Manuscript submitted to Journal of Petrology
Draft Manuscript for Review
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Ultra-refractory domains in the oceanic mantle lithosphere sampled as
mantle xenoliths at ocean islands
Manuscript ID:
Date Submitted by the
Author:
Original Manuscript
n/a
Simon, Nina; University of Oslo, Physics of Geological Processes
Neumann, Else-Ragnhild; Physics of Geological Processes
Bonadiman, C.; Università Di Ferrara
Coltorti, Massimo; Università Di Ferrara
Delpech, Guillaume; IDES, Université Orsay-Paris Sud, Faculté des
Sciences
Grégoire, Michel; Observatoire Midi-Pyrénées
Widom, Elisabeth; Miami University
Keyword:
ew
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Complete List of Authors:
JPET-Jun-07-0094.R2
er
Manuscript Type:
Journal of Petrology
Pe
Journal:
mantle xenoliths, ocean islands, recycled oceanic mantle
lithosphere, ultra-depleted mantle
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Ultra-refractory domains in the oceanic mantle lithosphere
sampled as mantle xenoliths at ocean islands
NINA S.C. SIMON, ELSE-RAGNHILD NEUMANN
Physics of Geological Processes, University of Oslo, P.O.BOX 1048 Blindern, N-0316 Oslo,
Norway; tel: +47-22856922, fax: +47-22855101, e-mail: n.s.c.simon@fys.uio.no,
e.r.neumann@geo.uio.no
CONSTANZA BONADIMAN, MASSIMO COLTORTI
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Earth Science Department, University of Ferrara, Via Saragat, 1, 44100 Ferrara, Italy; e-mail:
bdc@unife.it, clt@unife.it
GUILLAUME DELPECH
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UMR CNRS 8148 IDES, "Interactions et Dynamique des Environments de Surface", Faculté
des Sciences, Université Orsay-Paris Sud, BÂT.504, 91405 Orsay Cedex, France; e-mail:
delpech@u-psud.fr
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MICHEL GRÉGOIRE
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Observatoire Midi-Pyrénées, UMR 5562-DTP, OMP, Université Toulouse III, Avenue E.
Belin, 31400 Toulouse, France; e-mail: Michel.Gregoire@dtp.obs-mip.fr
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Manuscript submitted to Journal of Petrology
ELISABETH WIDOM
Department of Geology, Miami University, 114 Shideler Hall, Oxford, Ohio 45056, USA; email: widome@muohio.edu
Running title: Ultra-depleted oceanic lithosphere
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ABSTRACT
Many peridotite xenoliths sampled at ocean islands appear to have strongly refractory major
element and modal compositions. In order to better constrain the chemistry, abundance and
origin of these ultra-refractory rocks we compiled a large number of data for xenoliths from 9
different groups of ocean islands. The xenoliths were filtered petrographically for signs of
melt infiltration and modal metasomatism and the samples affected by these processes were
excluded. Most ocean islands are dominated by ultra-refractory harzburgites. Exceptions are
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Hawaii and Tahiti peridotites which are more fertile and contain primary clinopyroxene, and
Cape Verde, which contains both ultra-refractory and more fertile xenoliths. Ultra-refractory
harzburgites are characterized by the absence of primary clinopyroxene, low whole rock
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Al2O3, CaO, FeO/MgO and HREE concentrations, low Al2O3 in orthopyroxene (generally <3
wt%), high Cr# in spinel (0.3-0.8) and high forsterite contents in olivine (averages >91.5).
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They are therefore on average significantly more refractory than peridotites dredged and
drilled from mid-ocean ridges and fracture zones. Moreover, their compositions resemble
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those of oceanic fore-arc peridotites. The formation of ultra-refractory ocean island
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harzburgites requires potential temperatures above those normally observed at modern midocean ridges, and/or fluid fluxed conditions. Some ultra-refractory ocean island harzburgites
give high Os model ages (up to 3300 Ma), showing that their formation significantly pre-dates
the oceanic crust in the area. A genetic relationship with the host plume is considered unlikely
based on textural observations, equilibration temperatures and pressures, inferred physical
properties, and long-term depleted Os and Sr isotope ratios of some of the harzburgites.
Although we do not exclude the possibility that some ultra-refractory ocean island
harzburgites have formed at mid-ocean ridges, we favor a model in which they formed in a
process spatially and temporally unrelated to the formation of the oceanic plate and the host
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plume. Due to their whole rock compositions, ultra-refractory harzburgites have a very high
solidus, low densities and very high viscosities, and will tend to accumulate at the top of the
convecting mantle. They may be preserved as fragments in the convecting mantle over long
periods of time and be preferentially incorporated into newly formed lithosphere.
KEY WORDS: mantle xenoliths; ocean islands; recycled oceanic mantle lithosphere; ultradepleted mantle.
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INTRODUCTION
Detailed information about chemical composition is essential in understanding the formation
and evolution of different parts of the Earth’s mantle. Small variations in the major element
composition of the lithospheric mantle may strongly influence its physical properties and its
behavior in response to stress (e.g., Karato, 1986; Hirth & Kohlstedt, 1996; Afonso et al.,
2007; Simon & Podladchikov, 2008). It is therefore necessary to obtain reliable chemical
compositions of different mantle domains in order to develop realistic geodynamic models
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and correctly interpret geophysical data.
Mantle xenoliths sampled in ocean islands (OI) cover a wide range in compositions.
Some islands are dominated by spinel harzburgites that are more refractory than most
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peridotites sampled along mid-ocean ridges and fractures zones (MORP: mid-ocean ridge
peridotites). The xenolith series in a second group of islands are relatively fertile, whereas in a
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third group of ocean islands the xenoliths range from strongly refractory to fertile. In the third
group the fertile compositions are to a large extent the results of metasomatism associated
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with ocean island formation and/or melt-rock interaction during ascent, and many of the more
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fertile peridotites have strongly refractory pre-OI protoliths (e.g., Ryabchikov et al., 1995;
Wulff-Pedersen et al., 1996; Grégoire et al., 1997; Grégoire et al., 2000a; Grégoire et al.,
2000b; Moine et al., 2000; Neumann et al., 2002; Delpech, 2004; Neumann et al., 2004;
Bonadiman et al., 2005; Shaw et al., 2006).
The presence and abundance of ultra-refractory mantle not only in continental keels,
but also in oceanic domains, is an important observation. Most of this evidence for ‘depleted
compositions’ is based on isotope and trace element data on single minerals, samples and/or
locations. However, in order to use compositional variations in geodynamic modeling whole
rock major element compositions and modal mineralogy averaged over some larger area are
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actually the most important variables. This manuscript represents a first step in providing such
constraints for more petrologically consistent quantitative geodynamic modeling. Moreover,
the mechanism for the generation of these samples is far from clear and this review will serve
as a solid basis for further work.
Studies of mantle xenoliths in ocean islands have so far mainly been focused on
mantle metasomatism and melt-rock interaction during ascent. These are recent processes in
the evolution of the peridotites. We are interested in the pre-ocean island composition of
ocean island peridotites and particularly in the most refractory xenoliths. The aim of this
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paper is (i) to establish the distribution of strongly refractory peridotites in ocean islands, (ii)
to constrain their origin and mode of formation, and (iii) discuss the implications of strongly
refractory mantle material for geodynamics with respect to physical and chemical properties.
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We have collected available (published and unpublished) major element whole rock
and mineral compositions on 294 harzburgite and lherzolite xenoliths and their minerals from
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the Canary Islands (Neumann, 1991; Neumann et al., 1995; Wulff-Pedersen et al., 1996;
Neumann et al., 2002; Neumann et al., 2004; Abu El-Rus et al., 2006), Madeira (Munha et
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al., 1990), the Azores (Johansen, 1996), Cape Verde (Ryabchikov et al., 1995; Bonadiman et
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al., 2005; Shaw et al., 2006), Kerguelen (Grégoire et al., 1997; Grégoire et al., 2000a;
Grégoire et al., 2000b; Delpech, 2004), Grande Comore (Coltorti et al., 1999), Samoa (Hauri
et al., 1993; Hauri & Hart, 1994), Hawaii (e.g., Jackson & Wright, 1970; Sen, 1987; Goto &
Yokoyama, 1988; Sen & Leeman, 1991; Bizimis et al., 2003) and Tahiti (Tracy, 1980; Qi et
al., 1994). We refer to samples from the same island as a xenolith suite. The data are
presented in the Electronic data base (http://www.petrology.oxfordjournals.org). The data
base also gives major element on 450 MORP used for comparison.
Our study is based mainly on the most abundant major elements (SiO2, MgO, FeO,
Al2O3 and CaO) because these are less easily reset by fertilization processes than
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incompatible trace elements. In addition to spinel harzburgites and lherzolites, mantle
xenoliths from OI include dunites and wehrlites and pyroxenites. Dunites, wehrlites,
pyroxenites, etc. are not discussed here as these rock-types are closely associated with
fertilization processes (e.g., Frey, 1980; Sen, 1988; Neumann, 1991; Sen & Leeman, 1991;
Kelemen et al., 1992; Grégoire et al., 1997; Kelemen et al., 1998; Grégoire et al., 2000b;
Neumann et al., 2004). However, data on some of those samples are included in the data base.
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COMPOSITIONAL VARIATIONS
Classification of samples
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Since the focus of this paper is on pre-ocean-island (pre-OI) compositions we apply
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petrographic criteria to recognize OI peridotites that have been least affected by melt-rock
interaction processes. The effects of such processes range from very minor, reflected only in
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mild enrichment in strongly incompatible trace elements and resetting of radiogenic isotopes,
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to extensive petrographic changes such as formation of hydrous minerals, reactions involving
the growth of ±olivine±clinopyroxene±spinel (±ol±cpx±sp) at the expense of orthopyroxene
(opx), formation of poikilitic orthopyroxene and clinopyroxene, glass-rich reaction zones on
primary minerals (e.g. orthopyroxene and clinopyroxene) and sieve textured spinel and
clinopyroxene (Fig. 1), in addition to significant enrichment in incompatible trace elements.
Peridotites showing these textures are classified as OI2, indicating the presence of a second
generation of minerals related to melt-rock reaction (Table 1).
Those xenoliths that show no petrographic evidence of metasomatism (as listed above
and summarized in Table 1) are classified as OI1 (Fig. 2). These xenoliths typically show
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porphyroclastic to protogranular textures with deformed porphyroclasts of olivine and
exsolved orthopyroxene; in some xenolith series the porphyroclast assemblage comprises
exsolved clinopyroxene (e.g. from Hawaii and Tahiti) and/or spinel. The matrix consists of
less deformed or undeformed neoblasts of olivine + orthopyroxene ± spinel ± clinopyroxene.
In some cases the published petrographic descriptions do not give enough information to
classify a given sample; such samples are assigned to the OI2 group. Since this paper is aimed
at investigating the composition of the upper mantle before the onset of ocean island
magmatism the processes that have led to the petrographic changes recorded in the OI2
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samples are of no interest here, and are not discussed. We therefore make no effort to separate
between in-situ metasomatism in the mantle and interaction with the host magma during
ascent. Data points in the different figures are restricted to OI1 xenoliths. However, as it is of
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interest to the discussion to know the extent to which melt-rock interaction has changed preOI peridotite compositions, we show the ranges covered by OI2 xenoliths in some diagrams.
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In addition to the petrographic distinction into OI1 and OI2, the xenoliths from
different islands fall in two categories with respect to major element compositions (Table 1).
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One category comprises ultra-refractory harzburgites (OI-u) with low contents of
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clinopyroxene (<3 volume %) and other fusible components; the other category consists of
more fertile xenolith series (OI-f), in which lherzolites with primary clinopyroxene are
common. Essential petrographic and chemical differences between the two groups are
outlined in Table 1. The focus of this paper is on OI-u.
