JOURNAL OF PETROLOGY
VOLUME 49
NUMBER 6
PAGES 1223^1251
2008
doi:10.1093/petrology/egn023
Ultra-refractory Domains in the Oceanic
Mantle Lithosphere Sampled as Mantle
Xenoliths at Ocean Islands
NINA S. C. SIMON1*, ELSE-RAGNHILD NEUMANN1,
COSTANZA BONADIMAN2, MASSIMO COLTORTI2,
GUILLAUME DELPECH3, MICHEL GRE¤GOIRE4 AND
ELISABETH WIDOM5
1
PHYSICS OF GEOLOGICAL PROCESSES, UNIVERSITY OF OSLO, PO BOX 1048, BLINDERN, N-0316 OSLO, NORWAY
2
EARTH SCIENCE DEPARTMENT, UNIVERSITY OF FERRARA, VIA SARAGAT, 1, 44100 FERRARA, ITALY
UMR CNRS 8148 IDES, ‘INTERACTIONS ET DYNAMIQUE DES ENVIRONNEMENTS DE SURFACE’, FACULTE¤ DES
SCIENCES, UNIVERSITE¤ ORSAY-PARIS SUD, BA“T. 504, 91405 ORSAY CEDEX, FRANCE
3
4
OBSERVATOIRE MIDI-PYRE¤NE¤ES, UMR 5562-DTP, OMP, UNIVERSITE¤ TOULOUSE III, AVENUE E. BELIN,
31400 TOULOUSE, FRANCE
5
DEPARTMENT OF GEOLOGY, MIAMI UNIVERSITY, 114 SHIDELER HALL, OXFORD, OH 45056, USA
RECEIVED JUNE 29, 2007; ACCEPTED APRIL 11, 2008
ADVANCE ACCESS PUBLICATION MAY 3, 2008
Many peridotite xenoliths sampled at ocean islands appear to have
strongly refractory major element and modal compositions. To better
constrain the chemistry, abundance and origin of these ultrarefractory rocks we compiled a large number of data for xenoliths
from nine 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.The xenolith suites from most ocean islands are dominated by ultra-refractory
harzburgites. Exceptions are the Hawaii and Tahiti peridotites,
which are more fertile and contain primary clinopyroxene, and the
Cape Verde suite, which contains both ultra-refractory and more fertile xenoliths. Ultra-refractory harzburgites are characterized by
the absence of primary clinopyroxene, low whole-rock Al2O3 , CaO,
FeO/MgO and heavy rare earth element (HREE) concentrations,
low Al2O3 in orthopyroxene (generally53 wt %), high Cr-number
in spinel (03^08) and high forsterite contents in olivine
(averages4915). They are therefore on average significantly more
refractory than peridotites dredged and drilled from mid-ocean
ridges and fracture zones. Moreover, their compositions resemble
those of oceanic forearc peridotites. The formation of ultra-refractory
ocean island harzburgites requires potential temperatures above those
*Corresponding author. Telephone: þ47-22856922. Fax: þ47-22855101.
E-mail: [email protected]
normally observed at modern mid-ocean 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 the long-term depleted Os and Sr isotope compositions 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 plume. As a result of their whole-rock compositions, ultrarefractory harzburgites have a very high solidus temperature at a given
pressure, 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.
mantle xenoliths; ocean islands; recycled oceanic mantle
lithosphere; ultra-depleted mantle
KEY WORDS:
ß 2008 The Author(s)
This is an Open Access article distributed under the terms
of the Creative Commons Attribution Non-Commercial License
(http://creativecommons.org/licenses/by-nc/2.0/uk/) which permits
unrestricted non-commercial use, distribution, and reproduction in
any medium, provided the original work is properly cited.
JOURNAL OF PETROLOGY
VOLUME 49
I N T RO D U C T I O N
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 to
develop realistic geodynamic models 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
peridotites sampled along mid-ocean ridges and fractures
zones (MORP: mid-ocean ridge peridotites). The xenoliths
in a second group of islands are relatively fertile, whereas
in a 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 with ocean island formation
and/or melt^rock interaction during ascent, and many of
the more fertile peridotites have strongly refractory pre-OI
protoliths (e.g. Ryabchikov et al., 1995; Wulff-Pedersen
et al., 1996; Gre¤goire et al., 1997, 2000a, 2000b; Moine et al.,
2000; Neumann et al., 2002, 2004; Delpech, 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, to use compositional variations in geodynamic
modeling, whole-rock major element compositions and
modal mineralogy averaged over some larger area are
actually the most important variables. This study 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 study is (1) to establish the distribution of
strongly refractory peridotites in ocean islands, (2) to constrain their origin and mode of formation, and (3) to discuss the implications of strongly refractory mantle material
for geodynamics with respect to physical and chemical
properties.
NUMBER 6
JUNE 2008
We have compiled available (published and unpublished) major element whole-rock and mineral composition data for 294 harzburgite and lherzolite xenoliths
and their minerals from the Canary Islands (Neumann,
1991; Neumann et al., 1995, 2002, 2004; Wulff-Pedersen
et al., 1996; Abu El-Rus et al., 2006), Madeira (Munha
et al., 1990), the Azores (Johansen, 1996), Cape Verde
(Ryabchikov et al., 1995; Bonadiman et al., 2005; Shaw
et al., 2006), Kerguelen (Gre¤goire et al., 1997, 2000a, 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 an Electronic Appendix (available for downloading at http://www.petrology.oxfordjournals.org). The
database also gives the major element compositions of
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
incompatible trace elements. In addition to spinel harzburgites and lherzolites, mantle xenoliths from OI include
dunites, 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, 1998; Gre¤goire et al., 1997, 2000b;
Neumann et al., 2004). However, data on some of those
samples are included in the database.
C O M P O S I T I O N A L VA R I AT I O N S
Classification of samples
Because the focus of this study is on pre-ocean-island
(pre-OI) compositions we apply 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 mild
enrichment in strongly incompatible trace elements and
resetting of radiogenic isotopes, to extensive petrographic
changes such as those involving the 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
1224
SIMON et al.
ULTRA-DEPLETED OCEANIC LITHOSPHERE
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 polars). (b) Poikilitic opx in spinel lherzolite
(Canary Islands) enclosing corroded remnants of larger olivine grains and small oval olivine grains (crossed polars). (c) Poikilitic
harzburgite (Kerguelen) with poikilitic cpx with spongy rims produced by interaction between protogranular harzburgites and alkaline,
silica-undersaturated melt (crossed polars). (d) Poikilitic lherzolite (Canary Islands) showing cpx, locally with spinel exsolution lamellae,
enclosing oval opx grains and spinel (crossed polars). Spinel occurs as single grains and as clusters or chains of very small grains.
