JOURNAL OF PETROLOGY
VOLUME 42
NUMBER 1
PAGES 55–67
2001
Fluid and Element Cycling in Subducted
Serpentinite: a Trace-Element Study of the
Erro–Tobbio High-Pressure Ultramafites
(Western Alps, NW Italy)
MARCO SCAMBELLURI1∗, ELISABETTA RAMPONE2 AND
GIOVANNI B. PICCARDO1
1
DIPARTIMENTO PER LO STUDIO DEL TERRITORIO E DELLE SUE RISORSE, UNIVERSITÀ DI GENOVA,
CORSO EUROPA 26, 16132 GENOVA, ITALY
2
DIPARTIMENTO DI SCIENZE DELLA TERRA, UNIVERSITÀ DI MILANO, VIA BOTTICELLI 23, MILANO, ITALY
RECEIVED NOVEMBER 16, 1999; REVISED TYPESCRIPT ACCEPTED JULY 4, 2000
The oceanic serpentinization of peridotites and the influence of such
an alteration on element cycling during their subduction dewatering
are here investigated in a mantle slice (Erro–Tobbio peridotite), first
exposed to oceanic serpentinization and later involved in alpine
subduction, partial dewatering and formation of a high-pressure
olivine + titanian-clinohumite + diopside + antigorite assemblage
in the peridotites and in veins. Previous work indicates that highpressure veins include primary brines, representing a residue after
crystallization of the vein assemblage and containing recycled oceanic
Cl and alkalis. To reconstruct the main changes during oceanic
peridotite serpentinization and subsequent subduction, we analysed
samples in profiles from serpentinized oceanic peridotites to highpressure serpentinites, and from high-pressure ultramafites to veins.
Here we present results indicating that the main features of the
oceanic serpentinization are immobility of rare earth elements (REE),
considerable water increase, local CaO decrease and uptake of trace
amounts of Sr, probably as a function of the intensity of alteration.
Sr entered fine-grained Ca phases associated with serpentine and
chlorite. Trace-element analyses of mantle clinopyroxenes and highpressure diopsides (in country ultramafites and veins), highlight the
close similarity in the REE compositions of the various clinopyroxenes, thereby indicating rock control on the vein fluids and lack
of exotic components carried by externally derived fluids. Presence
of appreciable Sr contents in vein-forming diopside indicates cycling
of oceanic Sr in the high-pressure fluid. This, together with the
recognition of pre-subduction Cl and alkalis in the vein fluid,
indicates closed-system behaviour during eclogitization and internal
∗Corresponding author. Telephone: +39-010-3538315. Fax: +39010-352169. E-mail: [email protected]
cycling of exogenic components. Diopside and Ti-clinohumite are
the high-pressure minerals acting as repositories for REE and Sr,
and for high field strength elements (HFSE), respectively. The
aqueous fluid equilibrated with such an assemblage is enriched in
Cl and alkaline elements but strongly depleted in REE and
HFSE (less than chondrite abundances). Sr is low [(0·2–1·6)
× chondrites], although selectively enriched relative to light REE.
KEY WORDS: eclogite
facies; fluid; trace elements; serpentinite; subduction
INTRODUCTION
Cycling of crustal materials in subduction zones is accompanied by production of fluids and/or melts whose
transfer from the subducting plate to the overlying mantle
controls the onset of magmatism at convergent plate
margins. Distinctive signatures of arc basalts include
depletion in high field strength elements (HFSE) and
selective enrichment in large ion lithophile elements
(LILE), light rare earth elements (LREE), 10Be, B and
Cl. This indicates contamination of their mantle sources
by solute-rich agents containing several exogenic components, e.g. 10Be and Cl (Tera et al., 1986; Philippot et
al., 1998). Although some experimental studies constrain
 Oxford University Press 2001
JOURNAL OF PETROLOGY
VOLUME 42
NUMBER 1
JANUARY 2001
basin (Hoogerduijn Strating et al., 1993; Scambelluri et
al., 1995). During subduction, ductile deformation was
focused in serpentinite shear zones (high-strain domains)
surrounding volumes of serpentinized peridotite unaffected by the plastic deformation (low-strain domains).
The latter still preserve mineralogical and textural records
of both mantle and oceanic history (Fig. 1).
In low-strain domains, granular spinel lherzolites are
overprinted by tectonitic to mylonitic lherzolites showing
decompressional recrystallization to spinel-, plagioclaseand hornblende-bearing assemblages (Hoogerduijn
Strating et al., 1993). Subsequent intrusion of mid-ocean
ridge gabbros and basalts indicates the early shallow
exposure of this lithospheric mantle in an oceanic setting.
Later peridotite serpentinization (and local rodingitization of the mafic dykes) caused development of
chlorine-bearing chrysotile and lizardite, chlorite, magnetite, minor brucite, and chlorine- and alkali-bearing
phyllosilicates. The low-grade nature of such an assemblage and the mineral compositions point to peridotite
interaction with Cl- and alkali-bearing solutions, presumably seawater derived (Scambelluri et al., 1997). In
low-strain domains, subduction led to patchy replacement
of mantle and hydrothermal assemblages by radial aggregates of antigorite, fine-grained olivine, titanian-clinohumite, magnetite and diopside (shaded area around
profile C, Fig. 1). The peridotite volumes recording this
recrystallization display olivine + titanian-clinohumitebearing veins, coeval with an eclogitic foliation in the
associated mafic dykes. This indicates a high-pressure
origin of the olivine-bearing assemblages in ultramafites,
and production of eclogitic vein fluids.