We emphasize that we distinguish between refractory/fertile, which refers to major
element compositions and ultimately solidus temperatures, and the terms depleted/enriched,
which refer to trace element and isotope characteristics. These chemical characteristics may
be decoupled. A refractory harzburgite may, for example, be enriched in certain incompatible
elements, but still retains its high solidus and related physical properties. Most samples in the
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OI1-u group show some evidence of cryptic metasomatism (i.e., U-shaped REE-patterns). We
have, however, found a clear positive correlation between the extent of melt-rock interaction
processes as indicated by petrographic observations, and enrichment in strongly incompatible
trace elements. On the basis of numerical modeling, Herzberg et al. (2007) concluded that a
small fraction of trapped melt in a residual rock does not significantly affect the most
abundant major elements. For the rocks discussed here this conclusion is supported by an
excellent correlation between whole rock major element chemistry and parameters such as cr#
(cation proportion Cr/(Cr+Al)) and mg# (cation proportion Mg/(Mg+Fe2+)) in spinel, and the
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concentration of Al2O3 and Fo-content in cores of orthopyroxene and olivine porphyroclasts,
respectively.
In addition to chemical compositions we compare equilibration temperatures (Table 2)
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for the different peridotite suites. OI1-u harzburgites lack primary clinopyroxenes and
secondary clinopyroxenes in these rocks are commonly too small to analyze. OI2 xenoliths
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show two or more stages of mineral growth, and it is difficult to ascertain which mineral
compositions represent equilibrium. We therefore chose to use the single-mineral CaO-in-opx
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geothermometer of Sachtleben & Seck (1981), assuming that the presence of clinopyroxene
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exsolved from orthopyroxene satisfies the four-phase requirements, also for harzburgites
without primary clinopyroxene. This is in agreement with the conclusions of Neumann (1991)
and Neumann et al. (1995) based on the application of different geothermometers for OI1-u
harzburgites from the Canary Islands. The most important characteristics of the xenolith
suites from each ocean island are summarized below.
Canary Islands
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Xenoliths from the Canary Islands belong to the OI-u group (Table 1). They are dominated by
spinel harzburgites with protogranular, porphyroclastic or poikilitic textures (Neumann, 1991;
Siena et al., 1991; Neumann et al., 1995; Wulff-Pedersen et al., 1996; Neumann et al., 2002;
Neumann et al., 2004; Abu El-Rus et al., 2006; Neumann, unpublished data, 2007). The OI1u group is restricted to porphyroclastic to protogranular harzburgites with porphyroclasts of
olivine (ol1) and orthopyroxene (opx1); in some samples large equidimensional spinels may
also represent porphyroclasts (sp1). The groundmass consists of equant to irregular grains of
olivine, orthopyroxene and spinel, and minor amounts (generally <3 volume percent) of
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clinopyroxene. Clinopyroxene is present as exsolution lamellae in orthopyroxene and as small
neoblasts surrounding orthopyroxene porphyroclasts (opx1) and is interpreted as a younger
generation (cpx2) than the porphyroclasts (Fig. 2a). Rare, local occurrences of small,
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interstitial, vermicular cpx-sp intergrowths may represent melt infiltration.
OI2-u xenoliths (mainly from La Palma and Tenerife) comprise both spinel
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harzburgites and lherzolites. Clinopyroxene in OI2-u is found as small interstitial grains, as
reaction rims on orthopyroxene (Fig. 1a, e), in cpx+sp±opx symplectites along grain
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boundaries and/or as part of reaction textures (formation of clinopyroxene ± olivine at the
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expense of orthopyroxene; Fig. 1a, e), or as poikilitic grains enclosing round grains of olivine,
embayed grains of orthopyroxene (Fig. 1d) and small spinel grains. Many xenoliths from
Fuerteventura exhibit fibrous orthopyroxene. Some OI2 samples, particularly from La Palma
and Tenerife, carry trace amounts of phlogopite.
No primary clinopyroxene porphyroblasts (cpx1) have been observed in the Canary
Islands xenoliths; the clinopyroxene present (cpx2) is believed to have formed partly by
exsolution from orthopyroxene porphyroclasts (opx1) and subsequent recrystallization (Fig.
2a), partly from interaction with infiltrating melts (Fig. 1a, d, e; Neumann et al., 1995;
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Neumann et al., 2004). The samples with the highest modal proportions of clinopyroxene also
show the strongest textural evidence of metasomatism.
Whole rock major element relationships for Canary Islands OI1-u are plotted in Figure
3. We have also plotted orthopyroxene-olivine (opx-ol) tie-lines, as defined by OI1-u
xenoliths from Lanzarote. Major elements other than Na2O plot close to the opx-ol tie-line for
most OI1-u xenoliths, consistent with the absence of primary clinopyroxene and confirming
their strongly refractory composition. A few samples scatter at a moderate distance away from
the opx-ol tie-line, for example in the SiO2-CaO diagram. We interpret this scatter as the
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result of minor metasomatism and/or melt–rock interaction. The wide range in Na2O is
probably due to enrichment processes or contamination (Neumann, 1991; Siena et al., 1991;
Neumann et al., 1995; Neumann et al., 2004). The strongly melt-depleted nature of the OI1-u
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harzburgites is also reflected in low Al2O3 in orthopyroxene (0.3-3.4 wt%), Cr-rich spinels
(cr# = 0.29-0.95, with the majority of the samples in the range 0.53-0.71), and relatively high
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Fo-content (100Mg/(Mg+Fe)) in olivine (90.7-92.5; Fig. 4).
Many OI2-u xenoliths also plot close to the opx-ol tie-line (Fig. 3), although the
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majority have lower SiO2 and MgO, and higher CaO-FeO* contents than the OI1-u xenoliths
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and show wide ranges in TiO2 and Na2O. Compared to the OI1-u rocks, the OI2-u rocks show
a somewhat wider range in (Al2O3)opx, a similar range in cr#sp, but tend towards lower mg#
(Fig. 4). Estimated equilibration temperatures for OI1-u xenoliths are 873-1209 ºC with a
mean of 1024 ºC (Table 2). The OI2-u xenoliths cover a similar range (862-1215 ºC) but have
a significantly higher average, 1068 ºC.
Cores of OI1-u orthopyroxene porphyroclasts (opx1) are strongly depleted in REE
(YbN= 0.09-0.41, LaN= 0.002-1.2) and in light relative to heavy REE (LREE and HREE,
respectively; Neumann et al., 1995; Neumann et al., 2004; Fig. 5), which is typical for mantle
rocks that represent residues after extensive partial melting (e.g. Johnson et al., 1990;
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Neumann et al., 2004). A tendency for enrichment in LREE relative to the middle REE
(MREE) is most likely due to contamination (Neumann et al., 2002).
In contrast to the LREE-depleted opx1 in the OI1-u xenoliths, cpx2 in the same rocks
show a wide range in REE patterns and concentrations, ranging from LREE-depleted in the
clinopyroxenes with the lowest HREE, through U-shaped and flat patterns, to LREE-enriched
patterns in clinopyroxenes with the highest HREE (YbN= 0.54-1.1, LaN= 0.22-128; Fig. 6).
Madeira
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Mantle xenoliths from Madeira comprise refractory deformed harzburgites and lherzolites
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(OI-u; Fo90-91) and less deformed dunites, websterites, wehrlites and clinopyroxenites (Fo77-88;
Munha et al., 1990). Harzburgites and lherzolites are granular (most common) to
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porphyroclastic with deformed olivine and exsolved orthopyroxene porphyroclasts in a finegrained matrix of ol+opx+sp±cpx; minor amounts of phlogopite may be present interstitially
or as inclusions in clinopyroxene (in lherzolite).
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OI1-u xenoliths from Madeira (Munha et al., 1990) have lower (Al2O3)opx (1.5-3.6
wt%) than orthopyroxene in MORP, but fall within the depleted range of MORP with respect
to spinel compositions (cr#: 0.26-0.55) and Fool (90.4-91.2; Fig. 4). Temperature estimates for
the OI1-u xenoliths are 891-1017 ºC with a mean of 936 ºC (Table 2); OI2-u xenolith give the
wider range of 892-1182 ºC, with a mean of 1037 ºC.
The Azores
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Manuscript submitted to Journal of Petrology
12
Mantle xenoliths from the Azores comprise Mg-Cr-rich, porphyroclastic spinel harzburgites
(<2 volume percent clinopyroxene) and lherzolites ( 15 volume percent clinopyroxene) as
well as granular Fe-Ti-rich spinel wehrlites, clinopyroxenites and dunites (Johansen, 1996;
Neumann, unpublished data, 2007). Harzburgites predominate over lherzolites. The oldest
generation of minerals in the harzburgites and lherzolites are irregular, deformed
porphyroclasts of olivine and orthopyroxene with clinopyroxene exsolution lamellae, and
occasional large spinels. The matrix appears to consist of two generations of neoblasts. The
older neoblast generation consists of equigranular domains of polygonal neoblasts of olivine
rP
Fo
(occasionally with undulatory extinction) and rare unexsolved orthopyroxene; the younger
generation is made up by fine-grained, glass-rich domains of irregular neoblasts of olivine,
orthopyroxene, clinopyroxene, and spinel, without undulatory extinction (Johansen, 1996).
ee
Clinopyroxene occurs partly as small interstitial grains, partly as poikilitic grains. All Azores
xenoliths are classified as OI2.
rR
The Azores xenoliths are dominated by fertile whole rock major element
compositions, but are depleted with respect to (Al2O3)opx (2.1-2.3 wt%) and cr#sp (0.51-0.72;
ev
Fig. 4). Fool falls within the range 90.0-91.2 in all samples, with the exception of one with
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Fool of 86.5 (Fig. 4). Equilibration temperatures are high (1003-1232 ºC, mean: 1135 ºC;
Table 2). The textures of these rocks suggest that the contrast between whole rock and
mineral chemistry is the result of recent melt infiltration into a refractory peridotite protolith
which in general has not had time to affect the cores of porphyroclasts and large spinels.
Cape Verde
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13
Xenoliths from Sal in the Cape Verde archipelago have been described by Ryabchikov et al.
(1995), Bonadiman et al. (2005) and Shaw et al. (2006). The xenoliths have protogranular to
porphyroclastic and mylonitic textures and comprise refractory harzburgites without primary
clinopyroxene ( 3 % cpx) as well as mildly depleted spinel lherzolites with primary
clinopyroxene (8-18 volume % cpx). Many samples show extensive evidence of reactions,
e.g. glass-bearing sieve-textured rims on clinopyroxenes and spinels, and rims of
ol+sp+cpx+glass on orthopyroxene. Ryabchikov et al. (1995) also report phlogopite in some
samples. Glass is common, both interstitially and as small inclusions in olivine and
rP
Fo
orthopyroxene. Bonadiman et al. (2005) proposed that the strongly refractory spinel
harzburgites and the mildly depleted spinel lherzolites represent two unrelated suites. Both
suites include OI1 and OI2 samples (Figs. 2b, d and Fig. 1f, respectively). The presence of
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two distinct xenolith suites is supported by the relationship between modal proportions of
clinopyroxene and olivine (Fig. 7). Within overlapping ranges of olivine the lherzolites (OI-f)
rR
define a trend of strongly decreasing clinopyroxene with increasing olivine, whereas the
harzburgites (OI-u) show roughly constant, low clinopyroxene contents. Equilibration
temperatures are given in Table 2.
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The OI-u harzburgites are strongly refractory with whole rock major element contents
falling close to the opx-ol tie-line (e.g. 44.9-47.9 wt% MgO, 0.56-1.1 wt% CaO, <0.8 wt%
Al2O3; Fig. 3). Somewhat higher TiO2 in a few samples is probably due to minor
contamination.
