JOURNAL OF PETROLOGY
VOLUME 41
NUMBER 2
PAGES 257–283
2000
Melting of Refractory Mantle at 1·5, 2
and 2·5 GPa under Anhydrous and
H2O-undersaturated Conditions:
Implications for the Petrogenesis of High-Ca
Boninites and the Influence of Subduction
Components on Mantle Melting
TREVOR J. FALLOON∗, LEONID V. DANYUSHEVSKY6
SCHOOL OF EARTH SCIENCES, UNIVERSITY OF TASMANIA, GPO BOX 252-79, HOBART, TAS. 7001, AUSTRALIA
RECEIVED JULY 6, 1998; REVISED TYPESCRIPT ACCEPTED JULY 29, 1999
Boninites are an important ‘end-member’ supra-subduction zone
magmatic suite as they have the highest H2O contents and require
the most refractory of mantle wedge sources. The pressure–temperature
conditions of boninite origins in the mantle wedge are important to
understanding subduction zone initiation and subsequent evolution.
Reaction experiments at 1·5 GPa (1350–1530°C), 2 GPa
(1400–1600°C) and 2·5 GPa (1450–1530°C) between a model
primary high-Ca boninite magma composition and a refractory
harzburgite under anhydrous and H2O-undersaturated conditions
(2–3 wt % H2O in the melt) have been completed. The boninite
composition was modelled on melt inclusions occurring in the most
magnesian olivine phenocrysts in high-Ca boninites from the Northern
Tongan forearc and the Upper Pillow Lavas of the Troodos ophiolite.
Direct melting experiments on a model refractory lherzolite and a
harzburgite composition at 1·5 GPa under anhydrous conditions
(1400–1600°C) have also been completed. Experiments establish
a P, T ‘melting grid’ for refractory harzburgite at 1·5, 2 and 2·5
GPa and in the presence of 2–3 wt % H2O. The effect of 2–3
wt % dissolved H2O produces a liquidus depression in primary
boninite of >112 ± 19°C at a given temperature. The H2Obearing melts, recalculated to 100 wt % anhydrous, are >2–6
wt % higher in MgO, >1–2 wt % higher in SiO2 and >1–1·5
wt % lower in FeO, compared with nominally anhydrous melts at
the same P and T. These differences are consistent with a change
in the melting reaction, resulting in a higher contribution of orthopyroxene to the melt phase, compared with anhydrous conditions.
∗Corresponding author. Telephone: +61-3-62262454. Fax: +61-362232547. e-mail: [email protected]
We conclude that high-Ca boninite petrogenesis requires temperatures
as high as >1480°C at depths of >45 km in the mantle wedge;
these are constraints for any proposed model of intra-oceanic
subduction zones. A comparison of the results from the boninite–
harzburgite reaction experiments with the direct melting experiments
on refractory lherzolite and harzburgite indicates that the influence
of subduction components (included in the composition of the added
model boninite) is to cause high-pressure melting cotectics to move
towards the olivine apex (i.e. to relatively higher pressures) of the
molecular normative projection from diopside onto the base of the
‘basalt tetrahedron’ [Jd + CaTs + Lc–Qz–Ol] compared
with anhydrous melting of normal mantle in the absence of a
subduction component. The subduction component involved in highCa boninite petrogenesis in addition to H2O has relatively high
Al2O3 and Na2O contents. The experimental data from this and
other studies empirically quantify the absolute effect of dissolved
H2O (0·2–21 wt %) on the liquidus depression of olivine-saturated
basaltic melts with ><3 wt % total alkalis as follows:
olivine liquidus depression (° C) = 74·403 × (H2O wt %)0 · 352.
The equation that describes this empirical relationship is non-linear
with an error of >9 relative percent.
high-Ca boninites; mantle melting; H2O-undersaturated
melting; olivine liquidus depression; anhydrous melting
KEY WORDS:
 Oxford University Press 2000
JOURNAL OF PETROLOGY
VOLUME 41
INTRODUCTION
Magma genesis at intra-oceanic subduction zones is fundamentally different from that occurring at mid-ocean
ridges, as a result of the presence of H2O and other
volatiles derived from the dehydrating subducted slab,
the possible involvement of melts derived from the slab,
and the presence of refractory mantle sources in the
mantle wedge above the subducting slab (e.g. Gill, 1981;
Woodhead et al., 1993; Tatsumi & Eggins, 1995; Turner
et al., 1997; Green & Falloon, 1998).