In the high-strain domains, serpentinite mylonites display an antigorite foliation enclosing boudins of eclogitized metagabbro and metarodingite, and cut by olivinebearing veins and later olivine + antigorite shear bands
(Fig. 1). The olivine assemblage thus represents the
highest grade achieved during subduction: 2–2·5 GPa
and 550–600°C based on the eclogitic paragenesis in
metagabbros (Messiga et al., 1995). The altered Erro–
Tobbio peridotites thus underwent high-pressure partial
deserpentinization producing peak olivine + titanianclinohumite + fluid in presence of stable antigorite. The
high-pressure minerals lack Cl and alkalis, and veins
contain primary hypersaline fluid inclusions with sodium
and potassium chlorides as main daughter crystals (Scambelluri et al., 1997). This highlights partitioning of chlorine
and alkalis into the eclogitic fluid and cycling of oceanic
substances at eclogite facies.
the elemental fluxes in subduction zones through element
partitioning between minerals and fluids or melts (Brenan
et al., 1995a, 1995b; Ayers et al., 1997; Stalder et al.,
1998), little is known about the actual properties and
compositions of natural agents and about their release
during subduction. Studies of fluid inclusions in alpine
eclogite-facies rocks recognize saline aqueous solutions
as potential carriers of major and trace elements (Philippot
& Selverstone, 1991; Scambelluri et al., 1997). However,
these fluids are often residues after rock dehydration,
fluid–rock exchange and precipitation of eclogitic veins
(Scambelluri & Philippot, 2000), and at present there is
little evidence of pristine eclogitic fluids, and there are
few analyses of such fluids.
Concerning the process of fluid release at convergent
plate margins, subduction of hydrous oceanic mantle is
an important variable, as it represents a major source of
deep dehydration water. Serpentinized peridotites are
widespread in oceanic basins, and can contain 10–
13 wt % bulk H2O fixed in serpentine and associated
hydrous phases after interaction with seawater. Experiments and petrology of alpine peridotites indicate
that serpentine stability to very high pressures allows
considerable fluxing of surface waters into the mantle
(Scambelluri et al., 1995; Ulmer & Trommsdorff, 1995).
Formation of high-pressure veins with primary fluid
inclusions is compelling evidence of deep fluid release in
subducted serpentinites. The trapped fluid dissolves up
to 50 wt % chlorine and alkalis, i.e. components formerly
hosted in oceanic hydrous minerals and then inherited
by the subduction-zone fluid (Scambelluri et al., 1997).
The compositional relationships between oceanic and
high-pressure assemblages thus become crucial to define
deep fluid and element cycling by serpentinized peridotites and to assess the control of pre-subduction alteration on the composition of subduction fluids. Allied
questions concern the other components involved in the
transfer process, and the control exerted by the highpressure minerals on the trace-element composition of
associated fluids.
We investigate the above features in the alpine Erro–
Tobbio peridotite, a slice of hydrous oceanic mantle
involved in alpine subduction and eclogitization, by studying the compositional features related to oceanic serpentinization and later subduction of these rocks.
GEOLOGICAL AND PETROLOGICAL
BACKGROUND
The Erro–Tobbio peridotite (Ligurian Alps) corresponds
to subcontinental mantle first exhumed and hydrated
during opening of the Jurassic Tethyan ocean, and then
involved in alpine subduction and high-pressure recrystallization during subsequent closure of the oceanic
SAMPLE DESCRIPTION AND
ANALYTICAL PROCEDURES
To study the compositional variations during serpentinization and the element redistribution during deep
56
SCAMBELLURI et al.
FLUID AND ELEMENT CYCLING IN SUBDUCTED SERPENTINITE
Fig. 1. Internal structure of the Erro–Tobbio unit (not to scale), showing textures in low- and high-strain domains [redrawn after Scambelluri
et al. (1997)]. Circled numbers refer to mantle and subduction structures. Also shown are samples in profiles A and B from cores of serpentinized
peridotite (plagioclase peridotite in A and spinel tectonite in B) to serpentinite mylonites. Insets display the high-pressure recrystallization in lowand high-strain domains, together with location of samples in profiles from host-rock to wall-rock to vein (profile C is in a low-strain zone;
profile D is in a high-strain zone).
dewatering and veining, the analytical work was focused
on: (1) two profiles from cores of peridotites with variable
oceanic serpentinization (with relict mantle assemblages
and no high-pressure minerals), to high-pressure serpentinite mylonite shear zones (Fig. 1; profiles A and B);
(2) several vein to host-rock profiles in low- and highstrain domains (Fig. 1; profiles C and D, respectively).
In profiles A and B, samples were collected at regular
distances of about 15 m. Profile A is composed of ETA46,
ETA47 and ETA51. ETA46 is a serpentinized peridotite
tectonite core consisting of olivine, clinopyroxene,
orthopyroxene and spinel porphyroclasts partly recrystallized into plagioclase-facies assemblages. Olivine
is overgrown by oceanic chrysotile and lizardite, plagioclase is replaced by chlorite ± epidote, and thin
carbonate veinlets are locally present. No alpine highpressure minerals are present. In ETA47 (border of
the preserved lherzolite), the same mantle and oceanic
textures and assemblages as in ETA46 are statically
overgrown by the peak alpine assemblage, i.e. antigorite,
olivine, titanian-clinohumite and diopside. ETA51 (the
alpine serpentinite mylonite) shows a prograde foliation
of prevalent antigorite, chlorite, magnetite, subordinated
diopside and titanian-clinohumite. Porphyroclasts of relict mantle clinopyroxene are diffusely preserved. This
foliation is cut by shear bands with fine-grained olivine,
antigorite and titanian-clinohumite. Profile B is composed
of ETF1, ETF2, ETF3, ETF4, ETF6 and ETF7. ETF1,
ETF2 and ETF3 represent a preserved spinel tectonite
core. ETF1 (the inner part) has a tectonite foliation of
olivine, clinopyroxene, orthopyroxene and spinel porphyroclasts partly recrystallized in granoblastic aggregates
containing the same spinel-bearing assemblage. The
mantle minerals are overgrown by oceanic chrysotile and
lizardite. ETF2 and ETF3 have the same structure and
mineralogy as ETF1 and display an increasing intensity
of serpentinization (50 vol. % chrysotile and lizardite in
ETF2, 55 vol. % in ETF3, based on X-ray diffraction).
In ETF3 fine-grained Ca phases coexist in veins with
chlorite and serpentine. Samples ETF4, ETF6 and ETF7
form a shear-zone surrounding the spinel peridotite
(ETF7 being the inner part of the mylonite zone). They
contain some mantle clinopyroxene relics in a foliated
matrix of high-pressure antigorite, chlorite, magnetite
and minor diopside. This foliation is cut by olivine +
antigorite + titanian-clinohumite shear bands.