Mineral major element data are only available for one OI1-u harzburgite, which has
Cr-rich spinel (cr#=0.55), Fo-rich (91.3) olivine and Al2O3-poor orthopyroxene (2.6 wt %;
Fig. 4). Cores of orthopyroxene porphyroclasts in OI1-u have low HREE contents and show
strongly LREE-depleted to U-shaped REE-patterns (YbN= 0.04-0.29, LaN= 0.02-0.35; Fig. 5).
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14
Clinopyroxene in OI1-u shows LREE-depleted to LREE-enriched patterns (YbN= 0.7-4.5,
LaN= 0.3-12.2; Fig. 6).
The Cape Verde OI1-f lherzolites are relatively fertile with 39.7-45.5 wt% MgO, 1.24.6 wt% CaO and 0.5-2.0 wt% Al2O3 (Fig. 8); cr#sp= ~0.3, Fool= 87-89 and (Al2O3)opx= 4.16.0 wt% (Fig. 4). Orthopyroxene porphyroclasts in OI1-f are depleted in HREE and have
moderately LREE-depleted patterns (YbN= 0.23-0.36, LaN= 0.03-0.5; Fig. 5), whereas OI1-f
clinopyroxene have upwards convex REE-patterns (YbN= 0.9-1.5, LaN= 1.7-2.0; Fig. 6).
Kerguelen
ee
rP
Fo
Mantle xenoliths from Kerguelen belong to the OI-u group and the suite is dominated by
protogranular (OI1-u, Fig. 2a) and poikilitic harzburgites (OI2-u, Fig. 1c). Some OI2 samples
rR
are cut by thin mafic veinlets (of e.g. pyroxenite, glimmerite, hornblendite), which may show
reactions against spinel or orthopyroxene (e.g. Mattielli et al., 1996; Grégoire et al., 1997;
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Grégoire et al., 2000a; Grégoire et al., 2000b; Moine et al., 2000; Delpech, 2004). In OI1-u,
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clinopyroxene is generally present as small dispersed, interstitial neoblasts (cpx2), sometimes
associated with opx and vermicular spinels. In OI2, clinopyroxene also occurs in
symplectites, in or along veinlets, and in reaction zones around orthopyroxene and spinel. A
few poikilitic OI2 samples contain phlogopite and amphibole. Trace amounts of apatite,
feldspar, rutile, chromite, ilmenite and armalcolite have been observed in glass-bearing
domains or veinlets.
All the OI1-u and most of the OI2-u xenoliths fall close to the opx-ol tie-line in most
major element diagrams (Fig. 3). OI1-u have orthopyroxenes with low Al2O3 contents (1.82.4 wt%), high cr#sp (0.42-0.61) and high, uniform Fool of 91.0-91.8 (Fig. 4). Minerals in OI2-
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15
u tend towards more fertile compositions. OI1-u xenoliths give lower equilibration
temperatures (933-1104 ºC, average 1014 ºC; Table 2) than OI2-u (960-1193 ºC, average
1062 ºC).
Pyroxenes REE patterns indicate minor interaction with REE-enriched melts/fluids
(e.g. Mattielli et al., 1996; Grégoire et al., 1997; Mattielli et al., 1999; Grégoire et al., 2000a;
Grégoire et al., 2000b; Delpech, 2004). Most OI1-u orthopyroxenes show low concentrations
in HREE and MREE/HREE <<1, but some have irregular REE patterns (e.g. YbN: 0.03-0.23,
LaN= 0.011-0.93; Fig. 5). Clinopyroxenes in OI1-u xenoliths show a range in REE patterns
rP
Fo
from LREE-depleted, through U-shaped, to LREE-enriched, with YbN= 0.13-0.97 and LaN=
0.013-21 (Fig. 6).
Grande Comore
rR
ee
Mantle xenoliths from Grande Comore comprise spinel lherzolites, dunites and wehrlites; no
ev
harzburgites have been reported (Coltorti et al., 1999). The lherzolites have protogranular to
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porphyroclastic textures, and consist of large, deformed porphyroclasts of olivine and
orthopyroxene and small, strain-free grains of clinopyroxene and spinel which occur
interstitially or as inclusions in orthopyroxene (Ludden, 1977; Coltorti et al., 1999). The
primary textures have been overprinted by pyrometamorphic textures (e.g. melt infiltration),
resulting in reactions leading to formation of significant amounts of clinopyroxene, olivine
and spinel at the expense of orthopyroxene. Melt-rock interaction processes in the Grande
Comore lherzolites are also recorded in grain-to-grain compositional differences within single
samples, and zoning (Hauri et al., 1993; Hauri & Hart, 1994). All Grande Comore xenoliths
are classified as OI2.
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16
Despite the reaction textures, the Grande Comore OI2 xenoliths have low (Al2O3)opx
(0.8-3.4 wt%), high cr#sp (0.4-0.7; Fig. 4) and moderate Fool (90.0-91.3), classifying them as
OI2-u. Estimated equilibration temperatures fall in the range 965-1139 ºC, with a high mean
temperature of 1109 ºC (Table 2).
Samoa
rP
Fo
Mantle xenoliths from Savai’i (Samoa) are dominated by strongly refractory spinel
harzburgites with protogranular to porphyroclastic (most common) textures (Hauri et al.,
1993; Hauri & Hart, 1994) and belong to the OI-u group. Symplectites are common in the
ee
porphyroclastic rocks. Clinopyroxene occurs mostly along the edges of orthopyroxene+spinel
symplectites, but is also found as small interstitial grains rimming orthopyroxene and olivine.
rR
In some rocks, melt-rock interaction led to interstitial glass+clinopyroxene±apatite-bearing
patches in various stages of reaction with adjacent porphyroclasts (Hauri et al., 1993). Three
ev
harzburgites with porphyroclastic textures ( 0.5 percent cpx) have been classified as OI1-u.
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Whole rock major element data are only available for OI2-u harzburgites; these fall
close to the opx-ol tie-line (not shown). Mineral compositions are typical of OI1-u (1.3-2.6
wt% (Al2O3)opx; cr#sp: 0.49-0.70; Fool: 90.5-91.4; Fig. 4). OI2-u xenoliths tend towards more
fertile compositions, particularly less Mg-rich spinel and olivine. OI1- and OI2-u rocks give
roughly similar equilibration temperatures (983-1142 ºC with average 1063 ºC and 1045-1097
ºC with average 1072 ºC, respectively; Table 2).
Samoan OI1-u clinopyroxenes show a wide range in REE contents (YbN= 0.045-8.6,
LaN=0.011-2.5; Hauri et al., 1993; Hauri & Hart, 1994; Fig. 6) both between and within
samples. Hauri & Hart (1994) found a strong correlation between REEcpx and textures.
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17
Hawaii
The Hawaiian mantle xenoliths belong to the mildly depleted OI-f group. They range in rock
types from spinel- and spinel-plagioclase-bearing harzburgites and lherzolites, to dunites,
wehrlites, websterites and pyroxenites, some of which contain garnet (e.g., Jackson & Wright,
1970; Sen, 1987; Goto & Yokoyama, 1988; Sen & Leeman, 1991; Bizimis et al., 2003).
rP
Fo
Textures range from porphyroclastic to coarse-grained and fine-grained granular.
Clinopyroxene-rich spinel lherzolites are common (up to 24 vol.% cpx; e.g. Jackson &
Wright, 1970; Sen, 1987; Goto & Yokoyama, 1988; Sen, 1988; Sen & Leeman, 1991; Bizimis
ee
et al., 2003). Primary clinopyroxene porphyroclasts have spinel or orthopyroxene exsolution
lamellae (e.g. Sen, 1988).
rR
The whole rock major element compositions of all the Hawaiian harzburgites and
lherzolites plot far away from the opx-ol tie-lines in Figure 8; Hawaiian OI1 xenoliths include
ev
some of the lowest MgO (37-43 wt%), and highest CaO (1.1-3.9 wt%), TiO2 ( 0.43 wt%),
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Al2O3 ( 4 wt%) and Na2O ( 0.7 wt%) contents reported from ocean islands. OI1-f mineral
compositions are also relatively fertile (1.7-4.9 wt% (Al2O3)opx, cr#sp= 0.09-0.55, Fool: 88.592.0; Fig. 4). There is a strong tendency for lower Fool in the OI2-f compared to the OI1-f.
Equilibration temperatures are similar for OI1-f (914-1205 ºC with a mean of 1004 ºC) and
OI2-f (881-1161 ºC, mean 1005ºC; Table 2).
A large amount of clinopyroxene REE data from Hawaiian peridotites is published,
but no orthopyroxene REE data. Clinopyroxene in one OI1-f plagioclase-spinel harzburgite
and four OI1-f spinel lherzolites from Oahu exhibit smooth REE patterns with strong LREE
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18
depletion (YbN= 3.4-4.7, LaN= 0.02-1.0; Yang et al., 1998; Sen et al., 2005; Fig. 6). Minor
metasomatism is indicated by increasing La and Ce from cores to rims in a couple of samples.
In contrast to the OI-f xenoliths from the Hawaiian islands, mantle xenoliths from the
Loihi Seamount comprise dunites and spinel harzburgites, but no lherzolites (Clague, 1988).
The harzburgites have cataclastic textures. OI-u compositional characteristics are reflected in
low (Al2O3)opx (1.5-2.0 wt%), high cr#sp (0.53-0.75; Fig. 4) and high Fool (90.6-92.3; Fig. 4).
Estimated temperatures for the OI1-u are even higher than in Hawaiian OI1-f, 1139-1182 ºC,
with a mean of 1163 ºC (Table 2).
Tahiti
ee
rP
Fo
Tahiti mantle xenoliths are porphyroclastic to equigranular, mildly refractory OI-f spinel
rR
lherzolites (most abundant) and harzburgites (Tracy, 1980; Qi et al., 1994). Clinopyroxene
occurs partly as large porphyroclasts, partly in vermicular intergrowths with spinel and
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olivine. Some clinopyroxenes have spongy glass-bearing rims, interpreted as being a result of
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metasomatism (Qi et al., 1994). No other evidence of metasomatism is given by Tracy (1980)
or Qi et al. (1994). Samples with spongy glassy rims have been classified as OI2-f, the rest of
the Tahiti samples as OI1-f.
As the Hawaii xenoliths, most OI1-f from Tahiti have lower MgO and higher SiO2,
CaO, FeO* and Al2O3 than OI1-u harzburgites, but they have moderate TiO2 and Na2O
contents ( 0.12 and
0.05 wt %, respectively; Fig. 3). They also have relatively fertile
mineral compositions with (Al2O3)opx: 2.3-5.6 wt%, cr#sp: 0.12-0.41 and Fool: 84.6-90.7 (Fig.
4). OI1-f equilibration temperatures are 996-1108 ºC with a mean of 1058 ºC (1093 ºC for the
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single OI2-f sample; Table 2). Clinopyroxenes in OI1-f have uniform, mildly LREE-enriched
REE-patterns (YbN= 2.2-4.0, LaN= 1.6-3.4; Fig. 6).
DISCUSSION
Different ocean islands show significant contrasts in their OI1 mantle xenolith populations.
Xenoliths from the Canary Islands, Madeira, Kerguelen, Samoa and Loihi seamount are
rP
Fo
strongly refractory (OI1-u and OI2-u), whereas those from Hawaii and Tahiti are mildly
refractory (OI1-f and OI2-f). The textural and compositional characteristics of each group are
summarized in Table 1. Cape Verde xenoliths appear to comprise both categories. Although
ee
no data on OI1 xenoliths are available for the Azores and Grande Comore, the low (Al2O3)opx,
high cr#sp and high Fool in many OI2 samples suggest that ultra-refractory OI1-u domains
rR
may also be (have been) present in the upper mantle beneath these islands. Below we
summarize the characteristics of OI-u and OI-f xenoliths, compare the OI-u to MORP, and
ev
speculate on the origin and evolution of OI-u. Furthermore, we provide some estimates of the
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proportion of the uppermost mantle that might consist of ultra-refractory peridotites and
discuss some implications for geodynamics.