(e) Symplectitic intergrowth between cpx and spinel along grain boundaries in spinel harzburgite (Canary Islands). Small olivine domains
(orange and bluish green) within the symplectite are in optical continuity with the neighboring olivine porphyroclasts (crossed polars). The
texture suggests formation of the symplectite and cpx by reaction between melt and peridotite. (f) Spinel lherzolite from Cape Verde showing
sieve-textured opx.
in Table 1) are classified as OI1 (Fig. 2). These xenoliths
typically show 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 þ
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. Because this study 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
samples are of no interest here, and are not discussed. We
therefore make no effort to distinguish between in situ
metasomatism in the mantle and interaction with the host
magma during ascent. Data points in the figures are
restricted to OI1 xenoliths. However, as it is of interest
to the discussion to know the extent to which melt^rock
1225
JOURNAL OF PETROLOGY
VOLUME 49
Table 1: Summary of textural and geochemical characteristics of the peridotite groups
Ultra-refractory (OI-u)
Fertile (OI-f)
OI1-u
OI1-f
porphyroclastic to protogranular
textures as in OI1-u (Fig. 2)
no petrographic evidence of
no petrographic evidence of
melt–rock interaction (Fig. 2)
melt–rock interaction (Fig. 2)
deformed porphyroclasts of ol þ opx
deformed porphyroclasts of
matrix: less deformed or
matrix: less deformed or
ol þ opx þ cpx sp
undeformed neoblasts of
ol þ opx sp cpx
undeformed neoblasts of
ol þ opx sp cpx
53 vol. % cpx
up to 20 vol. % cpx
primary cpx appears absent
primary clinopyroxene common
high WR MgO: 44–48 wt %
low WR MgO: 38–45 wt %
low WR CaO: 04–12 wt %
high WR CaO: 09–46 wt %
low WR Al2O3: 02–13 wt %
high WR Al2O3: 06–37 wt %
(Al2O3)opx: 04–43 wt %; main
higher (Al2O3)opx: 17–60 wt %;
range 1–3 wt %
high Cr-no.sp: 03–09; main
range 04–07
narrow range in high Fool: 897–926
main range 2–6 wt %
low Cr-no.sp: 01–05; main
range 01–04
wide range in Fool: 836–920
low HREEopx YbN: 004–04
low HREEopx YbN: 02–04
LREE-depleted to U-shaped
LREE-depleted REEopx patterns
REEopx patterns
OI2-u
OI2-f
porphyroclastic/protogranular/poikilitic
petrographic evidence of melt–rock
petrographic evidence of
reactions, modal
melt–rock reactions, modal
metasomatism, etc. (Fig. 1)
metasomatism, etc. (Fig. 1)
primary cpx appears to be absent
lower WR MgO: 39–48 wt %
low WR MgO: 37–45 wt %
wide range in WR CaO: 03–36 wt %
high WR CaO: 08–39 wt %
wide range in WR Al2O3: 03–37 wt %
high WR Al2O3
wide range in (Al2O3)opx: 01–40 wt %
high (Al2O3)opx: 29–60 wt %
wide range in Cr-no.sp: 02–09
low Cr-no.sp: 003–03
wide range to low Fool: 846–915
wide range to low
common: 06–39 wt %
Fool: 836–917
higher HREEopx YbN: 01–2
LREE-depleted to U-shaped
REEopx patterns
WR, whole-rock.
interaction has changed pre-OI peridotite compositions, we
show the ranges covered by OI2 xenoliths in some diagrams.
In addition to the petrographic distinction into OI1
and OI2, the xenoliths from different islands fall in two
NUMBER 6
JUNE 2008
categories with respect to major element compositions
(Table 1). One category comprises ultra-refractory harzburgites (OI-u) with low contents of clinopyroxene
(53 vol. %) 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
study is on the OI-u xenoliths.
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 temperature and related physical properties. Most
samples in the OI1-u group show some evidence of cryptic
metasomatism [i.e. U-shaped rare earth element (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-number [cation proportion Cr/(Cr þAl)] and mg-number [cation proportion
Mg/(Mg þ Fe2þ)] in spinel, and the 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) for the various peridotite
suites. OI1-u harzburgites lack primary clinopyroxenes,
and secondary clinopyroxenes in these rocks are commonly too small to analyze. OI2 xenoliths 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
geothermometer of Sachtleben & Seck (1981), assuming
that the presence of clinopyroxene 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
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
1226
SIMON et al.
ULTRA-DEPLETED OCEANIC LITHOSPHERE
Fig. 2. Photomicrographs of typical OI1 textures. (a) Protogranular texture in OI1-u xenolith (Kerguelen) showing olivine with undulatory
extinction and exsolved opx (crossed polars). (b) Details of protogranular texture in spinel harzburgite (Cape Verde; crossed polars). (c) Opx
porphyroclasts with cpx exsolution lamellae (exs opx) in spinel harzburgite (Canary Islands; crossed polars). 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. (plane-polarized light.). Abbreviations as in Fig. 1.
(Neumann, 1991; Siena et al., 1991; Neumann et al., 1995,
2002, 2004; Wulff-Pedersen et al., 1996; Abu El-Rus et al.,
2006; E.-R. Neumann, unpublished data, 2007). The OI1-u
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 53 vol. %) of 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, interstitial, vermicular cpx^sp intergrowths may
represent melt infiltration.
OI2-u xenoliths (mainly from La Palma and Tenerife)
comprise both spinel harzburgites and lherzolites.
Clinopyroxene in OI2-u is found as small interstitial
grains, as reaction rims on orthopyroxene (Fig. 1a and e),
in cpx þ sp opx symplectites along grain boundaries
and/or as part of reaction textures (formation of
clinopyroxene olivine at the expense of orthopyroxene;
Fig. 1a and 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 and e;
Neumann et al., 1995, 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 Fig. 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 result of minor metasomatism and/or melt^rock interaction. The wide range
in Na2O is probably due to enrichment processes or
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JOURNAL OF PETROLOGY
VOLUME 49
Table 2: Equilibration temperatures estimated for ocean
island xenoliths using the Sachtleben & Seck (1981)
geothermometer based on CaO in orthopyroxene
Island
Group
n
Range (8C)
Canary Islands
OI1-u
48
873–1209
1024
OI2-u
41
862–1215
1068
OI1-u
7
891–1017
936
OI2-u
5
892–1182
1037
Azores
OI2-u
17
1003–1232
1135
Cape Verde
OI1-u
1
OI1-f
4
1047–1233
1190
Madeira
Average (8C)
OI2-f
2
1028–1227
1127
OI1-u
16
933–1104
1014
OI2-u
25
960–1193
1062
Grande Comore
OI2-u
8
965–1239
1109
Samoa
OI1-u
6
983–1142
1063
OI2-u
13
1008–1228
1072
Hawaii islands
OI1-f
9
914–1205
1004
OI2-f
23
881–1161
1005
Loihi seamount
OI1-u
4
1139–1182
1163
Tahiti
OI1-f
19
996–1108
1058
OI2-f
1
JUNE 2008
in LREE relative to the middle REE (MREE) is most probably due to contamination (Neumann et al., 2002).
In contrast to the LREE-depleted opx1 in the OI1-u
xenoliths, cpx2 in the same rocks shows a wide range in
REE patterns and concentrations, ranging from LREEdepleted in the clinopyroxenes with the lowest HREE,
through U-shaped and flat patterns, to LREE-enriched
patterns in clinopyroxenes with the highest HREE
(YbN ¼ 054^11, LaN ¼ 022^128; Fig. 6).