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1·87–2·68 wt %; CaO = 1·97–2·38 wt %), and preserve
Al/Ca ratios comparable with mantle values (Hofmann,
1988). One exception is the serpentinite ETA51, which
has 2·42 wt % Al2O3 coupled with rather low CaO
(1·67 wt %), probably resulting from a higher intensity
of serpentinization and Ca loss caused by partial clinopyroxene dissolution.
The trace-element compositions of serpentinized peridotites and serpentinite mylonites in both profiles (Table
1; Fig. 2) do not display significant REE variability. As
classically observed in orogenic peridotites, the studied
samples show overall depletion of LREE relative to heavy
REE (HREE) (CeN/YbN = 0·015–0·076). All analysed
peridotites display negative Zr anomalies with respect
to adjacent REE: comparable negative Zr anomalies
observed in other depleted peridotites (e.g. the Internal
Liguride peridotites of the Northern Apennines) were
ascribed to different solid–melt partition coefficients of
these elements during fractional melting (Rampone et al.,
1996). Thus, the Zr anomaly is probably inherited from
earlier mantle melting. On the other hand, Fig. 2 illustrates that ultramafites in profiles A and B display
increasing Sr from serpentinized oceanic peridotite to
serpentinite mylonites (e.g. samples ETA46 to ETA51).
Comparable positive Sr anomalies are not normally
observed in depleted mantle peridotites from the Alpine–
Apennine chain, where negative Sr anomalies are
acquired as a result of fractional melting (Rampone et
al., 1996). The measured positive Sr anomalies were thus
acquired during secondary, post-melting, enrichment.
Samples ETA71, ETA74, ETA75, ETA75A, ETA76,
ET42 and ET42A are veins and their wall- and hostrocks in low-strain domains. ETA71, ETA75 and ET42
are olivine + titanian-clinohumite + magnetite + diopside ± chlorite veins with variable mineral abundances: ETA71 contains diopside and accessory Ticlinohumite, olivine and magnetite; ETA75 and ET42
are richer in titanian-clinohumite and olivine, and display
comparable mineralogies and modal proportions.
ETA74, ETA75A and ET42A are peridotites collected
at centimetre distances from the veins (i.e. wall-rocks),
ETA76 is at metre distance (i.e. host-rock). Wall- and
host-rocks do not display significant petrographic differences; they preserve the mantle textures and some mantle
clinopyroxene and olivine. Their major constituents are
peak olivine + titanian-clinohumite + antigorite +
diopside grown as static pseudomorphic replacements.
ETF9, ETF10 and ETF11 constitute a profile from vein
(ETF10) to host-rock serpentinite (ETF9 and ETF11,
>1 m from the vein) in a mylonite zone. These hostrocks display an antigorite + chlorite foliation (with relict
porphyroclasts of mantle olivine and clinopyroxene) cut
by olivine + titanian-clinohumite + antigorite shear
bands. Vein ETF10 contains abundant chlorite.
Whole-rock major-element compositions were determined by X-ray fluorescence (XRF). Trace and rare
earth element (REE) concentrations were measured by
inductively coupled plasma-mass spectrometry (ICP-MS)
with a VG PQ2 instrument, at the Institut des Sciences
de la Terre, de l’Eau et de l’Espace de Montpellier [the
analytical procedure has been described by Ionov et al.
(1992)].
The major-element mineral analyses were performed
by energy-dispersive spectrometry with a Philips SEM
515 at the Dipartimento di Scienze della Terra, Genoa,
using accelerating potential of 15 kV, beam current of
20 nA, counting times of 100 s, and natural mineral
standards. Mineral analyses were also measured by wavelength-dispersive spectrometry using an ARL SEMQ
electron microprobe at 15 kV, sample current of 15 nA,
at the Dipartimento di Scienze della Terra, Milan. Traceelement mineral analyses were carried out with a Cameca
IMS 4F ion microprobe at the CNR-CSCC, Pavia [the
analytical procedure has been described by Bottazzi et
al. (1991)]. Trace-element abundances have been normalized to an average C1 chondrite composition (Anders
& Ebihara, 1982).
Profiles from high-pressure veins to hostrocks
The major-element compositions of veins and host-rocks
are reported in Table 2. Whole wall- and host-rocks
display Al2O3 and CaO variations comparable with those
of ultramafites from profiles A and B. In particular, the
wall- and host-rock samples ETA75 and ETA76 resemble
the serpentinite ETA51 in being depleted in CaO
(0·66 wt % and 0·79 wt %, respectively) relative to Al2O3
(1·72 wt % and 1·87 wt %). The veins display bulk
major-element compositions significantly different from
those of the surrounding peridotites. Moreover, veins
display a compositional variability that reflects their variable mineral abundances: vein ETA71, with the highest
CaO (13·21 wt %), has the highest diopside content (see
Table 2 for modal estimates); ETA75 and ETA72 have
higher TiO2 (1·23 wt % and 1·6 wt %) as a result of
higher titanian-clinohumite content. Vein ETF10 is H2O
and Al2O3 rich because of appreciable modal chlorite.
The absolute REE concentrations in wall- and hostrocks are lower than chondrite values (Table 2; Fig. 3)
and are comparable with those of ultramafites in profiles
BULK-ROCK COMPOSITIONS
Profiles from serpentinized peridotites to
serpentinite mylonites
Most peridotite samples in profiles A and B (Fig. 1; Table
1) display coherent variations in Al2O3 and CaO (Al2O3 =
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SCAMBELLURI et al.