Ultra-refractory OI peridotites (OI-u)
As shown in Figure 3 the whole rock major element compositions of OI1-u and OI2-u
xenoliths (the Canary Islands, Madeira, Cape Verde, Kerguelen, Samoa and Loihi) essentially
reflect different modal proportions of olivine and orthopyroxene; clinopyroxene generally
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20
makes up <2 vol.% and occurs as secondary neoblasts (Figs. 1, 2). OI1-u have high cr#sp, high
Fool, low (Al2O3)opx and low HREE concentrations in pyroxenes (Table 1; Figs. 4-6). Their
ultra-refractory compositions strongly suggest that they are residues after partial melting
beyond the stability of primary clinopyroxene. This is in agreement with the degrees of partial
melting proposed for example for Samoa (33-42% melting; Hauri & Hart, 1994) and the
Canary Islands (25-30% melting; Neumann et al., 2004). Furthermore, the OI1-u
compositions are similar to mantle residues formed by two stage melting experiments (Green
et al., 2004; Fig. 3) in which a residual, dry, chromite-bearing harzburgite with compositions
rP
Fo
close to MORP was melted by an additional 10% under H2O-CO2–fluxed conditions at 1.5
GPa and 1320-1350 ºC.
The secondary origin of clinopyroxene in OI1-u and OI2-u xenoliths from the Canary
ee
Islands is reflected in REE. Two OI1-u samples with mildly LREE-depleted cpx-patterns give
opx/cpx REE ratios close to the range of published equilibrium opx/cpx REE ratios in spinel
rR
peridotites (Fig. 9), supporting petrographic evidence (Fig. 2) that these clinopyroxenes
(cpx2) formed through exsolution from orthopyroxene porphyroclasts (opx1), followed by
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recrystallization. Many OI2-u, in contrast, have opx/cpx REE ratios that fall significantly
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below published REE partition coefficients (Fig. 9), implying disequilibrium between the
pyroxenes. We interpret these low ratios as the results of a much stronger influence of LREEenriched fluids/melts on clinopyroxene than on orthopyroxene; some of the most LREEenriched clinopyroxenes may have formed directly from LREE-rich infiltrating melts.
The less refractory nature of the OI2-u lherzolites does not imply that they represent
pristine mantle or mantle domains that have undergone lower degrees of partial melting than
the OI1-u harzburgites. Instead, we interpret the OI2-u as being formed by interaction
between melts and OI1-u protoliths. Evidence that interaction between melts and refractory
mantle is an important mode of formation of lherzolites comes from abyssal peridotites,
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Page 21 of 87
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peridotite massifs, ophiolites and mantle xenoliths (e.g. Hellebrand et al., 2002; Grégoire et
al., 2003; Müntener & Piccardo, 2003; Simon et al., 2003; Le Roux et al., 2007; Godard et
al., 2008), and from experimental work (Yaxley & Green, 1998; Yaxley, 2000; Bulatov et al.,
2002).
Mildly refractory OI peridotites (OI-f)
rP
Fo
The OI-f xenoliths (Hawaii, Tahiti, the lherzolite suite from Cape Verde) are significantly
more fertile than the OI1-u series (lower whole rock MgO and higher CaO-FeO*-TiO2-Al2O3
contents, higher (Al2O3)opx, lower cr#sp and lower Fool, common presence of primary
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clinopyroxene; Figs. 4, 6, 8; Table 1). A primary origin of clinopyroxenes in many xenoliths
from Hawaii is supported by their strongly LREE-depleted patterns (Yang et al., 1998; Fig.
rR
6). The OI1-f peridotites appear to have undergone a lower degree of partial melting than
OI1-u. Yang et al. (1998) estimated about 2-8% melting of a depleted mantle source for a
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group of spinel peridotites from Oahu. The OI1-f xenolith suites from Hawaii, Tahiti and
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Cape Verde indicate the presence of mildly refractory domains beneath these islands; the
upper mantle beneath Hawaii-Loihi and Cape Verde includes both strongly refractory and
more fertile domains.
Comparison between OI1-u and MOR peridotites
Peridotites that are dredged and drilled at ’normal’ segments along mid-ocean ridges and at
fracture zones (MORP) comprise a large part of the available oceanic peridotite database and
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22
are often regarded as representative of all oceanic lithospheric mantle. We therefore explore
here how MORP compare with OI1-u peridotites. We restricted the MORP samples
(recalculated to 100% dry) presented in Figure 10 to harzburgites and lherzolites (using modal
compositions as criteria). However, as modal data are not available for all samples, it is likely
that some of the rocks with the lowest SiO2 contents (<42 wt%) are in fact dunites, even
though they are classified as harzburgites in the literature. The samples which Seyler &
Bonatti (1997) and Tartarotti et al. (2002) described as strongly infiltrated by melts are
excluded from Figure 10.
rP
Fo
MORP show wide ranges in whole rock and mineral major element compositions.
Most samples have 35-49 wt% MgO, 40-49 wt% SiO2, <0.1-4.5 wt% CaO and 0.3-4.5 wt%
Al2O3 (Fig. 10). (Al2O3)opx fall in the range 1.7-6.1 wt%, cr#sp is 0.12-0.57, and in most
ee
MORP Fool is 90-91 (Fig. 4). It is generally accepted that MORP do not represent pristine
residues after extraction of mid-ocean ridge basalts, but have been chemically modified by
rR
open system melt migration and crystallization, melt/rock interaction by reactive porous flow,
metamorphism, serpentinization and marine weathering (e.g., Elthon, 1992; Hékinian et al.,
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1993; Cannat et al., 1997; Casey, 1997; Niu, 1997; Niu & Hékinian, 1997a; Asimow, 1999;
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Baker & Beckett, 1999; Tartarotti et al., 2002; Niu, 2004). The combination of these
processes explains much of the strong compositional heterogeneity observed in MORP. The
low CaO/Al2O3 ratios exhibited by many MORP (about 10 to <0.01; Fig. 11) are probably
due to leaching of Ca during serpentinization (e.g. Bach et al., 2004; Palandri & Reed, 2004).
OI1-u xenoliths have CaO/Al2O3 ratios close to 1.0, which is typical of harzburgites and
lherzolites that have not been subjected to serpentinization (Boyd, 1996).
Because of post-melting processes, the origin of clinopyroxene in MORP has been a
subject of debate. Elthon (1992) claimed that MORP experienced partial melting to the
exhaustion of clinopyroxene. He proposed that
2% clinopyroxene may be explained by
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23
granular exsolution from orthopyroxene, whereas proportions >2% most likely formed from
infiltrating melt. Other authors (e.g., Dick & Bullen, 1984; Edwards et al., 1996; Seyler &
Bonatti, 1997; Tartarotti et al., 2002; Seyler et al., 2003; Niu, 2004; Brunelli et al., 2006)
have presented extensive petrographic and geochemical evidence that a significant proportion
of the clinopyroxene found in MORP is primary implying melting within the stability field of
ol+opx+cpx. However, these authors acknowledge that some clinopyroxene has formed
through melt infiltration.
On average OI1-u harzburgites have higher MgO, lower (Al2O3)opx and higher cr#sp
rP
Fo
than MORP and are thus significantly more refractory (Fig, 4). Furthermore, OI1-u has no
residual clinopyroxene, whereas residual clinopyroxene is present in many MORP.
Orthopyroxenes in MORP (Tartarotti et al., 2002; Brunelli et al., 2003; Fig. 5) are, on
ee
average, less depleted in REE than those in OI1-u (e.g. YbN: 0.17-1.4 in MORP, YbN: <0.050.41 in OI1-u). Olivine in MORP has very uniform compositions with average Fo contents
rR
between 89.6 (Central Indian Ridge) and 90.7 (Mid-Atlantic Ridge; Fig. 12), whereas most
cores of olivine porphyroclasts in OI1-u have Fo contents between 91 and 92 (e.g., Hierro,
ev
Lanzarote, Kerguelen, Loihi). OI-u xenolith suites that include a significant proportion of
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samples with hydrous minerals or reaction textures (OI2-u) show a wider range in Fo contents
and lower Fo averages (e.g., La Palma, Tenerife, Cape Verde) than OI1-u suites. The tails
towards lower Fo can be explained by metasomatic overprinting with some addition of Fe in
OI2-u. Clear peaks at Fo contents between 92 and 93 (e.g. Gran Canaria and Samoa) in some
islands with heterogeneous olivine compositions reflect the ultra-refractory nature of the
protolith (OI1-u). Olivines in OI-f from Hawaii and Tahiti show wider ranges in Fo, and
lower averages than OI1-u.
The proportion of ultra-refractory peridotites among ocean island xenoliths is high.
Out of 241 OI peridotites 68% have Fo contents 90.5, i.e. Fo at least as high as or higher
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24
than average MORP. If the OI-f xenoliths from Hawaii and Tahiti are excluded, 81% of 191
xenoliths have Fo 90.5 (Fig. 12). These numbers illustrate that a significant proportion of
mantle sampled by ocean island volcanism is more refractory than MOR peridotites. In spite
of chemical changes induced by post-melting processes in MORP, the observed differences in
whole rock and mineral compositions (Figs. 4, 5, 6, 12) suggest that, on average, OI1-u
peridotites have experienced higher degrees of partial melting than MORP.
However, although MORP, on average, are significantly less refractory than OI1-u,
some MORP are as refractory as OI1-u (Fig. 10). This is, for example, true for domains along
rP
Fo
the East Pacific Rise and Mid-Atlantic Ridge (Niu & Hékinian, 1997a; Niu & Hékinian,
1997b). Niu & Hékinian (1997b) proposed that the extent of mantle melting beneath normal
mid-ocean ridges increases with increasing spreading rate. This may explain the presence of
ee
ultra-refractory peridotites along the East Pacific Rise. However, some of the most refractory
MORP known have been collected along one of the slowest-spreading segments of the Mid-
rR
Atlantic Ridge (north and south of the Fifteen-Twenty Fracture Zone; Harvey et al., 2006;
Godard et al., 2008). The formation of these ultra-refractory MORP is clearly not related to
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fast spreading.
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The presence of OI1-u peridotites along mid-ocean ridges and the compositional
overlap between OI1-u and the most refractory MORP (Figs. 4, 5, 10, 11) suggest that OI1-u
may form by mid-ocean ridge processes although this would require potential mantle
temperatures of at least 1400 ºC (Fig. 13). However, available Re-Os data on some OI1-u
(and some ultra-refractory MORP) are not compatible with a simple ridge-melting model.
Alternative models for the formation of OI1-u are discussed below.
The origin of ultra-refractory peridotites
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25
If some (or all) OI1-u did not form by ridge processes, what is the origin of these rocks and
how did they acquire their highly refractory nature? Different sets of data which throw light
on these problems are discussed in the next sections.
Old Re-Os isotope model ages in OI peridotites
rP
Fo
There is increasing evidence that many peridotite xenoliths collected in ocean islands are
older than the oceanic crust in the area. Unradiogenic Os isotope compositions and old Os
model ages are found for some of the most refractory OI1-u harzburgites, and in OI2 type
ee
xenoliths from the Kerguelen Islands (<1.6 GPa; Hassler & Shimizu, 1998; Delpech, 2004), a
group of lherzolites (OI-f) from Sal, Cape Verde ( 3.32±0.15 Ga; Bonadiman et al., 2005),
rR
OI1-u and OI2-u xenoliths from the Canary Islands ( 650 Ma; Table 2), and in spinel
lherzolite xenoliths (OI-f) from different localities in Oahu, Hawaii (90 Ma to 2 Ga; Griselin
ev
& Lassiter, 2001; Bizimis et al., 2004; Bizimis et al., 2007). Evidence of an old depleted
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mantle component has also been found in Azores lavas (2.5 Ga; Schaefer et al., 2002).
