolite exposures found from Serbia to Greece

along approximately 1,000 km correspond to infant arc

settings, suggesting that large-scale obduction processes

require the creation of new intra-oceanic subduction zones,

whether spontaneous or induced (Stern 2004; Gurnis et al.

2004). Recently, Agard et al. (2007) proposed a mechanism

in which large-scale obduction events are triggered by

intraplate instabilities resulting from sharp plate accelera-

tions, possibly in response to superplume events. These

authors demonstrated that the inception of the Jurassic

obduction, as testified by ages obtained for metamorphic

soles, also developed during a period of high convergence

velocities.

Conclusions

The crustal section of the Othris Ophiolite contains three

geochemically distinct groups of mafic rocks: N-MORB-

type basalts and basaltic andesites, E-MORB-type basalts,

and boninitic cumulates, indicative of both anhydrous

MOR-type and fluid-induced SSZ-type melting regimes.

Starting in the Triassic, a MORB-type oceanic lithosphere

was generated between the Korabi-Pelagonian and the

Apulian microcontinents. The occurrence of both N-

MORB- and E-MORB-type basalts in the Agrilia For-

mation suggests that the MORB source mantle was

heterogeneous. When oceanic extension rapidly changed

to convergence, intra-oceanic subduction started in the

Middle Jurassic and resulted in the almost contempora-

neous production of MORB-type and boninitic magmas,

as observed in the mafic rocks of the Mirna Group. The

excellent agreement between the melting conditions

inferred from the N-MORB-type basalts and the MOR-

type peridotites in Othris as well as the broadly com-

plementary nature of the boninitic cumulates and the

SSZ-type peridotites found in Othris suggest that the


section.

Acknowledgments We would like to express our gratitude to Arjan

Dijkstra and Gareth Davies for help with field work. This work was

funded in part by ISES grant 6.2.3 Dinaric-Hellenic Ophiolites.

MORB magmatism

Island Arc magmatism

Boninitic magmatism

ascendingasthenospheric

flowMelting zones:

lithosphere formedat a mid-ocean ridge

high-Ti intermediate to low-TiMORB boniniteIAT

lithosphere formedat a mid-ocean ridge

Jurassic: subduction initiation by intra-oceanic thrusting

slivers ofE-MORBsource

depleted mantle(N-MORB source)

Fig. 12 Cartoon (not to scale) showing the genesis of the Othris

Ophiolite by intra-oceanic subduction during the Middle–Late

Jurassic (modified from Insergueix-Filippi et al. 2000, and Barth

et al. 2008). The Mirdita-Pindos oceanic basin opened during the

Middle Triassic–Middle Jurassic. The subduction zone probably

originated at (or close to) the mid-ocean ridge axis of the Mirdita-

Pindos ocean (ridge collapse). The preservation of high temperatures

in the mantle wedge favors the setting of short-lived boninitic

magmatism in the earliest stages of subduction initiation, which are

partly contemporaneous with a progressive extinction of MORB

magmatism and initiation of arc magmatism. Enriched E-MORB

source material may occur as blobs or slivers in the depleted

N-MORB source mantle. The intermediate-Ti magmatism can result

from tapping of less depleted mantle but also from mixing of high-Ti

and low-Ti magmatism due to the close proximity of the respective

melting zones in the mantle wedge


123

Discussions with Dejan Prelevic and comments by two anonymous

reviewers helped to improve the manuscript. Tim Grove is thanked

for his constructive editorial handling.