FLUID AND ELEMENT CYCLING IN SUBDUCTED SERPENTINITE
Table 1: Bulk-rock major- and trace-element composition of samples in profiles from serpentinized peridotite
cores to serpentinite mylonite shear zones
ETA46
ETA47
ETA51
ETF1
ETF2
ETF3
ETF4
ETF6
ETF7
wt %
SiO2
41·40
39·63
39·56
40·40
38·88
39·65
38·77
39·86
38·81
TiO2
0·08
0·08
0·08
0·06
0·05
0·05
0·05
0·06
0·06
Al2O3
2·58
2·68
2·42
2·25
1·98
1·94
1·87
2·18
2·12
Fe2O3
8·57
8·01
8·52
8·13
8·10
7·88
7·77
7·34
7·97
MnO
0·12
0·12
0·12
0·11
0·11
0·11
0·10
0·10
0·11
MgO
37·40
35·60
36·12
38·62
38·54
38·86
38·21
37·81
37·42
2·08
CaO
2·38
2·29
1·67
2·24
2·10
2·05
1·97
2·13
Na2O
<0·04
<0·04
<0·04
0·06
0·06
0·06
0·05
0·06
0·05
H2O
6·70
10·68
10·80
9·10
10·73
10·26
11·76
11·20
12·15
Total
99·23
99·09
99·29
100·97
100·55
100·86
88·79
100·74
100·77
ppm
Ce
0·071
0·057
0·075
0·020
0·012
0·016
0·017
0·015
0·032
Nd
0·230
0·206
0·216
0·119
0·083
0·090
0·078
0·094
0·116
Sm
0·143
0·130
0·131
0·084
0·070
0·073
0·062
0·079
0·082
Eu
0·061
0·058
0·053
0·039
0·032
0·033
0·028
0·036
0·038
Gd
0·269
0·257
0·249
0·188
0·159
0·153
0·142
0·168
0·177
Tb
0·054
0·052
0·050
0·041
0·036
0·034
0·030
0·037
0·037
Dy
0·404
0·386
0·367
0·318
0·268
0·258
0·239
0·283
0·285
Ho
0·093
0·089
0·085
0·076
0·064
0·062
0·055
0·067
0·068
Er
0·290
0·281
0·259
0·238
0·203
0·195
0·180
0·215
0·210
Tm
0·045
0·042
0·039
0·035
0·032
0·030
0·029
0·033
0·033
Yb
0·296
0·279
0·258
0·243
0·205
0·202
0·190
0·212
0·214
Sc
12
12
12
Ti
525
465
495
V
Cr
64
62
60
2435
2552
2216
Sr
Y
2·44
—
4·06
—
8·80
—
13
12
—
—
45
39
—
12
—
—
11
—
38
—
35
—
13·5
13
—
13
—
44
—
41
—
4·05
3·64
4·84
9·6
9·4
2
2
2
2
2
2
Zr
1·26
0·99
1·22
0·64
0·43
0·52
0·48
0·53
0·58
Nb
0·022
0·021
0·019
0·024
0·023
0·027
0·018
0·038
0·021
A and B. The absolute REE concentrations of veins are
higher than those of wall- and host-rocks, with the
exception of vein ETF10 (Fig. 3b). In accordance with
the major-element compositions, the trace-element
abundances in veins are highly variable and are controlled
by the different mineralogies. This REE variability is
controlled by the diopside abundance, with vein ETA71
(the richest in diopside) having the highest REE values.
On the other hand, Ti variability depends on modal
titanian-clinohumite: Ti is, in fact, very low in veins
having trace amounts of this mineral (e.g. ETA71,
ETF10). Despite this heterogeneity in absolute concentrations, the trace-element patterns of veins are strikingly similar to those of their wall- and host-rocks (Fig.
3). Similar to REE, Sr is much more concentrated in
veins than in peridotites (except for sample ETF10; Fig.
4) and is broadly correlated with the diopside abundance.
MINERAL COMPOSITIONS
The major-element compositions of mantle and alpine
minerals have been documented in previous studies
(Hoogerduijn Strating et al., 1993; Scambelluri et al.,
1997). In this study major and trace-element analyses
mainly concern clinopyroxene, the most effective repository for trace elements in ultramafites. Analyses were
performed on (1) mantle clinopyroxene relics in profile
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characterized by higher middle REE (MREE) to HREE
concentrations, higher Zr (13–24 ppm), Y, V (205–287
ppm), Sc (53–78 ppm) and Ti (2491–4158 ppm), and
display more pronounced Sr depletion (0·91–1·79 ppm).
Comparable trace-element variations were observed in
clinopyroxenes from ophiolitic peridotites and ascribed to
interphase redistribution during metamorphic transition
from spinel- to plagioclase-facies assemblages, and clinopyroxene re-equilibration with plagioclase (Rampone et
al., 1993). The trace-element compositions of relict mantle
clinopyroxenes preserved in profile A, as well as in hostand wall-rocks (Table 4; Fig. 5a), have the same variability
as mantle clinopyroxenes defining the shaded fields in
Fig. 5a.
The high-pressure diopside displays significantly
different compositions. Compared with mantle clinopyroxenes, diopside has very low Na, Al, Ti and Cr, and
higher Si and Mg values. Concerning trace elements
(Fig. 5b), the REE and Sc concentrations in diopside
are similar to those of mantle clinopyroxene, although
diopside displays a slight HREE depletion (DyN/YbN =
1·4–1·8). The most striking difference between metamorphic diopside and mantle clinopyroxenes is the extreme HFSE depletion (Ti 11–618 ppm; Zr 0·2–1·8 ppm)
and Sr enrichment (13–119 ppm) of diopside.
Fig. 2. Bulk trace-element compositions of samples in profiles A and
B from serpentinized peridotites (open symbols) to serpentinite mylonites
(filled symbols), normalized to chondrite abundances (Anders & Ebihara,
1982).
Titanian-clinohumite
Titanian-clinohumite (Table 5) has 3·4–4·9 wt % TiO2
and F is absent. Abundances of most REE (not reported
in Table 5) are below the detection limits, with the
exception of Yb (0·07–0·35 ppm; Fig. 6a). Titanianclinohumite also incorporates appreciable amounts of
other HFSE, such as Nb (0·2–1·2 ppm) and Zr (1·05–12·2
ppm). Figure 6a compares the Erro–Tobbio titanianclinohumites with those from Val Malenco [Central Alps;
data after Weiss (1997)], formed at much lower pressures
(0·4–0·7 GPa): this figure indicates that titanian-clinohumite is a major repository of HFSE in meta-peridotites
and that the HFSE partitioning is not pressure dependent.
A and in wall- and host-rocks ETA74 and ETA75A,
and (2) high-pressure diopside. Their compositions are
reported in Tables 3 and 4, together with some representative analyses of spinel- and plagioclase-facies Erro–
Tobbio clinopyroxenes. Major and trace elements were
also analysed in high-pressure titanian-clinohumite
(Table 5) coexisting with diopside, to assess its potential
role as HFSE carrier.