Madeira
Mantle xenoliths from Madeira comprise refractory
deformed harzburgites and lherzolites (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 porphyroclastic
with deformed olivine and exsolved orthopyroxene porphyroclasts in a fine-grained matrix of ol þ opx þ sp cpx;
minor amounts of phlogopite may be present interstitially
or as inclusions in clinopyroxene (in lherzolite).
OI1-u xenoliths from Madeira (Munha et al., 1990) have
lower (Al2O3)opx (15^3 6 wt %) than orthopyroxene in
MORP, but fall within the depleted range of MORP with
respect to spinel compositions (cr-number: 026^055) and
Fool (904^912; Fig. 4). Temperature estimates for the OI1-u
xenoliths are 891^10178C with a mean of 9368C (Table 2);
OI2-u xenolith give the wider range of 892^11828C, with a
mean of 10378C.
1163
Kerguelen
NUMBER 6
1093
contamination (Neumann, 1991; Siena et al., 1991;
Neumann et al., 1995, 2004). The strongly melt-depleted
nature of the OI1-u harzburgites is also reflected in low
Al2O3 in orthopyroxene (03^34 wt %), Cr-rich spinels
(cr-number ¼ 029^095, with the majority of the samples
in the range 053^071), and relatively high Fo-content
[100Mg/(Mg þ Fe)] in olivine (907^925; Fig. 4).
Many OI2-u xenoliths also plot close to the opx^ol
tie-line (Fig. 3), although the majority have lower SiO2
and MgO, and higher CaO^FeO contents than the
OI1-u xenoliths and show wide ranges in TiO2 and Na2O.
Compared with the OI1-u rocks, the OI2-u rocks show a
somewhat wider range in (Al2O3)opx, a similar range in
cr-numbersp, but tend towards lower mg-number (Fig. 4).
Estimated equilibration temperatures for OI1-u xenoliths
are 873^12098C with a mean of 10248C (Table 2). The
OI2-u xenoliths cover a similar range (862^12158C) but
have a significantly higher average, 10688C.
Cores of OI1-u orthopyroxene porphyroclasts (opx1)
are strongly depleted in REE (YbN ¼ 009^041,
LaN ¼ 0002^12) and in light relative to heavy REE
(LREE and HREE, respectively; Neumann et al., 1995,
2004; Fig. 5), which is typical for mantle rocks that represent
residues after extensive partial melting (e.g. Johnson et al.,
1990; Neumann et al., 2004). A tendency for enrichment
The Azores
Mantle xenoliths from the Azores comprise Mg-Cr-rich,
porphyroclastic spinel harzburgites (52 vol. % clinopyroxene) and lherzolites (15 vol. % clinopyroxene) as well as
granular Fe^Ti-rich spinel wehrlites, clinopyroxenites and
dunites (Johansen, 1996; E. R. Neumann, unpublished
data, 2007). Harzburgites predominate over lherzolites.
The oldest generation of minerals in the harzburgites and
lherzolites is 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 (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). Clinopyroxene occurs partly as small
interstitial grains, partly as poikilitic grains. All Azores
xenoliths are classified as OI2.
The Azores xenoliths are dominated by fertile wholerock major element compositions, but are depleted with
respect to (Al2O3)opx (21^23 wt %) and cr-numbersp
(051^072; Fig. 4). Fool falls within the range 900^912
in all samples, with the exception of one with Fool of 865
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SIMON et al.
ULTRA-DEPLETED OCEANIC LITHOSPHERE
Fig. 3. OI1-u whole-rock major element variations. Data sources are as follows. Canary Islands: Neumann (1991), Wulff-Pedersen et al. (1996),
Neumann et al. (2002, 2004) and Abu El-Rus et al. (2006); Kerguelen: Gre¤goire et al. (1997, 2000a, 2000b) and Delpech (2004); Cape Verde:
Bonadiman et al. (2005) and C. Bonadiman (unpublished data, 2008). Dotted lines enclose the field of OI2-u xenoliths. The olivine (ol) and orthopyroxene (opx) fields and ol^opx tie-line (dashed grey 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, lherzolite) and residual compositions (per cent 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 partial melting of a residual chromite-bearing harzburgite at 15 GPa and 1320^13508C). (See text for discussion.)
(Fig. 4). Equilibration temperatures are high (1003^12328C,
mean: 11358C; 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 that in general has not had time to
affect the cores of porphyroclasts and large spinels.
Cape Verde
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
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Fig. 4. Relationships between Fo contents and Al2O3 in the cores of
coexisting olivine and orthopyroxene porphyroclasts, respectively,
and cr-number^mg-number in the cores of equidimensional spinel
grains in ultra-refractory OI1-u (upper panels) and more fertile OI1-f
(lower panels). The ranges covered by OI2 from the various islands
are indicated by colored dotted lines. Data sources are as follows.
Canary Islands: Neumann (1991), Neumann et al. (1995, 2002),
Wulff-Pedersen et al. (1996) and Abu El-Rus et al. (2006); Madeira:
Munha et al. (1990); Azores: Johansen (1996); Cape Verde:
Ryabchikov et al. (1995), Bonadiman et al. (2005) and Shaw et al.
(2006); Kerguelen: Gre¤goire et al. (1997) and Delpech (2004);
Samoa: Hauri & Hart (1994); Hawaii: Goto & Yokoyama (1988),
Sen (1988) and Bizimis et al. (2007); Loihi seamount: Clague (1988);
Tahiti: Tracy (1980) and Qi et al. (1994). Grey fields: compositional
ranges of minerals in MORP (data sources: Hamlyn & Bonatti,
1980; Komor et 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, 2006; Hellebrand
& Snow, 2003; Seyler et al., 2003). Because the ol^opx partition
coefficient for Mg/(Mg þ Fe) is close to unity for MORP (Seyler
et al., 2003) we have used [Mg/(Mg þ Fe)]opx instead of Fool where
olivine data are not available for the MORP. Stars marked
‘exp’ indicate the composition of minerals in the experimental
‘second-stage’ residue of Green et al. (2004). (See text for discussion.)
clinopyroxene (3 % cpx) as well as mildly depleted spinel
lherzolites with primary clinopyroxene (8^18 vol. % cpx).
Many samples show extensive evidence of reaction (e.g.
glass-bearing sieve-textured rims on clinopyroxenes and
spinels) and rims of ol þ sp þ cpx þ glass on orthopyroxene.
NUMBER 6
JUNE 2008
Fig. 5. Rare earth element patterns (normalized to Primordial
Mantle; McDonough & Sun, 1995) for orthopyroxene porphyroclasts
in (a) ultra-refractory OI1-u. Data sources: Canary Islands: Neumann
et al. (2002, 2004); Cape Verde: Bonadiman et al. (2005) and
C. Bonadiman (unpublished data, 2008); Kerguelen: Gre¤goire et al.
(2000b) and Delpech (2004); (b) fertile OI1-f from Cape Verde
(Bonadiman et al., 2005). Shaded field: orthopyroxene porphyroclasts
in MORP (data sources: Tartarotti et al., 2002; Brunelli et al., 2003;
Hellebrand et al., 2005).