References

Agard P, Jolivet L, Vrielynck B, Burov E, Monie P (2007) Plate

acceleration: the obduction trigger? Earth Planet Sci Lett

258:428–441

Baker MB, Stolper EM (1994) Determining the composition of high-

pressure mantle melts using diamond aggregates. Geochim

Cosmochim Acta 58:2811–2827

Barnes SJ, Roeder PL (2001) The range of spinel compositions in

terrestrial mafic and ultramafic rocks. J Petrol 42:2279–2302

Barth MG, Mason PRD, Davies GR, Dijkstra AH, Drury MR (2003)

Geochemistry of the Othris Ophiolite, Greece: evidence for

refertilization? J Petrol 44:1759–1785

Barth MG, Mason PRD, Davies GR, Drury MR (2008) The Othris

Ophiolite, Greece: a snapshot of subduction initiation at a mid-

ocean ridge. Lithos 100:234–254

Bebien J, Dimo-Lahitte A, Vergely P, Insergueix-Filippi D, Dupeyrat

L (2000) Albanian ophiolites I—magmatic and metamorphic

processes associated with the initiation of a subduction. Ofioliti

25:39–45

Beccaluva L, Macciotta G, Piccardo GB, Zeda O (1989) Clinopy-

roxene composition of ophiolite basalts as petrogenetic indicator.

Chem Geol 77:165–182

Beccaluva L, Coltorti M, Saccani E, Siena F (2005) Magma

generation and crustal accretion as evidenced by supra-subduc-

tion ophiolites of the Albanide–Hellenide Subpelagonian zone.

Isl Arc 14:551–563

Benoit M, Ceuleneer G, Polve M (1999) The remelting of

hydrothermally altered peridotite at mid-ocean ridges by intrud-

ing mantle diapirs. Nature 402:514–518

Bickle MJ (1982) The magnesium content of komatiitc liquids. In:

Arndt NT, Nisbet EG (eds) Komatiites. Allen & Unwin, London,

pp 479–494

Bizimis M, Salters VJM, Bonatti E (2000) Trace and REE content of

clinopyroxenes from supra-subduction zone peridotites. Impli-

cations for melting and enrichment processes in island arcs.

Chem Geol 165:67–85

Bortolotti V, Marroni M, Pandolfi L, Principi G, Saccani E (2002)

Interaction between mid-ocean ridge and subduction magmatism

in Albanian ophiolites. J Geol 110:561–576

Boudier F, Coleman RG (1981) Cross section through the peridotite

in the Samail ophiolite, southeastern Oman mountains.

J Geophys Res 86:2573–2592

Boudier F, Ceuleneer G, Nicolas A (1988) Shear zones, thrusts and

related magmatism in the Oman ophiolite: Initiation of thrusting

on an ocean ridge. Tectonophysics 151:275–296

Cameron WE, Nisbet EG (1982) Phanerozoic analogues of komatiitic

basalts. In: Arndt NT, Nisbet EG (eds) Komatiites. Allen &

Unwin, London, pp 29–50

Capedri S, Venturelli G, Bocchi G, Dostal J, Garuti G, Rossi A (1980)

The geochemistry and petrogenesis of an ophiolitic sequence

from Pindos, Greece. Contrib Mineral Petrol 74:189–200

Coleman RG (1977) Ophiolites. Springer, Heidelberg, p 229

Dijkstra A, Drury MR, Vissers RLM (2001) Structural petrology of

plagioclase peridotites in the West Othris Mountains (Greece):

melt impregnation in mantle lithosphere. J Petrol 42:5–24

Dijkstra AH, Barth MG, Drury MR, Mason PRD, Vissers RLM

(2003) Diffuse porous melt flow and melt-rock reaction in the

mantle lithosphere at a slow-spreading ridge: A structural

petrology and LA-ICP-MS study of the Othris Peridotite Massif

(Greece). Geochem Geophys Geosys 4:8613. doi:

10.1029/2001GC000278

Dilek Y (2003) Ophiolite concept and its evolution. In: Dilek Y,

Newcomb S (eds) Ophiolite concept and the evolution of

geological thought. Geol Soc Am Special Paper, vol 373.