Clinopyroxenes
DISCUSSION
Compositional variations during
serpentinization.
The Erro–Tobbio mantle clinopyroxenes display typical
decrease in Al and Na, and increase in Ti and Cr
from spinel- to plagioclase-facies clinopyroxenes (Table
3; Hoogerduijn Strating et al., 1993). Concerning the
trace-element compositions, the shaded fields of Fig. 5a
represent typical clinopyroxene compositions for Erro–
Tobbio spinel- and plagioclase-facies peridotites (Table
4). As a whole, these clinopyroxenes have similar REE
patterns (CeN/YbN = 0·0704–0·1168 and 0·0826–
0·1265, respectively) and share negative Sr, Zr and Ti
anomalies. The plagioclase-facies clinopyroxenes are
Profiles A and B (Fig. 1) include (1) mantle peridotites
with variable oceanic serpentinization, (2) peridotites
with static high-pressure recrystallization, and (3) highpressure serpentinite mylonites. Their analysis thus enables us to discuss the compositional changes that occurred during oceanic alteration and subduction of the
peridotites. In spite of their variable serpentinization
or high-pressure overprint, the analysed samples show
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FLUID AND ELEMENT CYCLING IN SUBDUCTED SERPENTINITE
Table 2: Major- and trace-element bulk composition of veins and their host- and wall-rock peridotites
ETA71
ETA74
ETA75
ETA75A
ETA76
ET42
ET42A
ETF9
ETF10
ETF11
vein
wall-
vein
wall-
host-
vein
wall-
host-
vein
host-
rock
rock
rock
rock
rock
rock
wt %
SiO2
44·53
40·75
38·78
40·96
40·93
37·00
40·63
38·36
38·98
39·24
TiO2
0·18
0·05
1·60
0·03
0·05
1·23
0·07
0·04
0·04
0·07
Al2O3
0·60
1·89
0·51
1·87
1·72
0·57
2·28
1·67
1·95
2·09
Fe2O3
11·34
8·83
17·91
7·99
8·69
17·06
7·81
7·51
8·19
7·82
MnO
0·14
0·12
0·25
0·09
0·12
0·31
0·11
0·11
0·12
0·11
MgO
26·91
37·85
33·19
37·46
38·25
39·79
36·42
38·79
38·52
37·68
CaO
13·21
1·50
5·81
0·79
0·66
0·77
2·00
1·80
1·68
2·31
Na2O
<0·04
<0·04
<0·04
<0·04
<0·04
<0·04
<0·04
0·04
0·05
0·06
H2O
2·56
8·40
2·30
9·84
9·47
3·40
9·26
12·61
11·31
11·20
Total
99·47
99·39
100·35
99·03
99·89
100·13
98·58
100·93
100·84
100·58
ppm
Ce
0·362
0·073
0·212
0·053
0·045
0·74
0·069
0·058
0·065
Pr
0·115
0·017
0·058
0·012
0·011
0·134
0·019
0·015
0·015
0·020
Nd
0·91
0·125
0·429
0·097
0·084
0·91
0·147
0·105
0·113
0·142
Sm
0·66
0·066
0·293
0·048
0·047
0·392
0·104
0·065
0·065
0·091
Eu
0·241
0·027
0·116
0·020
0·018
0·124
0·038
0·031
0·028
0·043
Gd
1·58
0·127
0·67
0·095
0·079
0·72
0·207
0·128
0·136
0·188
Tb
0·340
0·0259
0·138
0·0195
0·0158
0·12
0·042
0·029
0·029
0·040
Dy
2·68
0·199
1·07
0·143
0·124
0·82
0·324
0·223
0·218
0·317
Ho
0·621
0·046
0·236
0·034
0·029
0·178
0·073
0·054
0·052
0·074
Er
1·90
0·145
0·71
0·113
0·096
0·52
0·225
0·170
0·162
0·222
Tm
0·261
0·0226
0·099
0·0169
0·0154
0·076
0·034
0·027
0·027
0·035
Yb
1·60
0·152
0·64
0·115
0·104
0·57
0·229
0·167
0·191
0·233
Sc
Ti
53
7
41
8
9
1078
350
9590
175
270
33
—
12
440
V
53
49
126
47
44
82
62
Cr
813
2024
634
1886
1765
689
2336
Sr
Y
38·8
—
6·12
—
20·8
—
3·01
—
2·68
—
29·3
—
5·56
—
11
—
13
—
31
—
0·082
14
—
34
—
43
—
8·70
7·40
8·50
1·47
1·45
1·99
Zr
1·64
0·84
3·22
0·48
0·46
2·88
0·82
0·75
0·88
0·95
Nb
0·058
0·0181
0·458
0·0184
0·0187
0·282
0·018
0·023
0·028
0·030
Vein modes (%)
Ol
5·3
23·1
35·2
43·5
Di
53·6
25·5
6·0
6·4
Ti-cl
35·4
41·2
52·9
5·7
10·1
6·1
Mt
Chl
0·1
50
overlapping REE patterns, thus proving REE immobility
during the overall history. Also, peridotites affected by
oceanic low-grade alteration display considerable water
increase and ubiquitous Sr enrichment. Compared with
unaltered Northern Apennine depleted peridotites (Rampone et al., 1996), the Erro–Tobbio serpentinized peridotites have significantly higher Sr (up to 5 ppm compared
with 0·08–0·15 ppm Sr; Fig. 7). Moreover, oceanic serpentine and phyllosilicates occurring in these rocks can
contain appreciable chlorine and alkali contents (Scambelluri et al., 1997). Comparable REE behaviour and Sr
enrichment were observed by Menzies et al. (1993) after
peridotite–seawater experiments at 300°C. According
to Menzies et al. (1993) serpentinization of precursor
61
JOURNAL OF PETROLOGY
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JANUARY 2001
Fig. 4. Sr vs H2O diagram showing the bulk compositions of veins
and of associated host- and wall-rocks.