H2O not only acts as an incompatible element, and is
therefore a ‘flux’ for melting, but can also form a supercritical fluid capable of fractionating and variably
enriching subduction zone mantle sources in other incompatible elements. H2O affects the phase relationships
of fertile mantle peridotite in terms of both lowering
the solidus in pressure–temperature ( P–T ) space and
changing the nature of melting reactions (Green, 1973b,
1976; Gaetani & Grove, 1998; Green & Falloon, 1998).
Subduction zone magmas in general have geochemical
features consistent with a significant role of H2O in
magma genesis, i.e. higher degrees of mantle melting
and enrichment in mobile LILE (large-ion lithophile
elements) compared with MORB (mid-ocean ridge basalts) (e.g. Gill, 1981).
Boninites are a rare subduction-related magma type
with higher SiO2 and H2O, and lower TiO2, compared
with island arc tholeiite suites. They lack plagioclase
in rocks more mafic than andesite, and contain very
magnesian olivine (Ol ) phenocrysts (up to Fo94, Crawford
et al., 1989). It is generally accepted that the petrogenesis
of boninite requires T significantly higher (>100–300°C)
than can be reasonably expected (<800°C at <50 km) in
a mantle wedge on the basis of geophysical models (Hsui
& Toksov, 1979). Also, boninite petrogenesis is believed to
involve melting of refractory mantle sources (harzburgitic
consequent to prior extraction of MORB) through addition of an incompatible element enriched phase dominated by H2O. Boninites are in fact the most H2O-rich
magma type known from intra-oceanic subduction zones
(Danyushevsky et al., 1993; Sobolev & Chaussidon, 1996).
The so-called ‘boninite paradox’, is that the highest T
magmatic suite in intra-oceanic arcs also has the highest
H2O content (Sobolev & Chaussidon, 1996). Boninite
suites are therefore an ‘end-member’ subduction-related
suite as they (1) are the most H2O rich, and (2) require
the most refractory sources. A detailed understanding of
boninite petrogenesis is important as a constraint on the
geodynamics of subduction-related magmatic processes.
However, at present the T of boninite formation is
disputed. The controversy has arisen because of differing
opinions on the nature of boninite parental or primary
magma compositions, especially the MgO content of the
boninite magmas, which is reflected in the composition
NUMBER 2
FEBRUARY 2000
of Ol phenocrysts. As the MgO content of mafic magma
is a fundamental control on its liquidus T (e.g. Ford et
al., 1983), it is vital that the MgO content of parental or
primary boninite magmas is correctly established. Equally
important is the effect H2O has on lowering the liquidus
T of MgO-rich boninite magmas. Establishment of formation T of primary boninite is a vital first step in
constructing reasonable models of origin, i.e. do boninites
require deep mantle upwelling, or alternatively can they
be produced by contact melting of metasomatized cold
lithospheric mantle at shallow depths in the mantle
wedge?
At present there is a view that boninite magmas require
T ranging from 1150 to 1350°C, based on several experimental studies on boninite compositions ranging from
7 to 16 wt % MgO with experimental H2O conditions
ranging from anhydrous to water saturated (Tatsumi,
1981, 1982; Tatsumi & Maruyama, 1989; Umino &
Kushiro, 1989; Van der Laan et al., 1989).
On the other hand, on the basis of detailed studies of
melt inclusions hosted by magnesian Ol phenocrysts in
high-Ca boninite (HCB) suites, Sobolev et al. (1993) and
Sobolev & Danyushevsky (1994) concluded that HCB
primary magmas have MgO contents of 19–24 wt %,
H2O contents of 1–2 wt %, and T of formation of between
1450 and 1550°C. Such high Ts are not accounted for
by current geophysical models of mantle wedge thermal
structure and hence, if these high Ts are correct, it has
very important implications for the geodynamics of intraoceanic subduction zones. For example, Danyushevsky
et al. (1995) proposed that such high Ts require the
involvement of mantle plumes.
The conclusions of Sobolev et al. (1993) and Sobolev
& Danyushevsky (1994) were based on: (1) comparing
the compositions of the established primary or parental
HCB magma compositions with those of experimentally
determined, anhydrous partial melts of fertile mantle
compositions with an estimate of the effect of H2O on
the partial melt compositions; (2) an estimate, based on
available experimental data, on the effect of H2O in
lowering the liquidus T of the very magnesian HCB melt
compositions.