Boulder, Colorado, pp 1–16

Dilek Y, Flower MFJ (2003) Arc-trench rollback and forearc

accretion: 2. A model template for ophiolites in Albania,

Cyprus, and Oman. In: Dilek Y, Robinson PT (eds) Ophiolites

in Earth history. Geol Soc London Special Publication, vol 208,

pp 43–68

Dilek Y, Shallo M, Furnes H (2005) Rift-drift, seafloor spreading, and

subduction tectonics of Albanian ophiolites. Int Geol Rev

47:147–176

Economou-Eliopoulos M, Parakevopoulos GM (1989) Platinum-

group elements and gold in komatiitic rocks from the Agrilia

formation, Othrys ophiolite complex, Greece. Chem Geol

77:149–158

Falloon TJ, Danyushevsky LV (2000) Melting of refractory mantle at

1.5, 2 and 2.5 GPa under anhydrous and H2O-undersaturated

conditions: implications for the petrogenesis of high-Ca boni-

nites and the influence of subduction components on mantle

melting. J Petrol 41:257–283

Ferriere J (1985) Nature et developpement des ophiolites Helleniques

du secteur Othrys-Pelion. Ofioliti 10:255–278

Flower MFJ, Dilek Y (2003) Arc-trench rollback and forearc

accretion: 1. A collision-induced mantle flow model for Tethyan

ophiolites. In: Dilek Y, Robinson PT (eds) Ophiolites in Earth

history. Geol Soc London Special Publication, vol 218, pp 21–41

Gurnis M, Hall C, Lavier L (2004) Evolving force balance during

incipient subduction. Geochem Geophys Geosys 5, Q07001. doi:

10.1029/2003GC000681

Hoeck V, Koller F, Meisel T, Onuzi K, Kneringer E (2002) The

jurassic south Albanian ophiolites: MOR-vs SSZ-type ophiolites.

Lithos 65:143–164

Hynes A (1974) Notes on the petrology of some ophiolites, Othris

mountains, Greece. Contrib Mineral Petrol 46:233–239

Hynes AJ, Nisbet EG, Smith GA, Welland MJP, Rex DC (1972)

Spreading and emplacement ages of some ophiolites in the

Othris region (eastern central Greece). Z Deutsch Geol Ges

123:455–468

Insergueix-Filippi D, Dupeyrat L, Dimo-Lahitte A, Vergely P, Bebien

J (2000) Albanian ophiolites II—model of subduction zone

infancy at a mid-ocean ridge. Ofioliti 25:47–53

Johnson KTM (1998) Experimental determination of partition

coefficients for rare earth and high-field-strength elements

between clinopyroxene, garnet, and basaltic melt at high

pressures. Contrib Mineral Petrol 133:60–68

Johnson KT, Dick HJB, Shimizu N (1990) Melting in the oceanic

upper mantle: an ion microprobe study of diopsides in abyssal

peridotites. J Geophys Res 95:2661–2678

Jones G, Robertson AHF (1991) Tectono-stratigraphy and evolution

of the Mesozoic Pindos ophiolite and related units, Northwestern

Greece. J Geol Soc London 148:267–288

Kelemen PB, Shimizu N, Dunn T (1993) Relative depletion of

niobium in some arc magmas and the continental crust:

partitioning of K, Nb, La and Ce during melt/rock reaction in

the upper mantle. Earth Planet Sci Lett 120:111–134

Liang Y, Elthon D (1990) Evidence from chromium abundances in

mantle rocks for extraction of picrite and komatiite melts. Nature

343:551–553

McDonough WF, Sun S-S (1995) Composition of the earth. Chem

Geol 120:223–253

Menzies M (1976) Rifting of a Tethyan continent—rare earth

evidence of an accreting plate margin. Earth Planet Sci Lett

28:427–438


123

http://dx.doi.org/10.1029/2001GC000278

http://dx.doi.org/10.1029/2003GC000681

Menzies M, Allen C (1974) Plagioclase lherzolite-residual mantle

relationships within two eastern Mediterranean ophiolites. Con-

trib Mineral Petrol 45:197–213

Moores EM (1982) Origin and emplacement of ophiolites. Rev

Geophys Space Phys 20:735–760

Nehring F, Jacob DE, Barth MG, Foley SF (2008) Laser-ablation ICP-

MS analysis of siliceous rock glasses fused on an iridium strip

heater using MgO dilution. Microchim Acta 160:153–163

Nicolas A (1989) Structures of ophiolites and dynamics of oceanic

lithosphere. Petrology and structural geology.. Kluwer, Dordr-

echt, p 367

Nonnotte P, Ceuleneer G, Benoit M (2005) Genesis of andesitic-

boninitic magmas at mid-ocean ridges by melting of hydrated

peridotites: geochemical evidence from DSDP Site 334 gab-

bronorites. Earth Planet Sci Lett 236:632–653

Paraskevopoulos GM, Economou MI (1986) Komatiite-type ultra-

mafic lavas from the Agrilia formation, Othrys Ophiolite

complex, Greece. Ofioliti 11:293–304

Paraskevopoulos GM, Economou-Eliopoulos M (1997) Geochemistry

and geotectonic setting of basaltic rocks from the Othrys

ophiolite complex. Ann Geol Pays Hell 37:601–612

Pearce JA, Lippard SJ, Roberts S (1984) Characteristics and tectonic

significance of supra-subduction zone ophiolites. In: Kokelaar

PB, Howells MF (eds) Marginal Basin Geology. Geol Soc

London Special Publication, vol 16, pp 77–94

Pearce NJG, Perkins WT, Westgate JA, Gorton MP, Jackson SE, Neal

CR, Chenery SP (1997) A compilation of new and published

major and trace element data for NIST SRM 610 and NIST SRM

612 glass reference materials. Geostand Newsl 21:115–144

Photiades A, Saccani E, Tassinari R (2003) Petrogenesis and tectonic

setting of volcanic rocks from the subpelagonian ophiolitic

melange in the Agoriani area (Othrys, Greece). Ofioliti 28:121–

135

Python M, Ceuleneer G (2003) Nature and distribution of dykes and

related melt migration structures in the mantle section of the

Oman ophiolite. Geochem Geophys Geosys 4:8612. doi:

10.1029/2002GC000354

Python M, Ceuleneer G, Ishida Y, Barrat J-A, Arai S (2007) Oman

diopsidites: a new lithology diagnostic of very high temperature

hydrothermal circulation in mantle peridotite below oceanic

spreading centres. Earth Planet Sci Lett 255:289–305

Rampone E, Piccardo GB (2000) The ophiolite-oceanic lithosphere

analogue: New insights from the Northern Apennines (Italy). In:

Dilek Y, Moores EM, Elthon D, Nicolas A (eds) Ophiolites and

oceanic crust: New insights from field studies and the Ocean

Drilling Program. Geol Soc Am Special Paper, vol 349. Boulder,

Colorado, pp 21–34

Rassios A, Konstantopoulou G (1993) Emplacement tectonism and

the position of chrome ores in the Mega Isoma peridotites, SW

Othris, Greece. Bull Geol Soc Greece 28:463–474

Rassios A, Smith AG (2000) Constraints on the formation and

emplacement age of western Greek ophiolites (Vourinos, Pindos,

and Othris) inferred from deformation structures in peridotites.