altered peridotites that preserved an oceanic Sr imprint
during static eclogitization. Transition to serpentinite
mylonites is accompanied by further increase in water
and Sr (10–15 ppm), and two possible explanations
can be given for this enrichment: (1) it is an oceanic
inheritance, related to high water/rock ratios causing
higher Sr uptake in the strongly serpentinized horizons;
(2) it is an effect of Sr redistribution during high-pressure
metamorphism, caused by channelling of Sr-bearing
fluids into the shear zones. Although there is no definitive
argument against the second hypothesis, several lines of
evidence suggest that the Sr came primarily from an
oceanic source. On the basis of the oxygen isotope
analysis of the same rock samples as studied in this work,
Vallis (1997) and Frueh-Green et al. (2001) demonstrated
that high-pressure metaperidotites and serpentinite mylonites display widespread oxygen isotope heterogeneities
reflecting oceanic pre-subduction signatures. Comparable
heterogeneities are common in eclogites (Philippot et al.,
1998) and suggest that the high-pressure recrystallization
was not associated with isotopic re-equilibration caused
by large-scale fluid fluxing. Fluid redistribution was probably limited to centimetre to metre scales. Also, the Sr
values of our samples are well within the ranges of Sr
concentrations measured in abyssal serpentinites and
obtained experimentally. Consequently, the Sr and water
variability of Fig. 7 probably reflects pre-subduction
heterogeneities in serpentinization and in water distribution within the oceanic lithosphere.
Our study thus indicates that the main changes during
serpentinization consist of uptake of trace amounts of
strontium, chlorine and alkalis, associated with significant
addition of water. This exerted an important control on
the composition of fluids produced during subsequent
subduction and high-pressure recrystallization of the
Erro–Tobbio peridotite.
Fig. 3. Bulk trace-element compositions of samples in profiles from
host- to wall-rock ultramafites, to veins. (a) Low-strain domains; (b)
high-strain domains (see Fig. 1). Normalizing values as in Fig. 2.
lherzolite with Sr <8 ppm brings about an increase
of bulk Sr to levels of 20–60 ppm. Comparable Sr
concentrations were reported for abyssal serpentinites
(Bonatti et al., 1970). The Sr enrichment of the Erro–
Tobbio serpentinized peridotites therefore records an
interaction with seawater-derived solutions during
oceanic alteration. At this stage Sr was probably incorporated in Ca-rich phases (carbonate, tremolite, epidote and diopside) replacing mantle plagioclase and
clinopyroxene. The undeformed Erro–Tobbio high-pressure metaperidotites display Sr contents (2–6 ppm) similar
to those of the serpentinized peridotites with no highpressure imprint (Fig. 7): they could represent formerly
62
SCAMBELLURI et al.
FLUID AND ELEMENT CYCLING IN SUBDUCTED SERPENTINITE
Table 3: Representative major-element composition of mantle and high-pressure clinopyroxenes
Mantle cpx
High-pressure diopside
Reference mantle cpx
Peridotite–serpentinite profile
Wall-rocks
ET4/3
ETA46
ETA74
ETA75A
ETA71
ETA75
sp-facies sp-facies pl-facies
ET4/1
ET9/4
relict cpx relict cpx relict cpx
ETA47
ETA51
relict cpx relict cpx
ETA75A
rock
vein
vein
wt %
SiO2
50·83
50·84
52·71
50·51
50·13
50·57
49·60
51·96
55·37
55·21
54·69
TiO2
0·39
0·43
0·76
0·62
0·41
0·57
0·45
0·40
0·01
<0·01
<0·01
Cr2O3
0·92
0·96
0·99
1·19
1·02
1·12
1·04
1·15
0·20
0·41
0·37
Al2O3
7·04
6·53
3·92
5·51
6·74
5·22
7·34
3·55
0·18
0·13
0·19
FeO
3·13
2·72
2·78
3·31
3·34
3·20
3·03
2·72
0·98
1·58
1·29
MnO
0·09
0·09
0·00
0·11
0·12
0·11
0·12
0·10
0·05
0·10
0·01
MgO
17·11
15·95
16·18
15·80
15·65
16·11
14·33
16·05
17·32
17·17
17·37
CaO
19·67
21·66
22·80
21·75
22·14
21·80
23·01
24·25
25·52
25·43
25·34
Na2O
0·58
0·73
0·39
0·43
0·34
0·40
0·66
0·45
0·01
<0·01
<0·01
K2O
<0·01
<0·01
0·09
0·01
<0·01
<0·01
0·01
<0·01
0·01
<0·01
0·07
Total
99·76
99·91
100·62
99·24
99·89
99·10
99·59
100·63
99·65
100·03
99·33
0·90
0·91
0·91
0·90
0·90
0·90
0·90
0·91
0·97
0·95
0·96
XMg
Reference mantle cpx—clinopyroxenes from Erro–Tobbio spinel and plagioclase peridotites.
The high-pressure recrystallization
closed-system behaviour of the Erro–Tobbio ultramafites
during eclogitization and internal cycling of oceanic water
and components. The data presented here and the stable
isotope heterogeneities diffusely documented in highpressure and very high pressure rocks [Philippot et al.