Ryabchikov et al. (1995) also reported phlogopite in some
samples. Glass is common, both interstitially and as small
inclusions in olivine and 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 (Fig. 2b, d and Fig. 1f, respectively). The presence of
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) define a trend of strongly decreasing clinopyroxene
with increasing olivine, whereas the harzburgites (OI-u)
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SIMON et al.
ULTRA-DEPLETED OCEANIC LITHOSPHERE
Fig. 6. REE patterns for clinopyroxenes in OI1. (a) OI1-u from the Canary Islands (Neumann et al., 2002, 2004), Cape Verde (Bonadiman
et al., 2005; C. Bonadiman, unpublished data, 2008), Kerguelen (Mattielli et al., 1999; Gre¤goire et al., 2000a, 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).
Clinopyroxene in OI1-u shows LREE-depleted to LREEenriched patterns (YbN ¼ 07^45, LaN ¼ 03^122; Fig. 6).
The Cape Verde OI1-f lherzolites are relatively fertile
with 397^455 wt % MgO, 12^46 wt % CaO and
05^20 wt % Al2O3 (Fig. 8); cr-numbersp is 03,
Fool ¼ 87^89 and (Al2O3)opx ¼ 41^60 wt % (Fig. 4).
Orthopyroxene porphyroclasts in OI1-f are depleted in
HREE and have moderately LREE-depleted patterns
(YbN ¼ 023^036, LaN ¼ 003^05; Fig. 5), whereas OI1-f
clinopyroxene have upward-convex REE patterns
(YbN ¼ 09^15, LaN ¼ 17^2 0; Fig. 6).
Kerguelen
Fig. 7. Relationships between modal proportions of clinopyroxene
and olivine in mantle xenoliths from Cape Verde (data sources:
Bonadiman et al., 2005; C. Bonadiman, unpublished data, 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.
show roughly constant, low clinopyroxene contents.
Equilibration temperatures are given inTable 2.
The OI-u harzburgites are strongly refractory with
whole-rock major element contents falling close to the
opx^ol tie-line (e.g. 449^479 wt % MgO, 056^11wt %
CaO, 508 wt % Al2O3; Fig. 3). Somewhat higher TiO2 in
a few samples is probably due to minor contamination.
Mineral major element data are available only for
one OI1-u harzburgite, which has Cr-rich spinel
(cr-number ¼ 055), Fo-rich (913) olivine and Al2O3-poor
orthopyroxene (26 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 ¼ 004^029, LaN ¼ 002^035; Fig. 5).
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 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; Gre¤goire et al., 1997, 2000a, 2000b; Moine et al., 2000;
Delpech, 2004). In OI1-u, 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 (18^2 4 wt %), high cr-numbersp (042^061) and
high, uniform Fool of 910^918 (Fig. 4). Minerals in OI2u tend towards more fertile compositions. OI1-u xenoliths
give lower equilibration temperatures (933^11048C,
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Fig. 8. Major element variations in OI1-f from Cape Verde (Bonadiman et al., 2005; C. 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 al., 1994). Opx and ol
fields and opx^ol tie-lines (grey dashed lines) as in Fig. 3. Blue and red symbols and arrows show the results of mantle melting experiments by
Green et al. (2004). Light grey field, OI1-u.
average 10148C; Table 2) than OI2-u (960^11938C, average
10628C).
Pyroxene REE patterns indicate minor interaction with
REE-enriched melts or fluids (e.g. Mattielli et al., 1996,
1999; Gre¤goire et al., 1997, 2000a, 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 ¼ 003^023, LaN ¼ 0011^093;
Fig. 5). Clinopyroxenes in OI1-u xenoliths show a range in
REE patterns from LREE-depleted, through U-shaped,
to LREE-enriched, with YbN ¼ 013^097 and LaN ¼
0013^21 (Fig. 6).
Grande Comore
Mantle xenoliths from Grande Comore comprise spinel
lherzolites, dunites and wehrlites; no harzburgites have
been reported (Coltorti et al., 1999). The lherzolites have
protogranular to porphyroclastic textures, and consist
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SIMON et al.
ULTRA-DEPLETED OCEANIC LITHOSPHERE
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.
Despite the reaction textures, the Grande Comore
OI2 xenoliths have low (Al2O3)opx (08^34 wt %), high
cr-numbersp (04^07; Fig. 4) and moderate Fool
(900^913), classifying them as OI2-u. Estimated equilibration temperatures fall in the range 965^11398C, with a
high mean temperature of 11098C (Table 2).
Samoa
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 porphyroclastic rocks.
Clinopyroxene occurs mostly along the edges of orthopyroxene þ spinel symplectites, but is also found as small
interstitial grains rimming orthopyroxene and olivine.
In some rocks, melt^rock interaction led to interstitial
glass þ clinopyroxene apatite-bearing patches in various
stages of reaction with adjacent porphyroclasts (Hauri
et al., 1993). Three harzburgites with porphyroclastic textures (05% cpx) have been classified as OI1-u.
Whole-rock major element data are available only for
OI2-u harzburgites; these fall close to the opx^ol tie-line
(not shown). Mineral compositions are typical of OI1-u
(13^2 6 wt % (Al2O3)opx; cr-numbersp 049^070;
Fool 905^914; 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 (983^11428C with average 10638C and
1045^10978C with average10728C, respectively; Table 2).
Samoan OI1-u clinopyroxenes show a wide range in
REE contents (YbN ¼ 0045^86, LaN ¼ 0011^25; 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.
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). 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, 1988; Goto & Yokoyama, 1988; Sen & Leeman,
1991; Bizimis et al., 2003). Primary clinopyroxene porphyroclasts have spinel or orthopyroxene exsolution lamellae
(e.g. Sen, 1988).
The whole-rock major element compositions of all the
Hawaiian harzburgites and lherzolites plot far away
from the opx^ol tie-lines in Fig. 8; Hawaiian OI1 xenoliths
include some of the lowest MgO (37^43 wt %), and highest
CaO (11^39 wt %), TiO2 (043 wt %), Al2O3 (4 wt %)
and Na2O (07 wt %) contents reported from ocean
islands. OI1-f mineral compositions are also relatively
fertile [17^49 wt % (Al2O3)opx, cr-numbersp 009^055,
Fool 885^920; Fig. 4]. There is a strong tendency for lower
Fool in the OI2-f compared with the OI1-f. Equilibration
temperatures are similar for OI1-f (914^12058C with a mean
of 10048C) and OI2-f (881^11618C, mean10058C; Table 2).
A large amount of clinopyroxene REE data from
Hawaiian peridotites has been 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
depletion (YbN ¼ 34^47, LaN ¼ 002^10; 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 (15^20 wt %), high cr-numbersp (053^075;
Fig. 4) and high Fool (906^923; Fig. 4). Estimated temperatures for the OI1-u are even higher than in Hawaiian
OI1-f, 1139^11828C, with a mean of 11638C (Table 2).