Unradiogenic Os isotope ratios and high and heterogeneous Os model ages (up to
2
Ga) have also been reported for MORP from different ridge segments (e.g. Brandon et al.,
2000; Alard et al., 2006; Harvey et al., 2006). These data include some of the ultra-refractory
peridotites collected along the Mid-Atlantic ridge (Harvey et al., 2006). According to Harvey
et al. (2006), a possible explanation of the lack of evidence of such reservoirs in the Re-Os
chemistry of MORB may be that ultra-refractory segments of the upper mantle may be
resistant to further melting (sterile).
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26
The old-age peridotites sampled as ocean island xenoliths and along mid-ocean ridges
may represent recycled continental or oceanic lithosphere, or sub–arc wedge material. In
neither case are they complementary to local MORB, but represent material “exotic” in the
local oceanic mantle lithosphere. Alternatively, the ultra-refractory harzburgites could be
residues of large melting events episodically occurring over the history of the Earth. This
interpretation seems to be supported by a correlation between peaks in Re-Os age distribution
of Os-rich alloy grains measured in peridotite massifs and peaks in crustal ages (Pearson et
al., 2007). The tectonic setting of these large melting events, however, remains poorly
rP
Fo
constrained but could be related to large mantle upwellings following slab avalanches into the
lower mantle (Condie, 1998). In any case, the degree of melt depletion observed in OI1-u
exceeds the degree of melting at normal mid-ocean ridges. Hence, the formation of OI1-u
ee
requires either significantly hotter potential mantle temperatures or the availability of fluid
flux during melting in order to lower the solidus.
rR
Brandon et al. (2000) and Harvey et al. (2006) interpreted the old ages as evidence
that ancient melt depletion may be preserved over time in the convecting mantle. Old, ultra-
ev
refractory peridotites may thus be present in the convecting mantle and may be incorporated
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in the oceanic lithosphere, either during passive advection along mid-ocean ridges or by
freezing to its base as the mantle cools and the lithosphere thickens away from the mid-ocean
ridge. Alternatively, fragments of ultra-depleted peridotites may be brought to the base of the
abyssal lithosphere from the convecting mantle by rising plumes. However, the presence of
old, “exotic” OI1-u (and OI1-f) type material in the abyssal lithospheric mantle does not
exclude the possibility that some highly refractory OI1-u type peridotites may form by partial
melting along fast-spreading mid-ocean ridges, as proposed by Niu & Hékinian (1997a;
1997b).
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A plume origin?
It has been suggested previously that refractory peridotites with old Re-Os ages are
transported from the deeper mantle to lithospheric depths by rising plumes (Bizimis et al.,
2007). Such a mechanism supports the model by Parman (2007) that the unradiogenic He
isotope signature of many ocean island lavas originates from an old, depleted peridotite
component that is stored somewhere in the deep mantle, probably close to the core mantle
rP
Fo
boundary. Since P-T estimates can give important constraints on the evolution of OI-u we
investigate these and other models by looking at the temperatures and pressures at which the
ocean islands xenoliths last equilibrated.
ee
Available data imply that OI-u xenoliths last equilibrated at relatively low pressures
and temperatures. Minimum pressures obtained from geothermometry combined with
rR
densities of CO2 inclusions in mantle xenoliths from the Canary Islands, Samoa, Tahiti,
Kerguelen and Grand Comore fall in the range 0.7 to 1.25 GPa (Hansteen et al., 1991;
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Schiano et al., 1992; Schiano & Clocchiatti, 1994; Schiano et al., 1994; Neumann et al.,
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1995; Hansteen et al., 1998; Schiano et al., 1998). Grégoire et al. (2000a) estimated a
minimum pressure of formation of 1.3 GPa for vein-forming minerals in spinel harzburgite
xenoliths from Kerguelen. A few refertilized samples from Kerguelen and Lanzarote show
breakdown of spinel to plagioclase (e.g. Neumann et al., 1995; Delpech et al., 2004). Some
OI1 peridotites from Fuerteventura and Lanzarote (easternmost Canary Islands) show
petrographic evidence of serpentinization, later followed by heating and dehydration, and
clearly originate in the shallower parts of the mantle lithosphere, close to the MOHO (Abu ElRus et al., 2006).
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According to Schiano & Bourdon (1999) theoretical models of the elastic behavior of
silicate melt inclusions in mantle xenoliths indicate that only limited decompression will take
place in CO2 inclusions during ascent of the host peridotite. Based on a comparison between
the compositions of experimentally rehomogenized inclusions in peridotite xenoliths from
ocean islands (Azores, Kerguelen and Comores islands) and experimental data, Schiano &
Bourdon (1999) concluded that the original melts in the inclusions represent near-solidus
melts in equilibrium with peridotite assemblages at pressures of ca. 1 GPa. The xenolith series
from Hawaii comprise both plagioclase-spinel and spinel lherzolites and harzburgites; the
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highest pressures, 2.1 GPa, were obtained from spinel peridotites from Oahu (Yang et al.,
1998; Sen et al., 2005). All these pressure estimates would place the OI1 samples in the upper
to middle part of an intact, mature (>80 Ma) oceanic lithosphere (Fig. 13).
ee
This conclusion, however, assumes that the oceanic lithosphere is not deeply eroded
by the OI plume. If a significant portion of the original 60-100 km thick lithosphere is eroded
rR
and replaced by plume material, it might be possible that refractory material, which is brought
up by the rising plume, could equilibrate at such shallow depths. Such a model is proposed for
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Hawaiian peridotites that give old Re-Os ages (Bizimis et al., 2007) and is supported by some
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geophysical data (Li et al., 2004) and certain numerical models (e.g. Moore et al., 1999).
A plume origin, however, is hard to reconcile with the low equilibration temperatures
estimated for the OI1-u xenoliths, 860-1200 ºC (Table 1). These temperatures would require
plume material to stagnate at shallow depths for sufficiently long times to allow reequilibration and phase changes during cooling to the low (lithospheric) temperatures
recorded in the xenoliths. Moore et al. (1999) have shown numerically that significant
lithosphere erosion only happens if the plume is at least 140° C hotter than the background
temperature. Such very hot temperatures should leave a trace in the peridotites (e.g. resetting
of the Hf isotopic clock), which is not observed (Bizimis et al., 2007). Furthermore, if the
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OI1-u were part of the rising plume and transported to pressures below 1.3 GPa, they should
not only be hot, but also show more sign of melt infiltration and deformation. Even though the
OI1-u harzburgites probably would not melt because of their high solidus, more fertile plume
rocks would definitely be partially molten at 1.3 GPa. We would therefore expect plume melts
to have invaded stretched and folded OI1-u fragments (Farnetani et al., 2002). Because of
their strong depletion, OI1-u are highly susceptible to reaction with/contamination by
enriched melts and would react quickly. We do not observe this. Furthermore, OI2-u
xenoliths, which do show melt–peridotite reactions, tend to give higher average equilibration
rP
Fo
temperatures than OI1-u xenoliths from the same island (Table 2). This suggests local heating
of cold, lithospheric peridotites by plume melts rather than regional cooling of accreted plume
material.
ee
Another argument against transport of OI-u peridotites from the deep mantle are their
low densities (Simon & Podladchikov, 2008). Since the ultra-refractory OI-u peridotites are
rR
strongly buoyant such material is not expected to be stored for long periods of time deep in
the lower mantle, but rather to float to the top of the convecting mantle. OI1-u material is
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therefore unlikely to be present in large quantities in the source region of deep, thermally
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activated plumes. However, some plumes may themselves consist of low density residual
mantle that has been subducted to some depth and has become buoyant relative to the
surrounding more fertile peridotite when heated to ambient temperatures (Oxburgh &
Parmentier, 1977; Presnall & Helsley, 1982; Lee & Chen, 2007).
Evidence against a plume origin for OI-u xenoliths is also provided by Sr isotope data
on xenoliths from the Canary Islands. In general, Sr, Nd, Pb, Hf isotope analyses of OI mantle
xenoliths have isotopic signatures similar to those of the local ocean island basalts (e.g.,
Vance et al., 1989; Whitehouse & Neumann, 1995; Mattielli et al., 1996; Mattielli et al.,
1999; Sen et al., 2003; Neumann et al., 2004). Sr-Nd-Pb isotope data on whole rock samples
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and clinopyroxene separates from mantle xenoliths from the Canary Islands follow this
general pattern (Whitehouse & Neumann, 1995; Neumann et al., 2004; Fig. 14). However, insitu laser ablation analyses of
87
Sr/86Sr in clinopyroxenes show a range in isotope ratios
(Neumann et al., 2004). Clinopyroxenes in OI1-u xenoliths have
87
Sr/86Sr isotope ratios of
0.70266 and 0.70276, which is significantly below those of Canary Islands basalts and
requires long term evolution in a depleted, low Rb/Sr environment. The ultra-refractory nature
of the Canarian OI1-u peridotites must thus be an old feature caused by processes unrelated to
the present plume. One OI2-u sample gave the intermediate value of 0.70288. Only
rP
Fo
clinopyroxene in OI2-u xenoliths gave ratios of 0.7031-0.7033, which, like those of the whole
rock and clinopyroxene separates, is within the range of Canary Islands basalts field. The
original old depleted isotopic signature has therefore recently been overprinted by
ee
metasomatism by melts/fluids originating in the Canary Islands plume (Neumann et al.,
2004).
rR
In conclusion, we consider a plume origin for the ocean island xenoliths as highly
unlikely. Based on the P-T estimates, the inferred physical properties (high viscosity and low
ev
density) and the in-situ Sr isotope data, we favor the idea that the OI peridotites were part of
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the oceanic lithosphere before the onset of the plume processes that brought them to the
surface.
Did the OI1 peridotites acquire their highly depleted nature by plume processes?
Although we concluded above that the OI-u peridotites are not brought up by the plume, the
lithospheric mantle beneath ocean islands is clearly affected by plume–related processes. The
most obvious one is metasomatism by plume-derived melts and fluids, which is reflected in
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OI2 xenoliths, and which has also caused some enrichment in incompatible elements and
radiogenic isotopes in OI1 xenoliths (e.g. Figs. 1-6, 14). However, plume-related processes
have also been proposed to be a cause of more refractory mantle compositions by the
following processes: (1) reactions between plume melts/fluids and peridotite wall–rock may
extract lithophile components from the wall–rock minerals and cause orthopyroxene to form
at the expense of olivine and clinopyroxene (e.g. Kelemen et al., 1992); (2) the plume may
lead to thermal erosion of the base of the oceanic lithosphere and trigger partial melting in its
remaining parts (e.g., Moore et al., 1999; Li et al., 2004; Bizimis et al., 2007).
rP
Fo
Microtextural evidence of melt-wall–rock reactions is common in OI2 peridotites.
They show the reaction (opx+melt
ol+cpx). This is opposite to the reaction described by
Kelemen et al. (1992), but is compatible with infiltration by transitional to strongly alkaline
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ocean island melts saturated in clinopyroxene. Infiltration of ocean island xenoliths by
alkaline melts typically leads from harzburgite (OI1-u)
rR
lherzolite (OI2-u)
dunite/wehrlite. In addition, the Ti-Al-Fe-rich nature of clinopyroxenes in some OI dunites,
wehrlite and clinopyroxenites suggests crystallization directly from alkali basaltic magmas
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(e.g. Neumann et al., 2004). All these reactions are by definition absent from OI1 xenoliths.