(1998), and reference cited therein), point to restricted
fluid mobility during eclogitization. This contradicts the
large-scale mobility required to flux slab fluids into the
mantle, and suggests that fluids may remain entrapped
in the slab, until they are injected into the upper mantle
at depths greater than those attained by most exposed
eclogites. The latter may not represent the levels at which
fluids become extracted from the slab and higher-pressure
rocks need to be exhumed (or discovered) to find the
regions of important fluid loss from the slab. Watson et
al. (1990) proposed that transition from immobile to
mobile fluids in the mantle is related to concurrent
deformation, and to decrease of wetting angles below a
critical value of 60°, which occurs at higher pressure and
temperature. This points to a depth-related change in
fluid mobility with continuing subduction. Alternatively,
recent experiments of Watson & Wark (1997) indicate
that SiO2 diffusion into an immobile fluid can occur over
kilometre-scale distances at rates comparable with those
of thermal conductivity. These results have been taken
by Philippot & Rumble (2000) to explain both limited fluid
circulation and mantle metasomatism in deep subduction
Petrography and compositions of the Erro–Tobbio highpressure ultramafites and veins provide constraints on
the mobility and composition of fluids released during
partial deserpentinization at eclogite facies. Veins and
host- and wall-rocks have the same high-pressure mineralogy, indicating control by the country peridotite on
the components dissolved in the fluid. Similarity between
the REE patterns of clinopyroxenes in the veins and in
the host-rock also reflects such a rock-control on the
fluid. In particular, the bulk REE concentrations in
veins are dominated by diopside, whose REE absolute
concentrations and patterns are, in turn, very similar to
those of rock-forming mantle clinopyroxene relics and
high-pressure diopside in peridotites. These comparable
geochemical signatures in clinopyroxene suggest a lack
of infiltration of external fluids transporting exotic components during the high-pressure metamorphism. Rather,
internally produced fluids equilibrated with, and were
compositionally controlled by, the surrounding ultramafites. Further indication of local derivation of rock
components in the fluid is the occurrence of Sr-bearing
diopside in the veins: this also represents compelling
evidence for cycling of marine Sr in the high-pressure
fluid. These data, together with previous recognition that
pre-subduction chlorine and alkalis were released in
the dehydration fluid (Scambelluri et al., 1997), indicate
63
JOURNAL OF PETROLOGY
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Table 4: REE and trace-element composition of mantle and high-pressure clinopyroxenes
Mantle cpx
High-pressure diopside
Reference mantle cpx
Peridotite–serpentinite profile wall-rocks
ET4/3
ET4/3
ET4/1
ET4/1
ET9/4
ET9/4
ET9/4
ETA46
ETA47
ETA51
ETA74
ETA75A
ETA75A
ETA75
ETA75
ETA71
ETA71
sp-
sp-
sp-
sp-
pl-
pl-
pl-
relict
relict
relict
relict
relict
rock
vein
vein
vein
vein
facies
facies
facies
facies
facies
facies
facies
cpx
cpx
cpx
cpx
cpx
ppm
Ce
0·43
0·37
0·49
0·42
0·88
1·27
1·09
0·70
0·61
0·62
0·35
0·93
0·27
0·19
0·43
0·31
0·43
Nd
1·22
1·41
1·63
1·30
3·99
4·71
2·85
3·43
3·50
3·40
2·03
2·91
0·87
0·81
1·40
1·07
1·23
Sm
0·94
0·69
1·16
1·04
1·98
2·68
1·66
1·90
2·22
2·14
1·37
1·87
0·76
0·60
1·40
0·86
1·42
Eu
0·79
0·56
0·441
0·46
0·89
1·07
0·71
0·80
0·75
0·71
0·56
0·50
0·52
0·24
0·54
0·31
0·42
Gd
—
—
—
—
—
—
—
3·42
3·20
3·30
2·40
3·10
1·83
0·95
2·94
1·86
3·42
Dy
2·27
1·86
2·22
2·26
5·53
5·84
3·41
4·30
4·20
4·20
2·80
3·60
2·61
1·35
4·84
2·37
2·43
Er
1·45
1·10
1·32
1·49
2·83
3·16
2·18
2·56
2·79
2·46
1·76
2·00
1·70
0·79
2·95
1·49
2·34
Yb
1·60
1·35
1·09
1·30
2·78
3·10
2·23
2·18
2·48
1·97
1·62
1·80
1·34
0·52
2·18
0·86
Sc
1·98
48
51
48
52
78
63
53
73
66
67
52
60
69
51
83
19
41
Ti
1900
2073
1888
1997
4158
2492
2892
2997
2656
3247
2180
2299
11
15
618
15
15
V
188
201
192
185
287
223
206
299
331
346
253
294
14
12
Cr
4887
5162
5776
4921
5687
4239
4545
6732
7511
7504
6981
7521
213
298
114
110
233
119
22
43
13
22
Sr
2·95
3·18
2·48
2·67
0·91
1·41
1·79
Zr
7·13
7·86
7·21
7·44
24·37
12·96
15·94
Nb
—
—
—
—
—
—
—
0·53
15
—
1·12
10
0·94
16
0·13
0·11
8·3
0
0·09
1·64
11
0·18
1·06
0·20
0·02
0·01
8·33
1·8
—
6·0
0·89
0·02
7·81
0·76
—
Reference mantle cpx—clinopyroxenes from Erro–Tobbio spinel and plagioclase peridotites.
zones. Whether or not the eclogitic fluids may flux the
overlying mantle is thus matter of debate and remains
an open problem.
Ti-clinohumite–diopside partition coefficients (Fig. 6b)
indicate Sr partitioning into the high-pressure pyroxene
and the affinity of titanian-clinohumite for HFSE (mainly
Ti and Zr); their crystallization was thus accompanied
by incorporation of Sr and HFSE into the respective
mineral phases. Sr was originally held in Ca-rich oceanic
phases. Negative HFSE in the alpine diopside, despite
the lack of negative HFSE anomaly in the whole rocks,
indicates that these elements, primarily hosted in mantle
clinopyroxene that survived subduction, were released
during its breakdown in presence of fluid and partitioned
in titanian-clinohumite. This mineral thus plays a role
in fractionating HFSE from LILE at high pressures,
similarly to other Ti phases such as rutile and ilmenite
(Brenan et al., 1995b; Ayers et al., 1997). Capability of
titanian-clinohumite to store HFSE was firstly envisaged
by Weiss & Muentener (1996) for the low-pressure titanian-clinohumites from Val Malenco peridotite (Central
Alps); our study confirms this hypothesis and shows that
clinohumite acts as a repository for HFSE independent
of pressure conditions.
The vein mineralogies indicate that the dehydration
fluid dissolved appreciable amounts of silicate components from surrounding rocks, and probably corresponded to an aqueous solute-rich solution. Presence
of diopside, titanian-clinohumite and hypersaline fluid
inclusions in the veins indicates that Sr, HFSE, Cl and
alkalis were also soluble in this fluid. Precipitation of vein
minerals consumed silicate components from the fluid,
and left a residual solution enriched in excess water,
which became trapped in the primary fluid inclusions
inside vein minerals (Scambelluri et al., 1997). Crystallization of diopside and titanian-clinohumite thus fixed
Sr and HFSE, respectively, and controlled the traceelement composition of the coexisting aqueous residue.