Tahiti
Tahiti mantle xenoliths are porphyroclastic to equigranular, mildly refractory OI-f spinel lherzolites (most abundant) and harzburgites (Tracy, 1980; Qi et al., 1994).
Clinopyroxene occurs partly as large porphyroclasts,
partly in vermicular intergrowths with spinel and olivine.
Some clinopyroxenes have spongy glass-bearing rims,
interpreted as being a result of metasomatism (Qi et al.,
1994). No other evidence of metasomatism has been given
by Tracy (1980) or Qi et al. (1994). Samples with spongy
glassy rims have been classified as OI2-f, and the rest of
the Tahiti samples as OI1-f.
Like 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
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Na2O contents (012 and 005 wt %, respectively; Fig. 3).
They also have relatively fertile mineral compositions
with (Al2O3)opx 23^56 wt %, cr-numbersp 012^041 and
Fool 846^907 (Fig. 4). OI1-f equilibration temperatures
are 996^11088C with a mean of 10588C (10938C for the
single OI2-f sample; Table 2). Clinopyroxenes in OI1-f
have uniform, mildly LREE-enriched REE patterns
(YbN ¼ 22^40, LaN ¼16^3 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 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 no data on OI1 xenoliths are available for the
Azores and Grande Comore, the low (Al2O3)opx, high
cr-numbersp and high Fool in many OI2 samples suggest
that ultra-refractory OI1-u domains may also be present
in the upper mantle beneath these islands. Below we summarize the characteristics of OI-u and OI-f xenoliths, compare the OI-u with MORP, and speculate on the origin
and evolution of OI-u. Furthermore, we provide some estimates of the 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 Fig. 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 makes up
52 vol. % and occurs as secondary neoblasts (Figs 1 and
2). OI1-u have high cr-numbersp, 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 residues formed in two-stage melting experiments (Green
et al., 2004; Fig. 3) in which a residual, dry, chromitebearing harzburgite with a composition close to MORP
was melted by an additional 10% under H2O^CO2-fluxed
conditions at 15 GPa and 1320^13508C.
The secondary origin of clinopyroxene in OI1-u and
OI2-u xenoliths from the Canary Islands is reflected in
their REE characteristics. Two OI1-u samples with mildly
Fig. 9. Opx/cpx REE ratios for selected samples from the Canary
Islands, compared with published partition coefficients for pyroxenes
in equilibrium in spinel harzburgites and lherzolites (Eggins et al.,
1998; Garrido et al., 2000; Green et al., 2000; Hellebrand et al., 2005).
Continuous lines, OI1-u; dashed lines, OI2-u. The opx/cpx ratios for
OI1 and some OI2 fall within the range of the experimental data,
indicating equilibrium between the two pyroxenes. The ratios
obtained for the most REE-rich OI2 clinopyroxenes fall far below
the experimental range, indicating REE enrichment in clinopyroxene
relative to orthopyroxene.
LREE-depleted cpx patterns have opx/cpx REE ratios
close to the range of published equilibrium opx/cpx REE
ratios in spinel peridotites (Fig. 9), supporting petrographic
evidence (Fig. 2) that these clinopyroxenes (cpx2) formed
through exsolution from orthopyroxene porphyroclasts
(opx1), followed by recrystallization. Many OI2-u, in contrast, have opx/cpx REE ratios that fall significantly below
published REE partition coefficients (Fig. 9), implying disequilibrium between the pyroxenes. We interpret these
low ratios as the result of a much stronger influence of
LREE-enriched fluids or melts on clinopyroxene than on
orthopyroxene; some of the most LREE-enriched 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,
peridotite massifs, ophiolites and mantle xenoliths (e.g.
Hellebrand et al., 2002; Gre¤goire et al., 2003; Mu«ntener &
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)
The OI-f xenoliths (Hawaii, Tahiti, the lherzolite suite
from Cape Verde) are significantly more fertile than the
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SIMON et al.
ULTRA-DEPLETED OCEANIC LITHOSPHERE
OI1-u series [lower whole-rock MgO and higher CaO^
FeO^TiO2^Al2O3 contents, higher (Al2O3)opx, lower
cr-numbersp and lower Fool, common presence of primary
clinopyroxene; Figs 4, 6 and 8; Table 1]. A primary origin
of clinopyroxenes in many xenoliths from Hawaii is supported by their strongly LREE-depleted REE patterns
(Yang et al., 1998; Fig. 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 group of spinel peridotites
from Oahu. The OI1-f xenolith suites from Hawaii, Tahiti
and 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) make up a large part of the available oceanic
peridotite database and 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 Fig. 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 (542 wt %) are in
fact dunites, even though they are classified as harzburgites
in the literature. The samples that Seyler & Bonatti (1997)
and Tartarotti et al. (2002) described as strongly infiltrated
by melts are excluded from Fig. 10.
MORP show wide ranges in whole-rock and
mineral major element compositions. Most samples have
35^49 wt % MgO, 40^49 wt % SiO2, 501^45 wt %
CaO and 03^45 wt % Al2O3 (Fig. 10). (Al2O3)opx falls
in the range 17^61wt %, cr-numbersp is 012^057, and in
most 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 open-system melt migration and crystallization, melt^rock interaction by reactive porous flow,
metamorphism, serpentinization and marine weathering
(e.g. Elthon, 1992; He¤kinian et al., 1993; Cannat et al., 1997;
Casey, 1997; Niu, 1997, 2004; Niu & He¤kinian, 1997a;
Asimow, 1999; Baker & Beckett, 1999; Tartarotti et al.,
2002). The combination of these processes explains much
of the strong compositional heterogeneity observed in
MORP. The low CaO/Al2O3 ratios exhibited by many
MORP (ranging from about 10 to 5001; 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 10, 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 granular exsolution
from orthopyroxene, whereas proportions 42% most
probably formed from infiltrating melt. Other workers
(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 þ opx þ cpx. However, these researchers have acknowledged that some clinopyroxene has formed through melt
infiltration.
On average OI1-u harzburgites have higher MgO, lower
(Al2O3)opx and higher cr-numbersp 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 average, less depleted in REE than those
in OI1-u (e.g. YbN 017^14 in MORP; YbN 5005^041 in
OI1-u). Olivine in MORP has very uniform compositions
with average Fo contents between 896 (Central Indian
Ridge) and 907 (Mid-Atlantic Ridge; Fig. 12), whereas
most cores of olivine porphyroclasts in OI1-u have Fo
contents between 91 and 92 (e.g. Hierro, Lanzarote,
Kerguelen, Loihi). OI-u xenolith suites that include a significant proportion of 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 in some islands with heterogeneous
olivine compositions (e.g. Gran Canaria and Samoa)
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 905; that is, Fo at least as high as
or higher than average MORP. If the OI-f xenoliths from
Hawaii and Tahiti are excluded, 81% of 191 xenoliths have
Fo 905 (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 and 12) suggest that,
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Fig. 10. Variations in bulk-rock major element compositions of MORP from various ridge^transform systems (data sources: Miyashiro et al.,
1969; Aumento & Loubat, 1971; Dick, 1989; Blum, 1990; Cannat et al., 1995; Snow & Dick, 1995; Casey, 1997; Niu & He¤kinian, 1997a; Stephens,
1997; Brandon et al., 2000; Lee et al., 2003; Seyler et al., 2003; Niu, 2004; Harvey et al., 2006; Godard et al., 2008). Grey field, OI1-u xenoliths.