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Finally, interaction between peridotites and melts is expected to reset the Sr isotope ratio in
the peridotites to those of the melts. As reported above, this is seen in all whole rock samples
and clinopyroxene separates in OI-u xenoliths from the Canary Islands, as well as in in-situ
analyses of clinopyroxenes in OI2-u rocks (Neumann et al., 2004). However, the low 87Sr/86Sr
ratios obtained from in-situ analyses of clinopyroxenes in Canarian OI1-u xenoliths (Fig. 14)
are robust evidence against interaction with plume melts.
The Sr isotope ratios in Canarian OI1-u rocks do not exclude the possibility that
heating by a thermal plume caused partial melting of the lithospheric mantle. However, the
low equilibration temperatures obtained for the OI1-u (and OI2-u) xenoliths make this
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possibility highly unlikely. In Figure 13, the estimated P-T conditions for the ultra-refractory
xenoliths are compared to estimated solidi for peridotites of average OI1-u and more fertile
compositions. The equilibration temperatures obtained from the OI-u xenoliths are far below
their estimated solidus temperatures. Heating and melting related to the present plumes
therefore seems an implausible explanation for the formation of the OI1-u harzburgites. Since
neither plume-induced melting nor reaction with plume melts appear to be responsible for the
ultra-refractory compositions of the OI1-u peridotites we conclude that this composition was
acquired a significant period of time before the onset of plume processes. This is in agreement
rP
Fo
with the conclusions that the old Os model ages obtained in some ultra-refractory OI
peridotites and MORP reflect ancient melting events (e.g. Hassler & Shimizu, 1998; Brandon
et al., 2000; Griselin & Lassiter, 2001; Delpech, 2004; Bizimis et al., 2004; Bonadiman et al.,
ee
2005; Alard et al., 2006; Harvey et al., 2006; Bizimis et al., 2007).
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Comparison of OI1-u to refractory peridotites from other settings
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Although some OI-u peridotites may have formed along mid-ocean ridges, we are left with
the hypothesis that many OI1-u peridotites represent old, refractory material that has been
passively accreted to the oceanic lithosphere either along mid-ocean ridges or during cooling
and thickening. By which kind of process did the latter acquire their ultra-refractory
composition if they were not formed in their host plume? We compare OI1-u compositions to
those of sub-continental mantle xenoliths and peridotites from oceanic sub-arc mantle (Figs.
15 and 16) in order to determine if there are similarities that might suggest a related origin.
Peridotite suites from both types of setting include highly refractory samples, whose large
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Page 33 of 87
33
degree of melt extraction has been ascribed to either higher mantle temperatures, flux by
hydrous fluids, or both (see e.g. review by Pearson et al., 2003).
The most refractory sub-continental mantle xenoliths are cratonic peridotites from
South Africa (Boyd & Mertzman, 1987; Cox et al., 1987; Nixon et al., 1987; Winterburn et
al., 1990; Boyd et al., 1999; Simon et al., 2003; Simon et al., 2007) and East Greenland
(Bernstein et al., 1998). However, these xenoliths differ significantly from the OI1-u
xenoliths by lower FeO*, and higher SiO2 and MgO (Fig. 15). Figure 15 also shows the
comparison of OI1-u xenoliths to some refractory off-craton continental mantle xenoliths,
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Fo
which are characterized by large scatter in major elements. Some off-craton xenoliths fall
within the OI1-u field, but the majority are more fertile than OI1-u. We did not, however,
filter the continental peridotite data for metasomatized samples which might explain a large
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part of the scatter. We can therefore not exclude the possibility that OI1-u represent fragments
of continental lithospheric mantle that was eroded from the base of the continental lithosphere
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or sheared off during rifting and passive margin formation. The latter scenario has been
proposed as the ‘raft’ model by Müntener & Manatschal (2006) to explain the occurrence of
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some strongly refractory spinel peridotites in the apparently oceanic domain at the
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Newfoundland margin. Such a model, however, does not explain how the presumed
continental lithospheric mantle reached its high degree of depletion. Müntener & Manatschal
(2006) proposed that the original melt depletion event took place in an arc setting. Extensive
partial melting in an arc setting has also been proposed as a mechanism to form Archean
cratonic lithosphere (e.g. Kesson & Ringwood, 1989; De Wit et al., 1992; Westerlund et al.,
2006; Simon et al., 2007).
In Figure 16 we compare OI1-u to some mantle peridotite series from modern oceanic
sub-arc settings. As a group, the arc samples also show considerable compositional variations,
but the peridotite suites from the Izu-Bonin-Mariana forearc, the Torishima Forearc Seamount
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34
(data from Parkinson & Pearce, 1998) and from Lihir, Papua New Guinea (data from McInnes
et al., 2001) are strikingly similar to OI1-u. This similarity suggests that OI1-u type
peridotites may form by similar processes. A more detailed comparison between OI1-u
peridotites and sub-arc mantle is given elsewhere (Neumann & Simon, 2008).
In conclusion, the highly refractory OI1-u xenoliths most likely represent residues of
shallow melting at elevated temperatures, water contents, or both. Following melting they
have been recycled to the upper part of the convecting mantle, where they circulated and
became accidentally trapped in the oceanic lithosphere before they were sampled by the ocean
rP
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island magma. However, some may have formed along hot mid-ocean ridges and been
directly incorporated in the oceanic mantle lithosphere.
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CONCLUSIONS AND IMPLICATIONS FOR GEODYNAMICS
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The major element compositions of oceanic island peridotites have been reviewed. After
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petrographically filtering for the effects of mantle metasomatism and melt-rock interaction
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during transport, a large proportion of peridotite xenoliths sampled in ocean islands proved to
be ultra-refractory harzburgites. These samples are characterized by partial melting beyond
the cpx-out reaction and are on average significantly more refractory (higher Fool and cr#sp,
lower CaO and (Al2O3)opx) than abyssal peridotites sampled by dredging and drilling along
mid-ocean ridges and at fracture zones. We have shown that the strong melt depletion in the
OI1-u harzburgites is not related to melt-rock interaction, and that the OI1-u are not
genetically related to their host ocean island plume. These rocks must thus be residues of
large melting events at temperatures significantly higher than present day normal potential
mantle temperatures, or they are residues formed during extensive fluid-fluxed melting,
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35
possibly in an arc setting. The last possibility is supported by their compositional similarity to
modern oceanic fore–arc peridotites. If the melt depletion of the OI1-u is caused by fluidfluxed melting the textures related to the fluid-rock interaction, which are commonly
observed in recent arc-related peridotites, have been erased by complete textural and chemical
re-equilibration.
We propose that ultra-refractory OI1-type harzburgites represent the end-product of
mantle differentiation, in the sense that they have reached a maximum degree of melt
extraction. Although we cannot provide an estimate on how much of the mantle consists of
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strongly refractory harzburgite, sterile OI-u material makes up more than two thirds of all
harzburgites and lherzolites sampled by ocean island magmas (Fig. 12) and is thus common in
the mantle beneath ocean islands. This is in accordance with a model of a strongly
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heterogeneous mantle, as supported by recent geochemical and geophysical observations and
theoretical models (e.g., Phipps Morgan & Morgan, 1999; Brandon et al., 2000; Helffrich &
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Wood, 2001; Meibom et al., 2002; Anderson, 2006; Harvey et al., 2006; Helffrich, 2006;
Armienti & Gasperini, 2007; Sobolev et al., 2007). The presence of such material has
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consequences for geodynamics due to its distinct geochemical and geophysical properties.
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Strong depletion in fusible components (and increase in forsterite content), as in OI1-u
peridotites, leads to high solidus temperatures and decreased water contents, which cause
higher viscosities and greater resistance to deformation (Karato, 1986; Hirth & Kohlstedt,
1996). Manga (1996) has shown numerically that highly viscous material (10-100 times more
viscous that the surrounding mantle, such as strongly melt-depleted peridotite; Karato, 1986;
Hirth & Kohlstedt, 1996) can persist relatively undisturbed for many mantle overturns in the
convecting mantle and has a tendency to aggregate to form larger heterogeneities from
smaller blobs. Because of their high solidus temperatures such rocks will be “sterile” and are
not expected to participate in further partial melting. These rocks are therefore not expected to
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36
lend their geochemical signatures to partial melts formed either along mid-ocean ridges or in
ocean islands, except if temperatures are extraordinarily high (Fig. 13).
Due to their high Mg/Fe ratios and low Al contents, the OI1-u harzburgites will have
relatively low densities and therefore strong buoyancy, and will tend to rise above less
refractory, denser mantle material. The accumulation of refractory mantle at the top of the
convecting mantle over time is confirmed by 3-D mantle convection simulations (Tackley &
Xie, 2002). Highly refractory material may therefore be abundant in regions where new plates
form and may preferentially be incorporated into the cooling and growing oceanic
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Fo
lithospheric plate. We thus expect the proportion of highly refractory harzburgites to be
significantly higher in the oceanic mantle lithosphere and the underlying asthenosphere, than
in the deeper parts of the convecting mantle, in agreement with the model of Phipps Morgan
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& Morgan (1999). An important consequence of a high proportion of low density OI1-u
material in the oceanic plate is that it will be quite difficult to subduct, particularly in the
initial
stages
of
subduction.
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Some
preliminary
calculations
with
Perple_X
(http://www.perplex.ethz.ch/perplex.html, see Connolly & Petrini, 2002, for methods) show
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that a mantle column (100 km tall) with the average composition of the Canary Island OI1-u
and a 6.5 km thick basaltic crust (
crust
= 2950 kg/m3) will have a density of 3200 kg/m3 close
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to the ridge, compared to 3220 kg/m3 for the average composition previously given for
oceanic peridotite (Table 4). Thickening of the lithosphere by simple half space cooling leads
to a 100 km thick, old, thermally equilibrated lithosphere with average column densities of
3245 kg/m3 and 3263 kg/m3 for Canarian OI1-u and average MORP, respectively. Afonso et
al. (2007) developed a more sophisticated model for the calculation of oceanic lithosphere
density structure and buoyancy over time. Their model results in an 106 km thick mature
lithosphere, where the upper 60-80 km consist of variably melt-depleted material and the
lower 20-40 km are made up by fertile mantle, depending on the melt model used. Afonso et
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al. (2007) conclude that such a lithosphere develops smaller negative buoyancies (~ 40 kg/m3)
than commonly assumed, and thus that the role of intrinsic negative buoyancy may not be as
critical for the onset of subduction. If the oceanic lithosphere contains a significant amount of
ultra-refractory peridotite, as proposed here, buoyancy will be further enhanced, making
initiation of subduction even more difficult. This is, of course, assuming that the
asthenosphere on average consists of more fertile, denser mantle, which may not be realistic
given the presumably large amount of buoyant recycled residues in the convecting mantle. If
the asthenosphere is rich in ultra-refractory material the cooling oceanic lithosphere may
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Fo
become negatively buoyant because its upper to middle part is made up of denser (relative to
OI1-u) MORP. Lenses of material with different density in the lithosphere will also enhance
the formation of gravitational instabilities and small scale convection. Better constraints on
ee
models for the composition and stratification of the oceanic lithosphere may be derived from
comparison of calculated geophysical properties (e.g. gravity, geoid-anomalies, topography,
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seismic velocities) with measured geophysical data (e.g. Afonso et al., 2007).
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ACKNOWLEDGEMENTS
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Dante Canil generously gave us a copy of his data base on mantle rocks, which was the
foundation of our own data base. The work has profited from discussions with a number of
colleagues, in particular Y. Podladchikov, L. Rüpke, J. Phipps Morgan, O. Alard, J. C.