High Cl and alkali concentrations in the primary hypersaline fluid inclusion indicate partitioning of these
large ions in the fluid residue, rather than in the compact
structures of vein minerals (Scambelluri et al., 1997).
Additional information concerning the trace-element
composition of such a fluid can be achieved using
experimental clinopyroxene–fluid partition coefficients,
which are available for a range of P–T conditions. Partitioning data determined at 2 GPa and 900°C (Ayers et
al., 1997; Brenan et al., 1995a, 1995b), the conditions
64
SCAMBELLURI et al.
FLUID AND ELEMENT CYCLING IN SUBDUCTED SERPENTINITE
closest to P–T estimates of the Erro–Tobbio high-pressure
recrystallization, indicate that most trace elements preferentially enter the clinopyroxene. The coexisting aqueous fluid has extremely low REE and HFSE contents in
the range (0·01–0·1) × chondrite abundances. Sr is also
low—although selectively enriched relative to LREE—
and ranges from (0·2–0·4) × chondrites (Brenan et al.,
1995a) to (0·8–1·6) × chondrites (Ayers et al., 1997). On
the other hand, the fluid is strongly enriched in Cl and
alkaline elements (Scambelluri et al., 1997).
Table 5: Major- and trace-element composition
of Ti-clinohumite from veins and host peridotites
ETA75A
ETA71
ETA75
ETA75
peridotite
vein
vein
vein
wt %
SiO2
36·12
35·84
36·19
36·27
TiO2
3·82
3·86
3·39
3·64
Cr2O3
0·05
<0·01
<0·01
0·10
Al2O3
0·13
<0·01
0·09
0·06
FeO
12·14
12·98
12·86
11·25
MnO
0·47
0·38
0·38
MgO
44·47
43·87
44·56
CaO
0·07
<0·01
<0·01
H2O
Total
SUMMARY
0·44
Our study of the alpine Erro–Tobbio eclogitized peridotite indicates that pre-subduction serpentinization and
alteration had a relevant control on the composition
of fluids evolved during later burial and high-pressure
recrystallization. Serpentinization was accompanied by
REE immobility, uptake of trace amounts of Cl, Sr and
alkalis, and by significant addition of water. Formation of
a high-pressure paragenesis made up of olivine, antigorite,
diopside and Ti-clinohumite in country ultramafites and
veins was the result of one major pulse of fluid released
during subduction. The close similarity in the REE
compositions of mantle clinopyroxenes and high-pressure
diopside indicates rock control on the vein fluids and
absence of exotic components carried by externally derived fluids. Presence of appreciable Sr contents in veinforming diopside indicates cycling of oceanic Sr in the
high-pressure fluid. This, together with the presence
45·6
0·06
3·61
3·55
3·81
3·74
100·88
100·48
101·28
101·16
ppm
Ce
0·415
<0·01
0·367
0·49
Yb
0·07
0·35
0·25
0·33
Sc
Ti
<3
33
55
22
36911
28232
34690
28692
V
39
16
24
20
Cr
186
176
306
190
Sr
0·63
Zr
1·05
12·2
0·26
0·23
7·4
Nb
0·20
1·2
0·65
0·74
10·1
0·75
Fig. 5. Trace-element compositions of clinopyroxenenes from the Erro–Tobbio peridotite. (a) Reference mantle clinopyroxenes (shaded fields:
clinopyroxenes from Erro–Tobbio spinel and plagioclase peridotites), and mantle clinopyroxene relics from profiles of serpentinized peridotites
to serpentinite mylonites, and from wall-rocks encasing the high-pressure veins. (b) High-pressure diopside from host-rocks and veins. Normalizing
values as in Fig. 2.
65
JOURNAL OF PETROLOGY
VOLUME 42
NUMBER 1
JANUARY 2001
Fig. 7. Bulk-rock Sr vs water diagram reporting samples of unaltered
depleted peridotites of the Northern Apennines [data after Rampone
et al. (1996)], and of Erro–Tobbio serpentinized peridotites (ETA46,
ETF1, ETF2 and ETF3), undeformed high-pressure peridotites
(ETA47, ETA74, ETA75A, ETA76 and ET42), and high-pressure
serpentinite mylonites (ETA51, ETF4, ETF6, ETF7, ETF9 and ETF11).
(Brenan et al., 1995a; Ayers et al., 1997), this fluid has
very low REE and HFSE content, is selectively enriched
in Sr relative to LREE, and is Cl and alkali rich.
ACKNOWLEDGEMENTS
We are grateful to Alberto Zanetti and Riccardo Vannucci (Pavia) for ion probe analyses, Jean-Louis Bodinier
(Montpellier) for ICP-MS analyses, and Danilo Biondelli
(Milan) and Laura Negretti (Genoa) for assistance during
electron microprobe and scanning electron microscope
analyses. We thank Jeffrey Alt, Marguerite Godard and
Ian Parkinson for careful and constructive reviews. Special thanks go to Jean-Louis Bodinier for helpful comments and for the editorial handling of this paper. This
work was funded by the Italian MURST within the
project ‘Transformations in subducted materials and mass
transfer to the mantle wedge’.
Fig. 6. (a) Trace-element analyses of titanian-clinohumite from the
Erro–Tobbio ultramafites (data points) and from the Malenco peridotite
(bars; data after Weiss, 1997). Normalizing values as in Fig. 2. (b)
Titanian-clinohumite–diopside partition coefficients, calculated from
measured high-pressure pairs in Erro–Tobbio veins and wall-rocks,
and from low-pressure pairs in the Malenco peridotite [data after Weiss
(1997)].
of pre-subduction Cl and alkalis in brines trapped as
inclusions in vein minerals, indicates internal cycling of
exogenic components into the fluid phase.
Diopside and Ti-clinohumite are the high-pressure
minerals acting as the most effective trace-element repositories: Ti-clinohumite hosts HFSE, diopside mainly
incorporates REE and Sr, and lacks HFSE. The brine
inclusions hosted in vein diopside and Ti-clinohumite
represent a fluid residue after mineral deposition: they
correspond to the aqueous fluid phase that equilibrated
with such an assemblage and that was eventually evolved
at this stage of deep peridotite recrystallization. On the
basis of our trace-element analysis and clinopyroxene–
fluid partition coefficients available in the literature
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