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 the East Pacific Rise and Mid-Atlantic
Ridge (Niu & He¤kinian, 1997a, 1997b). Niu & He¤kinian
(1997b) proposed that the extent of mantle melting
beneath normal mid-ocean ridges increases with increasing spreading rate. This may explain the presence of
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-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 fast
spreading.
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 and 11) suggest that
OI1-u may form by mid-ocean ridge processes although this
would require potential mantle temperatures of at least
14008C (Fig. 13). However, available Re^Os data (see
Table 3) on some OI1-u (and some ultra-refractory MORP)
are not compatible with a simple ridge-melting model.
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SIMON et al.
ULTRA-DEPLETED OCEANIC LITHOSPHERE
Fig. 11. Variations in CaO/Al2O3 ratios of (a) MORP and (b) OI1-u.
Data sources as in Figs 3 and 10. MORP show a strong tendency for
lower CaO/Al2O3 ratios than OI1-u, probably as a result of
serpentinization.
Alternative models for the formation of OI1-u are
discussed below.
The origin of ultra-refractory peridotites
If some (or all) OI1-u did not form by ridge processes,
what is the origin of these peridotites and how did they
acquire their highly refractory nature? Different sets of
data that throw light on these problems are discussed in
the following sections.
Old Re^Os isotope model ages in OI peridotites
There is increasing evidence that many peridotite xenoliths collected from ocean islands are older than the adjacent 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 xenoliths from the Kerguelen Islands (516 GPa;
Hassler & Shimizu, 1998; Delpech, 2004), a group of lherzolites (OI-f) from Sal, Cape Verde (332 015 Ga;
Bonadiman et al., 2005), OI1-u and OI2-u xenoliths from
the Canary Islands (650 Ma; Table 2), and in spinel lherzolite xenoliths (OI-f) from various localities in Oahu,
Hawaii (90 Ma to 2 Ga; Griselin & Lassiter, 2001; Bizimis
et al., 2004, 2007). Evidence of an old depleted mantle
component has also been found in Azores lavas (25 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 various 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 ultrarefractory segments of the upper mantle are resistant to
further melting (i.e. sterile).
The old-age peridotites sampled as ocean island xenoliths and along mid-ocean ridges may represent recycled
continental or oceanic lithosphere, or sub-arc mantle
wedge material. In neither case are they complementary
to local MORB, but represent material ‘exotic’ in the local
oceanic mantle lithosphere. Alternatively, the ultrarefractory 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 Osrich 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 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 midocean ridges. Hence, the formation of OI1-u requires
either significantly higher potential mantle temperatures
or the availability of a fluid flux during melting to lower
the solidus.
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-refractory peridotites may thus be present in the convecting mantle and may be incorporated
into 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 mantle plumes. However, the presence of old,
‘exotic’ OI1-u (and OI1-f) type material in the oceanic
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 & He¤kinian
(1997a, 1997b).
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Fig. 12. Fo contents of cores in olivine porphyroclasts in OI spinel peridotite xenoliths compared with those in MORP. Olivine porphyroclasts
in OI1-u (dry) xenoliths show peaks in 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 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 with average Fo
between 896 and 906. The dashed line shows the MORP average of 905. Data for OI peridotites 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), Gre¤goire et al. (1997), Coltorti et al. (1999),
Delpech (2004), Bonadiman et al. (2005) and Shaw et al. (2006). Data for 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) and Seyler et al. (2003). AAR, American^
Antarctic Ridge^transform system in the Indian Ocean; CIR, Central Indian Ridge^transform system; MAR, Mid-Atlantic Ridge; SWIR,
Southwest Indian Ridge^transform system.
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 boundary.
Because 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 island xenoliths last equilibrated.
Available data imply that OI-u xenoliths last equilibrated at relatively low pressures and temperatures.
Minimum pressures obtained from geothermometry combined with densities of CO2 inclusions in mantle xenoliths
from the Canary Islands, Samoa, Tahiti, Kerguelen and
Grand Comore fall in the range 07 to 125 GPa
(Hansteen et al., 1991, 1998; Schiano et al., 1992, 1994, 1998;
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SIMON et al.
ULTRA-DEPLETED OCEANIC LITHOSPHERE
Fig. 13. Pressure^temperature diagram showing the stability fields
of plagioclase-, spinel- and garnet-peridotites 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 end-members to the Holland and
Powell database (Holland & Powell, 1998; Connolly & Petrini, 2002;
Klemme et al., in preparation). Ultra-refractory OI1-u peridotite is
assumed to contain 005 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 & Wood (1998; increase of 78C 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 as
OI1-u have lost all cpx, 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 higher pressures for OI1-u. Also shown is the simplified
geotherm at a mid-oceanic ridge. The temperature at the Moho is
taken as 5508C and T at the base of the lithosphere (Tast) is 13308C.
Grey fields show the ranges of pressure and temperature estimates
for OI-u (Table 2).
Schiano & Clocchiatti, 1994; Neumann et al., 1995).
Gre¤goire et al. (2000a) estimated a minimum pressure of
formation of 13 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, followed by heating
and dehydration, and clearly originate in the shallower
parts of the mantle lithosphere, close to the Moho (Abu
El-Rus et al., 2006).
Based on a comparison between the compositions of
experimentally rehomogenized melt 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 c. 1GPa. The xenolith
series from Hawaii comprise both plagioclase^spinel and
spinel lherzolites and harzburgites; the highest pressures,
21GPa, 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 (480 Ma) oceanic lithosphere
(Fig. 13).
This conclusion, however, assumes that the oceanic
lithosphere is not deeply eroded by an underlying mantle
plume. If a significant portion of the original 60^100 km
thick lithosphere is eroded 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 has been proposed
for Hawaiian peridotites that give old Re^Os ages
Table 3: Bulk-rock Re^Os isotope data for Canary Island harzburgite and lherzolite xenoliths (E.-R. Neumann &
E. Widom, unpublished data)
Sample
Type
Os (ppb)
Re (ppb)
187
Os/188Os (2)
187
Re/188Os
Tma (Ma)
Trd (Ma)
LA8-10
OI1
2363
0067
012409(19)
0136
65452
43457
LA1-7
OI1
1251
0003
012613(20)
0010
13460
13130
LA1-13
OI1
7576
0003
012423(11)
0002
41575
41381
PAT2-68
OI1
1271
0036
012554(13)
0134
32910
21917
H1-5
OI2
3631
0002
012652(18)
0002
7358
7314
PAT2-75
OI2
1750
0030
012486(16)
0082
40209
32028
TF14-38
OI2
2189
0011
012303(27)
0025
63061
59155
TF14-47
OI2
1206
0012
012678(12)
0049
3913
3434
TF14-48
OI2
2273
0021
012486(32)
0045
36077
32087
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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JOURNAL OF PETROLOGY
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(Bizimis et al., 2007) and is supported by some 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, 860^12008C (Table 1). These temperatures
would require plume material to stagnate at shallow
depths for sufficiently long times to allow thermal and
chemical re-equilibration 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 happens only if the plume is
at least 1408C hotter than the background temperature.