Afonso and E. Hellebrand. H. Dawson, G. Pearson and two anonymous reviewers are thanked
for their comments that help to focus the ideas expressed in this paper. We acknowledge the
Norwegian Research Council (NFR) for financial support, including a post-doctoral
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38
fellowship to Nina S.C. Simon. The work has also been supported by a Center of Excellence
grant from the Norwegian Research Council to PGP.
Appendix A and B. Supplementary material
The
databases
on
which
this
paper
is
based
http://www.petrology.oxfordjournals.org.
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Figure captions
Fig. 1. Photomicrographs of typical OI2 textures. (a) Melt-rock reaction in spinel harzburgite
(Canary Islands) leading to growth of clinopyroxene (cpx) and olivine (ol) at the expense of
orthopyroxene (opx; g: glass; crossed Nicols). (b) Poikilitic opx in spinel lherzolite (Canary
Islands) enclosing corroded remnants of larger olivine grains and small oval olivine grains
(crossed Nicols). (c) Poikilitic harzburgite (Kerguelen) with poikilitic cpx with spongy rims
produced by interaction between protogranular harzburgites and alkaline, silicaundersaturated melt (crossed Nicols). (d) Poikilitic lherzolite (Canary Islands) showing cpx,
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locally with spinel exsolution lamellae, enclosing oval opx grains and spinel (crossed Nicols).
Spinel occurs as singly grains and as clusters or chains of very small grains. (e) Symplectitic
intergrowth between cpx and spinel along grain boundaries in spinel harzburgite (Canary
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Islands). Small olivine domains (orange and bluish green) within the symplectite are in optic
continuation with the neighboring olivine porphyroclasts (crossed Nicols). The texture
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suggests formation of the symplectite and cpx by reaction between melt and peridotite. (f)
Spinel lherzolite from Cape Verde showing sieve textured cpx and opx.
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Fig. 2. Photomicrographs of typical OI1 textures. (a) Overview of protogranular texture in
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OI1-u xenolith (Kerguelen) with olivine with undulatory extinction and exsolved opx (crossed
Nicols). (b) Details of protogranular texture in spinel harzburgite (Cape Verde; crossed
Nicols). (c) Opx porphyroclasts with cpx exsolution lamellae in spinel harzburgite (Canary
Islands; crossed Nicols). Locally cpx exsolution lamellae have developed into larger domains
(cpx2). Small equant cpx2 are scattered along the rims of the opx porphyroclasts. (d) Spinel
lherzolite from Cape Verde with primary cpx.
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Fig. 3. OI1-u whole rock major element variations [Canary Islands: (Neumann, 1991; WulffPedersen et al., 1996; Neumann et al., 2002; Neumann et al., 2004; Abu El-Rus et al., 2006);
Kerguelen: (Grégoire et al., 1997; Grégoire et al., 2000a; Grégoire et al., 2000b; Delpech,
2004); Cape Verde: (Bonadiman et al., 2005; Bonadiman, unpublished data, 2008)]. Dotted
lines enclose the ranges covered by OI2-u xenoliths. The olivine (ol) and orthopyroxene (opx)
fields and ol-opx tie-line are based on mineral data from Lanzarote. Blue and red symbols and
arrows show the results of mantle melting experiments (Green et al., 2004). The blue arrows
connect source (H: fertile Hawaiian pyrolite; T: Tinaquillo depleted lherzolite; MM3:
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lherzolite) and residual compositions (percent melting given by blue numbers) in “first-stage
melting” experiments, aimed at modeling residues after extraction of MORB melts. The red
arrows connect to residues of “second-stage” experiments (additional 10%, H2O-CO2-fluxed
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partial melting of a residual chromite-bearing harzburgite at 1.5 GPa and 1320-1350 ºC). See
text for discussion.
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Fig. 4. Relationships between Fo-contents and Al2O3 in the cores of coexisting olivine and
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orthopyroxene porphyroclasts, respectively, and cr#-mg# in the cores of equidimensional
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spinel grains in ultra-refractory OI1-u (upper panels) and more fertile OI1-f (lower panels).
The ranges covered by OI2 from the different islands are indicated by colored dotted lines.
Data sources Canary Islands: Neumann (1991), Neumann et al. (1995; 2002), Wulff-Pedersen
et al. (1996), Abu El-Rus et al. (2006); Madeira: Munha et al. (1990); Azores: Johansen
(1996); Cape Verde: Ryabchikov et al. (1995), Bonadiman et al. (2005), Shaw et al. (2006);
Grande Comore: Coltorti et al. (1999); Kerguelen: Grégoire et al. (1997), Delpech (2004);
Samoa: Hauri & Hart (1994); Hawaii: Goto & Yokoyama (1988), Sen (1988), Bizimis et al.
(2007); Loihi seamount: Clague (1988); Tahiti: Tracy (1980), Qi et al. (1994). Grey fields:
compositional ranges of minerals in MORP (data from Hamlyn & Bonatti, 1980; Komor et
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Page 63 of 87
63
al., 1990; Bonatti et al., 1992; Johnson & Dick, 1992; Cannat et al., 1995; Snow & Dick,
1995; Arai & Matsukage, 1996; Dick & Natland, 1996; Niida, 1997; Ross & Elthon, 1997;
Stephens, 1997; Hellebrand et al., 2002; Brunelli et al., 2003; Hellebrand & Snow, 2003;
Seyler et al., 2003; Brunelli et al., 2006). Since the ol-opx partition coefficient for
Mg/(Mg+Fe) is close to unity for MORP (Seyler et al., 2003) we have used [Mg/(Mg+Fe)]opx
instead of Fool where olivine data are not available for the MORP. Stars marked “exp”: the
composition of minerals in the experimental “second-stage” residue of Green et al. (2004).
See text for discussion.
rP
Fo
Fig. 5. Rare earth element patterns (normalized to Primordial Mantle; McDonough & Sun,
1995) for orthopyroxene porphyroclasts in (a) OI1-u. Data sources, Canary Islands: Neumann
ee
et al. (2002; 2004); Cape Verde: Bonadiman et al. (2005), Bonadiman, unpublished data
(2008); Kerguelen: Grégoire et al. (2000b), Delpech (2004); (b) OI1-f from Cape Verde
rR
(Bonadiman et al., 2005). Shaded field: orthopyroxene porphyroclasts in MORP (data from
Tartarotti et al., 2002; Brunelli et al., 2003; Hellebrand et al., 2005).
ev
Fig. 6. REE patterns for clinopyroxenes in OI1. (a) OI1-u from the Canary Islands (Neumann
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et al., 2002; Neumann et al., 2004), Cape Verde: Bonadiman et al. (2005), Bonadiman
(unpublished data, 2008), Kerguelen (Mattielli et al., 1999; Grégoire et al., 2000a; Grégoire et
al., 2000b; Moine et al., 2000; Delpech, 2004), and Samoa (Hauri et al., 1993). (b) OI1-f
from Cape Verde (Bonadiman et al., 2005), Hawaii (Sen et al., 1993; Yang et al., 1998) and
Tahiti (Qi et al., 1994).
Fig. 7. Relationships between modal proportions of clinopyroxene and olivine in mantle
xenoliths from Cape Verde (data from Bonadiman et al., 2005; Bonadiman, unpublished data,
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Manuscript submitted to Journal of Petrology
64
2008). The different positions and slopes of linear regression lines R1 and R2 calculated for
the harzburgites and lherzolites, respectively, strongly support suggestions by Doucelance et
al. (2003) and Bonadiman et al. (2005) that the harzburgites and lherzolites represent two
unrelated rock suites.
Fig. 8. Major element variations of OI1-f from Cape Verde (Bonadiman et al., 2005;
Bonadiman, unpublished data, 2008), Hawaii (Jackson & Wright, 1970; Sen, 1987; Goto &
Yokoyama, 1988; Sen & Leeman, 1991; Bizimis et al., 2003) and Tahiti (Tracy, 1980; Qi et
rP
Fo
al., 1994). Opx and ol fields and opx-ol tie-lines as in Figure 3. Blue and red symbols and
arrows show the results of mantle melting experiments by Green et al. (2004). Light grey
field: OI1-u.
ee
Fig. 9. Opx/cpx REE ratios for selected samples from the Canary Islands, compared to
rR
published partition coefficients of pyroxenes in equilibrium in spinel harzburgites and
lherzolites (Eggins et al., 1998; Garrido et al., 2000; Green et al., 2000; Hellebrand et al.,
ev
2005). Solid lines: OI1-u; dashed lines: OI2-u. The opx/cpx ratios obtained for OI1 and some
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Page 64 of 87
OI2 fall within the range of experimental data, indicating equilibrium between the two
pyroxenes. The ratios obtained for two most REE-rich OI2 clinopyroxenes with fall far below
the
experimental
range
indicating
REE-enrichment
in
clinopyroxene
relative
to
orthopyroxene.
Fig. 10. Variations in bulk-rock major element compositions of MORP from different ridgetransform systems (data from Myashiro et al., 1969; Aumento & Loubat, 1971; Dick, 1989;
Blum, 1990; Cannat et al., 1995; Snow & Dick, 1995; Casey, 1997; Niu & Hékinian, 1997a;
Stephens, 1997; Brandon et al., 2000; Lee et al., 2003; Seyler et al., 2003; Niu, 2004; Harvey
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Page 65 of 87
65
et al., 2006; Godard et al., 2008). AAR: American-Antarctic Ridge-transform system in the
Indian Ocean; CIR: Central Indian Ridge-transform system; MAR: Mid-Atlantic Ridge; PAR:
Pacific-Antarctic Ridge-transform system; SWIR: Southwest Indian Ridge-transform system;
grey field: OI1-u xenoliths.
Fig. 11. Variations in CaO/Al2O3 ratios of (a) MORP and (b) OI1-u. Data sources as in
Figures 3 and 10. MORP show a strong tendency for lower CaO/Al2O3 ratios than OI1-u,
probably due to serpentinization.
rP
Fo
Fig. 12. Fo-contents in cores in olivine porphyroclasts in OI spinel peridotite xenoliths
compared to those in MORP. Olivine porphyroclasts in OI1-u (dry) xenoliths show peaks in
ee
Fo contents between 91 and 92 (e.g. Hierro, Lanzarote, Cape Verde, Kerguelen). Islands with
a high proportion of samples with hydrous minerals (OI2-u) show wider ranges in Fo contents
rR
and lower averages (e.g. La Palma, Tenerife). Olivine in the fertile OI-f from Hawaii and
Tahiti shows lower ranges in Fo contents. Olivine in MORP has very uniform Fo contents
ev
with average Fo between 89.6 and 90.6. The dashed line shows the MORP average of 90.5.
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Data on OI are from Jackson & Wright (1970), Tracy (1980), Clague (1988), Goto &
Yokoyama (1988), Sen (1988), Munha et al. (1990), Neumann (1991), Hauri & Hart (1994),
Qi et al. (1994), Ryabchikov et al. (1995), Neumann et al. (1995; 2002), Johansen (1996),
Wulff-Pedersen et al. (1996), Grégoire et al. (1997), Coltorti et al. (1999), Delpech (2004),
Bonadiman et al. (2005), Shaw et al. (2006). Data on MORP are from Hamlyn & Bonatti
(1980), Dick (1989), Komor et al. (1990), Snow & Dick (1995), Niida (1997), Ross & Elthon
(1997), Stephens (1997), Brunelli et al. (2003), Seyler et al. (2003).