In: Dilek Y, Moores E, Elthon D, Nicolas A (eds) Ophiolites and

oceanic crust: new insights from field studies and the ocean

drilling program. Geol Soc Am Special Paper, vol 349. Boulder,

Colorado, pp 473–484

Robertson AHF (2002) Overview of the genesis and emplacement of

Mesozoic ophiolites in the eastern Mediterranean Tethyan

region. Lithos 65:1–67

Robertson AHF, Karamata S (1994) The role of subduction-accretion

processes in the tectonic evolution of the Mesozoic Tethys in

Serbia. Tectonophysics 234:73–94

Robertson A, Shallo M (2000) Mesozoic-tertiary tectonic evolution of

Albania in its regional Eastern Mediterranean context. Tectono-

physics 316:197–254

Robertson AHF, Clift PD, Degnan PJ, Jones G (1991) Palaeogeo-

graphic and palaeotectonic evolution of the Eastern

Mediterranean Neotethys. Palaeogeogr Palaeoclimatol Palaeo-

ecol 87:289–343

Saccani E, Photiades A (2004) Mid-ocean ridge and supra-subduction

affinities in the Pindos ophiolites (Greece): implications for

magma genesis in a forearc setting. Lithos 73:229–253

Saccani E, Beccaluva L, Coltorti M, Siena F (2004) Petrogenesis and

tectono-magmatic significance of the Albanide-Hellenide Sub-

pelagonian ophiolites. Ofioliti 29:75–93

Shallo M, Dilek Y (2003) Development of the ideas on the origin and

of Albanian ophiolites. In: Dilek Y, Newcomb S (eds) Ophiolite

concept and the evolution of geochemical thought. Geol Soc Am

Special Paper, vol 373. Boulder, Colorado, pp 351–363

Shervais JW (1982) Ti-V plots and the petrogenesis of modern and

ophiolitic lavas. Earth Planet Sci Lett 59:101–118

Sinton JM, Detrick RS (1992) Mid-ocean ridge magma chambers.

J Geophys Res 97:197–216

Smith AG (1993) Tectonic significance of the Hellenic-Dinaric

ophiolites. In: Prichard HM, Alabaster T, Harris NBW, Neary

CR (eds) Magmatic processes and plate tectonics. Geol Soc

London Special Publication, vol 76. Blackwell, Oxford,

pp 213–243

Smith AG, Rassios A (2003) The evolution of ideas for the origin and

emplacement of the western Hellenic ophiolites. In: Dilek Y,

Newcomb S (eds) Ophiolite concept and the evolution of

geological thought. Geol Soc Am Special Publication, vol 373.

Boulder, Colorado, pp 337–350

Smith AG, Spray JG (1984) A half-ridge transform model for the

Hellenic-Dinaric ophiolites. In: Dixon JE, Robertson AHF (eds)

The geological evolution of the Eastern Mediterranean. Geol Soc

London Special Publication, vol 17. Blackwell, Oxford, pp 629–

644

Smith AG, Hynes AJ, Menzies M, Nisbet EG, Price I, Welland MJ,

Ferriere J (1975) The stratigraphy of the Othris Mountains,

eastern central Greece: a deformed Mesozoic continental margin

sequence. Eclogae Geol Helv 68:463–481

Spray JG, Bebien J, Rex DC, Roddick JC (1984) Age constraints on

the igneous and metamorphic evolution of the Hellenic-Dinaric

ophiolites. In: Dixon JE, Robertson AHF (eds) The geological

evolution of the Eastern Mediterranean. Geol Soc London

Special Publication, vol 17. pp 619–627

Stampfli GM, Borel GD (2002) A plate tectonic model for the

Paleozoic and Mesozoic constrained by dynamic plate bound-

aries and restored synthetic oceanic isochrons. Earth Planet Sci

Lett 196:17–33

Staudigel H (2003) Hydrothermal alteration processes in the oceanic

crust. In: Holland HD, Turekian KK (eds) Treatise on geochem-

istry, vol 3, pp 511–535

Stern RJ (2004) Subduction initiation: spontaneous and induced.

Earth Planet Sci Lett 226:275–292

Sun S-S, McDonough WF (1989) Chemical and isotopic systematics

of ocean basalts: implications for mantle composition and

processes. In: Saunders AD, Norry MJ (eds) Magmatism in the

ocean basins. Geol Soc London Special Publication, vol 42,

pp 313–345

Varne R, Brown AV, Falloon T (2000) Macquarie Island: its geology

and structural history, and the timing and tectonic setting of its

N-MORB to E-MORB magmatism. In: Dilek Y, Moores EM,

Elthon D, Nicolas A (eds) Ophiolites and oceanic crust: new

insights from field studies and the ocean drilling program, vol

349. Geological Society of America, Boulder, pp 301–320

Walter MJ (1998) Melting of garnet peridotite and the origin of

komatiite and depleted lithosphere. J Petrol 39:29–60


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