Such very high 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
OI1-u were part of the rising plume and transported to
pressures below 13 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 temperatures, more
fertile plume components would definitely be partially
molten at 13 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 or 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 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.
Another argument against transport of OI-u peridotites
from the deep mantle is their low densities (Simon &
Podladchikov, 2008). Because the ultra-refractory OI-u
peridotites are 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 therefore unlikely to be present in large quantities in the source region of deep,
thermally 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 for xenoliths from the
Canary Islands. In general, Sr, Nd, Pb and Hf isotope
data for OI mantle xenoliths indicate isotopic signatures
similar to those of the local ocean island basalts (e.g.
Vance et al., 1989; Whitehouse & Neumann, 1995; Mattielli
et al., 1996, 1999; Sen et al., 2003; Neumann et al., 2004).
NUMBER 6
JUNE 2008
Fig. 14. Sr^Nd isotope data for whole-rock samples (Vance et al.,
1989; Rolfsen, 1994; Ovchinnikova et al., 1995) and clinopyroxene
mineral separates (Whitehouse & Neumann, 1995; Neumann et al.,
2002), compared with laser ablation microprobe analyses of 87Sr/86Sr
ratios in clinopyroxenes in OI1 and OI2 xenoliths from the Canary
Islands (Neumann et al., 2004). Fields of Canary Islands basalts (data
from Hoernle et al., 1991, 1995; Ovchinnikova et al., 1995; Thirlwall
et al., 1997; Simonsen et al., 2000), and Mid-Atlantic Ridge N-MORB
(Ito et al., 1987; Dosso et al., 1991) are shown for comparison. In situ Sr
isotope ratios in clinopyroxenes in two OI1-u samples (continuous
lines) show depleted Sr isotope ratios below those of the Canary
Islands basalts. Laser ablation 87Sr/86Sr of OI2-u cpx (dashed lines)
and a wehrlite cpx (dotted line) overlap with the Canary Island
basalt field. Hz-lz, harzburgite^lherzolite.
Sr^Nd^Pb isotope data on whole-rock samples and clinopyroxene separates from mantle xenoliths from the
Canary Islands follow this general pattern (Whitehouse
& Neumann, 1995; Neumann et al., 2004; Fig. 14).
However, in situ laser ablation analyses of 87Sr/86Sr in clinopyroxenes show a range in isotope ratios (Neumann et al.,
2004). Clinopyroxenes in two OI1-u xenoliths have
87
Sr/86Sr isotope ratios of 070266 and 070276, which are
significantly below those of Canary Islands basalts and
require 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. Only clinopyroxenes in OI2-u xenoliths gave ratios of 07031^07033, which,
like those of the whole-rock and clinopyroxene separates,
is within the range of Canary Islands basalts field. One
OI2-u sample gave an intermediate value of 070288.
The original old depleted isotopic signature has therefore
recently been overprinted by metasomatism by melts or
fluids originating in the Canary Islands plume (Neumann
et al., 2004).
In conclusion, we consider a recent plume origin for the
ocean island xenoliths highly unlikely. Based on the P^T
estimates, the inferred physical properties (high viscosity
and low density) and the in situ Sr isotope data, we favor
the idea that the OI peridotites were part of the oceanic
1240
SIMON et al.
ULTRA-DEPLETED OCEANIC LITHOSPHERE
lithosphere before the onset of the plume processes that
brought them to the surface.
Did the OI1 peridotites acquire their highly depleted
compositions by plume processes?
Although we concluded above that the OI-u peridotites are
not brought up by the associated plume, the lithospheric
mantle beneath ocean islands is clearly affected by plumerelated processes. The most obvious one is metasomatism
by plume-derived melts and fluids, which is reflected in
OI2 xenoliths, and which has also caused some enrichment
in incompatible elements and radiogenic isotopes in OI1
xenoliths (e.g. Figs 1^6 and 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 or 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).
Microtextural evidence of melt^wall-rock reactions is
common in OI2 peridotites. They show the reaction
(opx þ melt ! ol þ cpx). This is opposite to the reaction
described by Kelemen et al. (1992), but is compatible with
infiltration by transitional to strongly alkaline ocean
island melts saturated in clinopyroxene. Infiltration of
ocean island peridotites by alkaline melts typically leads
from harzburgite (OI1-u) to lherzolite (OI2-u) to dunite^
wehrlite. In addition, the Ti^Al^Fe-rich nature of the clinopyroxenes in some OI dunites, wehrlite and clinopyroxenites suggests crystallization directly from alkali basaltic
magmas (e.g. Neumann et al., 2004). Such reactions are by
definition absent from OI1 xenoliths. Aditionally, 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 of 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 possibility
highly unlikely. In Fig. 13, the estimated P^T conditions
for the ultra-refractory xenoliths are compared with
estimated solidi for peridotites of average OI1-u and more
fertile compositions. The equilibration temperatures
obtained for 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.
As neither plume-induced melting nor reaction with
plume melts appears to be responsible for the ultrarefractory 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 with the conclusions that the old Os model
ages obtained for some ultra-refractory OI peridotites and
MORP reflect ancient melting events (e.g. Hassler &
Shimizu, 1998; Brandon et al., 2000; Griselin & Lassiter,
2001; Bizimis et al., 2004, 2007; Delpech, 2004; Bonadiman
et al., 2005; Alard et al., 2006; Harvey et al., 2006).
Comparison of OI1-u with refractory peridotites from
other settings
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 of the lithosphere as it ages. By which kind of
process did OI1-u peridotites acquire their ultra-refractory
composition if they were not formed in their host plume?
We compare OI1-u compositions with those of subcontinental mantle xenoliths and peridotites from oceanic
sub-arc mantle (Figs 15 and 16) 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 degree of melt extraction has
been ascribed to either higher mantle temperatures or
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, 2007) and East Greenland (Bernstein et al., 1998).
However, these xenoliths differ significantly from the OI1u xenoliths by lower FeO, and higher SiO2 and MgO
(Fig. 15). Figure 15 also shows the comparison of OI1-u
xenoliths with some refractory off-craton continental
mantle xenoliths, 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
part of the scatter. We cannot therefore exclude the possibility that OI1-u represent fragments of continental lithospheric mantle that was eroded from the base of the
continental lithosphere or sheared off during rifting and
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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 Brazil (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 (Boyd & Mertzman, 1987; Cox et al., 1987; Nixon et al., 1987; Winterburn et al., 1990; Boyd et al., 1999; Simon et al., 2003, 2007).
The off-craton mantle xenolith series show wide compositional ranges, including fertile compositions.
passive margin formation. The latter scenario has been
proposed as the ‘raft’ model by Mu«ntener & Manatschal
(2006) to explain the occurrence of some strongly refractory spinel peridotites in the apparently oceanic domain
at the Newfoundland margin. Such a model, however,
does not explain how the presumed continental lithospheric mantle reached its high degree of depletion.