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Manuscript submitted to Journal of Petrology
66
Fig. 13. Phase diagram showing the stability fields of plagioclase-, spinel- and garnetperidotites for different bulk compositions (Hawaiian pyrolite, average MORP, and average
OI1-u; Table 4). Spinel stability extends into the garnet and plagioclase fields. Phase
boundaries are computed with Perple_X including the recent addition of Cr-bearing endmembers to the Holland and Powell database (Holland & Powell, 1998; Connolly & Petrini,
2002; Klemme et al., 2008). Ultra-refractory OI1-u peridotite is assumed to contain 0.05 wt%
Na2O (Table 4). The solidi are taken from Green & Ringwood (1970) for wet and dry pyrolite
and extrapolated to more refractory compositions using the approximation of Robinson &
rP
Fo
Wood (1998; increase of 7 °C per degree of melt extraction) and assuming that OI1-u have
experienced in total 30% melt extraction from a pyrolite source. The resulting OI1-u solidus
probably gives minimum temperatures for the onset of melting since OI1-u have lost all cpx,
ee
which will increase solidus temperatures further than estimated by the simple extrapolation
used here. The diagram shows that the calculated garnet-in boundary is shifted to significantly
rR
higher pressures for OI1-u. Also shown is the simplified geotherm at a mid-oceanic ridge. The
temperature at the MOHO is taken as 550 °C and T at the base of the lithosphere (Tast) is 1330
ev
°C. Grey fields show the ranges of pressure and temperature estimates for OI-u (Table 2).
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Page 66 of 87
Fig. 14. Sr-Nd isotope data on whole rock samples (Vance et al., 1989; Rolfsen, 1994;
Ovchinnikova et al., 1995) and clinopyroxene separates (Whitehouse & Neumann, 1995;
Neumann et al., 2002), compared to laser ablation microprobe analyses of 87Sr/86Sr ratios in
clinopyroxenes in OI1 and OI2 from the Canary Islands (Neumann et al., 2004). Fields of
Canary Islands basalts (CI basalts; data from Hoernle et al., 1991; Hoernle et al., 1995;
Ovchinnikova et al., 1995; Thirlwall et al., 1997; Simonsen et al., 2000), and mid-Atlantic
Ridge N-MORB (MAR; Ito et al., 1987; Dosso et al., 1991) are shown for comparison. In situ
Sr isotope ratios in clinopyroxenes in two OI1 samples (solid lines) show depleted Sr isotope
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Page 67 of 87
67
ratios below those of the Canary Islands basalts. Laser ablation
87
Sr/86Sr of OI2 cpx (dashed
lines) and a wehrlite cpx (dotted line) overlap with the Canary Island basalt field.
Fig. 15. Comparison between OI1-u xenoliths (grey fields) and continental mantle xenolith
suites with a high proportion of harzburgites. Data from The Thumb, New Mexico
(Ehrenberg, 1982), Patagonia, South America (Skewes & Stern, 1979; Laurora et al., 2001),
NE Brasil (Rivalenti et al., 2000), Sardinia (Dupuy et al., 1987; Beccaluva et al., 2001),
Namibia (Boyd et al., 2004), East Greenland (Bernstein & Brooks, 1998) and South Africa
rP
Fo
(Boyd & Mertzman, 1987; Cox et al., 1987; Nixon et al., 1987; Winterburn et al., 1990; Boyd
et al., 1999; Simon et al., 2003; Simon et al., 2007). The off-craton mantle xenolith series
show wide compositional ranges, including fertile compositions.
ee
Fig. 16. Comparison between OI1-u xenoliths (grey fields) and mantle peridotites from the
rR
Izu-Bonin-Mariana forearc, the Torishima Forearc Seamount (ODP Leg 125; data from
Parkinson & Pearce, 1998) and Lihir (Papua New Guinea; data from McInnes et al., 2001).
ev
These forearc peridotites (individual symbols) have major element compositions similar to
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Manuscript submitted to Journal of Petrology
OI1-u. Also shown is the compositional range covered by the majority (>95%) of peridotites
from active oceanic arcs (dashed line; data from Berry, 1981; Jaques et al., 1983; Aoki, 1987;
Pearce et al., 2000; Takazawa et al., 2000; Franz et al., 2002; in addition to the references
cited above).
http://www.petrology.oupjournals.org/
Manuscript submitted to Journal of Petrology
Table 1: Summary of textural and geochemical characteristics of the different groups.
Ultra-refractory (OI-u)
Fertile (OI-f)
OI1-u
- porphyroclastic to protogranular
- no petrographic evidence of melt-rock interaction (Fig.
2)
- deformed porphyroclasts of ol+opx
- matrix: less deformed or undeformed neoblasts of
ol+opx±sp±cpx
- <3 vol% cpx
- primary cpx appears absent
- high WR MgO: 44-48 wt%
- low WR CaO: 0.4-1.2 wt%
- low WR Al2O3: 0.2-1.3 wt%
- (Al2O3)opx: 0.4-4.3 wt%; main range 1-3 wt%
- high Cr#sp: 0.3-0.9; main range 0.4-0.7
- narrow range in high Fool: 89.7-92.6
- low HREEopx YbN: 0.04-0.4
- LREE-depleted to U-shaped REEopx-patterns
OI1-f
- textures as in OI1-u (Fig. 2)
- no petrographic evidence of melt-rock interaction
(Fig. 2)
- deformed porphyroclasts of ol+opx+cpx±sp
- matrix: less deformed or undeformed neoblasts of
ol+opx±sp±cpx
- up to 20 vol% cpx
- primary clinopyroxene common
- low WR MgO: 38-45 wt%
- high WR CaO: 0.9-4.6 wt%
- high WR Al2O3: 0.6-3.7 wt%
- higher (Al2O3)opx: 1.7-6.0 wt%; main range 2-6 wt%
- low Cr#sp: 0.1-0.5; main range 0.1-0.4
- wide range in Fool: 83.6-92.0
- low HREEopx YbN: 0.2-0.4
- LREE-depleted REEopx-patterns
OI2-u
- porphyroclastic/protogranular/poikilitic
- petrographic evidence of melt-rock reactions, modal
metasomatism, etc. (Fig. 1)
- primary cpx appears to be absent
- low WR MgO: 39-48 wt%
- low WR CaO: 0.4-1.2 wt%
- wide range in WR Al2O3: 0.2-3.7 wt%
- wide range in (Al2O3)opx: 0.1-4.0 wt%
- wide range in Cr#sp: 0.2-0.9
- wide range to low Fool: 84.6-91.5
- higher HREEopx YbN: 0.1-2
- LREE-depleted to U-shaped REEopx-patterns
OI2-f
r
Fo
er
Pe
- petrographic evidence of melt-rock reactions, modal
metasomatism, etc. (Fig. 1)
Re
- low WR MgO: 37-45 wt%
- high WR CaO: 0.8-3.9 wt%
- high WR Al2O3 common: 0.6-3.9 wt%
- high (Al2O3)opx: 2.9-6.0 wt%
- low Cr#sp: 0.03-0.3
- wide range to low Fool: 83.6-91.7
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Page 69 of 87
Table 2. Equilibration temperatures estimated for different ocean islands xenoliths using the
Sachtleben & Seck (1991) geothermometer based on CaO in orthopyroxene.
Island
Group
N
Range
Average
Canary Islands
OI1-u
OI2-u
48
41
873-1209
862-1215
1024
1068
Madeira
OI1-u
OI2-u
7
5
891-1017
892-1182
936
1037
Azores
OI2-u
17
1003-1232
1135
Cape Verde
rP
Fo
OI1-u
OI1-f
OI2-f
1
4
2
1047-1233
1028-1227
1163
1190
1127
OI1-u
OI2-u
16
25
933-1104
960-1193
1014
1062
Grande Comore
OI2-u
8
965-1239
1109
Samoa
OI1-u
OI2-u
6
13
983-1142
1008-1228
1063
1072
Hawaii islands
OI1-f
OI2-f
9
23
914-1205
881-1161
1004
1005
Loihi seamount
OI1-u
4
1139-1182
1163
Tahiti
OI1-f
OI2-f
19
1
996-1108
Kerguelen
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Manuscript submitted to Journal of Petrology
1058
1093
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Manuscript submitted to Journal of Petrology
Table 3: Bulk rock Re-Os isotope data for Canary Island harzburgite and lherzolite
xenoliths (Neumann & Widom, unpublished data).
type
LA8-10a
LA1-7
LA1-13
PAT2-68
H1-5a
PAT2-75
TF14-38
TF14-47
TF14-48
OI1
OI1
OI1
OI1
OI2
OI2
OI2
OI2
OI2
Os
(ppb)
Re
(ppb)
2.363
1.251
7.576
1.271
3.631
1.750
2.189
1.206
2.273
0.067
0.003
0.003
0.036
0.002
0.030
0.011
0.012
0.021
187
Os/188Os
(±2 sigma)
0.12409(19)
0.12613(20)
0.12423(11)
0.12554(13)
0.12652(18)
0.12486(16)
0.12303(27)
0.12678(12)
0.12486(32)
187
Re/188Os
0.136
0.010
0.002
0.134
0.002
0.082
0.025
0.049
0.045
Tma
(Ma)
Trd
(Ma)
654.52
134.60
415.75
329.10
73.58
402.09
630.61
39.13
360.77
434.57
131.30
413.81
219.17
73.14
320.28
591.55
34.34
320.87
er
Pe
a
SAMPLE
r
Fo
measured July 1995. All others June 1993 at Department of Terrestrial Magnetism, Carnegie Institution of
Washington. Tma: Re-Os mantle model age. Trd: Re-depletion age, assuming all Re was extracted during
melting.
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Page 71 of 87
Table 4. Different peridotite major element compositions (recalculated to 100% without
Na2O, K2O, P2O5) compared to OI1 averages.
Rock type
N
SiO2
TiO2
Al2O3
Cr2O3
FeO*
MnO
MgO
CaO
Prim. mantle models
HPa
Tinb
Exp
Resc
45.62
0.72
3.57
0.43
8.55
0.14
37.85
3.11
44.42
0.01
0.83
0.25
5.70
Table 4 continued
47.76
1.02
OI1-u averages
Can.Isl.
Cape V.h Kerguel.i
28
7
12
43.76
43.62
45.18
0.02
0.02
0.01
0.63
0.44
0.83
0.31
0.39
0.19
8.04
7.76
7.54
0.13
0.12
0.12
46.50
46.97
45.39
0.61
0.68
0.74
OI1-u av.
g
47
44.10
0.02
0.65
0.29
7.87
0.13
46.29
0.65
rR
ee
Rock type
N
SiO2
TiO2
Al2O3
Cr2O3
FeO*
MnO
MgO
CaO
45.20
0.08
3.24
0.45
7.62
0.14
40.26
3.01
MORP averages
MARd
Ind. Oc.e Ind.Oc.f
158
48
44.83
46.19
44.26
0.02
0.04
0.03
1.45
2.13
1.88
0.40
0.28
0.51
8.47
8.61
8.18
0.13
0.13
0.13
44.15
40.83
43.62
0.55
1.79
1.40
rP
Fo
HP: Hawaiian pyrolite; Tin: Tinaquillo lherzolite, Exp Res: Final residue obtained after
“second-stage” melting in two-stage melting experiments (Green et al., 2004). MORP: MOR
peridotite; MAR: average of mid-Atlantic ridge peridotites sampled at ODP 153-920; Ind.Oc.: Indian
Ocean; Can.Isl.: Canary Islands; Cape V.: Cape Verde; Kerguel.: Kerguelen.
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References: aRingwood (1966); bJaques & Green (1980); cGreen et al. (2004), destimated on
the basis of data from Cannat et al. (1995), Casey (1997), Stephens (1997), Brandon et al.
(2000); eSnow & Dick (1995); fSeyler et al. (2003);gbased on data from Neumann (1991),
Neumann et al. (1995), Wulff-Pedersen et al. (1996), Abu El-Rus et al. (2006); hBonadiman et
al. (2005); iGrégoire et al. (1997, 2000a, 2000b), and Delpech (2004).
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Manuscript submitted to Journal of Petrology
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135x216mm (600 x 600 DPI)
http://www.petrology.oupjournals.org/