Mu«ntener & 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 Fig. 16 we compare OI1-u with some mantle
peridotite series from modern oceanic sub-arc settings.
As a group, the arc samples also show considerable
1242
SIMON et al.
ULTRA-DEPLETED OCEANIC LITHOSPHERE
Fig. 16. Comparison between OI1-u xenoliths (grey fields) and mantle peridotites from the Izu^Bonin^Mariana forearc, the Torishima Forearc
Seamount (data from Ishii et al., 1992; Parkinson & Pearce, 1998; Zanetti et al., 2006) and Lihir (Papua New Guinea; data from McInnes et al.,
2001). These forearc peridotites (individual symbols) have major element compositions similar to OI1-u. Also shown is the compositional
range covered by the majority (495%) of peridotites from active oceanic arcs (dotted 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).
compositional variations, but the peridotite suites from
the Izu^Bonin^Mariana forearc, the Torishima Forearc
Seamount (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 will be given elsewhere (Neumann &
Simon, in press).
In conclusion, the highly refractory OI1-u xenoliths
most probably represent residues of shallow melting processes at elevated temperatures or water contents, or both.
Following melting these residues were recycled within the
upper part of the convecting mantle and became accidentally trapped in the oceanic lithosphere before they were
sampled by the ocean island basalts. However, some may
have formed along hot mid-ocean ridges and been directly
incorporated in the oceanic mantle lithosphere.
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JOURNAL OF PETROLOGY
VOLUME 49
CONC LUSIONS A N D
I M P L I C AT I O N S F O R
GEODY NA M IC S
The major element compositions of oceanic island peridotites have been reviewed. After petrographically filtering
for the effects of mantle metasomatism and melt^rock
interaction during transport, a large proportion of the
peridotite xenoliths sampled in ocean islands proved to be
ultra-refractory harzburgites. These samples are indicative
of partial melting beyond the cpx-out reaction and are on
average significantly more refractory [higher Fool and
cr-numbersp, lower CaO and (Al2O3)opx] than abyssal
peridotites sampled by dredging and drilling along midocean ridges and at oceanic 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 associated ocean
island plume. These peridotites must therefore be residues
of large-degree melting events at temperatures significantly higher than present-day normal mantle potential
temperatures, or they are residues formed during extensive
fluid-fluxed melting, possibly in an arc setting. The last
possibility is supported by their compositional similarity
to modern oceanic forearc peridotites. If the melt depletion
of the OI1-u is caused by fluid-fluxed melting any textures
associated with fluid^rock interaction, which are commonly observed in recent arc-related peridotites, must
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 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 heterogeneous mantle, as supported
by recent geochemical and geophysical observations and
theoretical models (e.g. Phipps Morgan & Morgan, 1999;
Brandon et al., 2000; Helffrich & 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 consequences for geodynamics because of its distinct geochemical and physical
properties.
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
NUMBER 6
JUNE 2008
(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 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).
Because of their high Mg/Fe ratios and low Al contents,
the OI1-u harzburgites 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
(i.e. mid-ocean ridges) and may preferentially be incorporated into the cooling and growing oceanic lithosphere. We
thus expect the proportion of highly refractory harzburgite
to be significantly higher in the oceanic mantle lithosphere
and underlying asthenosphere than in the deeper parts
of the convecting mantle, in agreement with the model of
Phipps Morgan & Morgan (1999). An important consequence of a high proportion of low-density OI1-u material
in the oceanic lithosphere is that it will be difficult to subduct, particularly in the initial stages of subduction. Some
preliminary calculations with Perple_X [http://www.
perplex.ethz.ch/perplex.html; see Connolly & Petrini
(2002) for methods] show that a mantle column (100 km
thick) with the average composition of the Canary Island
OI1-u xenoliths overlain by a 65 km thick basaltic crust
(crust ¼ 2950 kg/m3) will have a density of 3200 kg/m3
close to the ridge, compared with 3220 kg/m3 for the average composition of 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 a 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 of fertile mantle, depending on
the melt model used. Afonso et al. (2007) concluded 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,
1244
SIMON et al.
ULTRA-DEPLETED OCEANIC LITHOSPHERE
Table 4: Peridotite major element compositions (recalculated to 100% without Na2O, K2O, P2O5) compared with
OI1 averages
Primitive mantle models
Rock type:
HP1
Tin2
Exp.
MORP averages
res.3
MAR4
n:
Ind.Oc.5
Ind.Oc.6
158
48
SiO2
4562
4520
4442
4483
4619
4426
TiO2
072
008
001
002
004
003
Al2O3
357
324
083
145
213
188
Cr2O3
043
045
025
040
028
051
FeO
855
762
570
847
861
818
MnO
014
014
013
013
013
MgO
3785
4026
4776
4415
4083
4362
CaO
311
301
102
055
179
140
OI1-u averages
Rock type:
Can.Isl.7
Cape V.8
Kerguelen9
OI1-u av.
n:
28
SiO2
4376
4362
12
47
4518
TiO2
002
4410
002
001
Al2O3
002
063
044
083
065
Cr2O3
031
039
019
029
FeO
804
776
754
787
MnO
013
012
012
013
MgO
4650
4697
4539
4629
CaO
061
068
074
065
7
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 Ocean Drilling Program (ODP) Leg 153, Site 920; Ind.Oc., Indian Ocean; Can.Isl., Canary Islands; Cape V.,
Cape Verde.
References: 1Ringwood (1966); 2Jaques & Green (1980); 3Green et al. (2004), 4estimated on the basis of data from Cannat
et al. (1995), Casey (1997), Stephens (1997) and Brandon et al. (2000); 5Snow & Dick (1995); 6Seyler et al. (2003); 7based
on data from Neumann (1991), Neumann et al. (1995), Wulff-Pedersen et al. (1996) and Abu El-Rus et al. (2006);
8
Bonadiman et al. (2005); 9Grégoire et al. (1997, 2000a, 2000b) and Delpech (2004).
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 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 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, seismic velocities)
with measured geophysical data (e.g. Afonso et al., 2007).
AC K N O W L E D G E M E N T S
Dante Canil generously gave us a copy of his database
on mantle rocks, which was the foundation of our own
database. The work has profited from discussions with a
number of colleagues, in particular Y. Podladchikov,
L. Ru«pke, J. Phipps Morgan, O. Alard, J. C. Afonso and
E. Hellebrand. H. Downes, G. Pearson and two anonymous reviewers are thanked for their comments that
1245
JOURNAL OF PETROLOGY
VOLUME 49
helped to focus the ideas expressed in this paper. We
acknowledge the Norwegian Research Council (NFR) for
financial support, including a post-doctoral fellowship to
N.S.C.S. The work has also been supported by a Center of
Excellence grant from the Norwegian Research Council
to Physics of Geological Processes, University of Oslo.
Funding to pay the Open Access publication charges for
this article was provided by Norwegian Research Council
Project 166535/V30.
S U P P L E M E N TA RY DATA
Supplementary data for this paper are available at Journal
of Petrology online.
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