Here we attempt to refine the T estimates of
HCB petrogenesis by establishing a ‘melting grid’
for refractory peridotite (harzburgite) affected by a
subduction-related component (SC) under both anhydrous and H2O-undersaturated conditions, and by
directly estimating the liquidus depression caused by
the presence of H2O on such highly magnesian primary
magma compositions.
Our experimental results confirm that remarkably
high Ts (>1480°C) at relatively shallow depths
(>45 km) in the mantle wedge are required for HCB
petrogenesis.
258
FALLOON AND DANYUSHEVSKY
MELTING OF REFRACTORY MANTLE
HIGH-Ca BONINITE PRIMARY
MAGMA COMPOSITION
Sobolev & Danyushevsky (1994) presented a detailed
mineralogical, geochemical and experimental study of
an HCB suite from the northern termination of the
Tonga Trench (Sharaskin et al., 1983; Falloon et al., 1987,
1989; Falloon & Crawford, 1991; Danyushevsky et al.,
1995). Primary melt compositions for this suite were
established on the basis of an experimental study of melt
inclusions in Ol phenocrysts, and numerical modelling of
the reverse of fractional crystallization. The approach of
Sobolev & Danyushevsky (1994) was to search for the
most magnesian Ol phenocrysts present in the boninites,
by crushing the samples and mounting >200–300 Ol
grains for microprobe analysis. The most magnesian Ol
were then examined for suitable melt inclusions for
optically controlled homogenization experiments. Successfully homogenized melt inclusions were analysed by
electron microprobe. The primary melts established by
this technique have high MgO contents ranging from 19
to 24 wt %.
RATIONALE OF EXPERIMENTAL
STUDY
Sobolev et al. (1993) and Sobolev & Danyushevsky (1994)
estimated the P and T of primary HCB magma genesis
(2–2·5 GPa, 1450–1550°C; 2–3 GPa, 1380–1550°C,
respectively) on the basis of:
(1) comparison with anhydrous partial melting experiments on fertile peridotite compositions (Takahashi
& Kushiro, 1983; Falloon & Green, 1987, 1988) using
the normative projection of Walker et al. (1979);
(2) the assumption that the mantle melting cotectics,
defined by the anhydrous experimental data, were linear
in the Walker et al. (1979) projection and would continue
to be so towards more refractory mantle compositions;
(3) an estimate of the effect of 2 wt % H2O additions
based on the experimental studies of Green (1973b, 1976)
and Kushiro (1990), which were also performed on fertile
peridotite compositions.
The rationale of our experimental study is to determine
more accurately the P and T of primary HCB magma
genesis by determining a ‘melting grid’ appropriate for
refractory mantle in the presence of an SC, and to
directly determine the effect of >2 wt % H2O on the
liquidus T of primary HCB magma compositions.
As previous petrological studies on HCB petrogenesis
suggest that the residual mantle at the P and T of magma
generation is clinopyroxene (Cpx) free (Crawford et al.,
1989; Falloon et al., 1989; Sobolev & Danyushevsky,
1994), we have performed reaction experiments between
a model primary HCB and a model refractory harzburgite. The rationale of the experiments is to allow the
boninite to equilibrate with harzburgite at various P and T
values, thus defining the positions of Ol + orthopyroxene
(Opx) + liquid (L) cotectics for refractory peridotite
affected by an SC. The use of reaction experiments
ensures that large areas of glass, in the case of nominally
anhydrous experiments, or glass and quench pyroxenes
in the case of hydrous experiments, are available after
the experiment, which can be analysed free from the
effects of metastable quench crystallization on primary
crystal phases (Fig. 1a), and can also be analysed by
Fourier transform IR (FTIR) spectroscopy for H2O contents.
We should be able to demonstrate that at some P and
T the primary HCB identified by melt inclusion studies
is saturated in mantle minerals (specifically Ol and Opx),
and the composition of the saturated melt is close to
that determined by the melt inclusion technique. Two
significant technical problems that could affect the outcome of the melt inclusion studies are: (1) overheating of
the inclusions during homogenization causing estimated
primary magma compositions to have incorrectly high
MgO contents; (2) poor quenching, especially in H2Obearing MgO-rich melt compositions. An additional
problem is the limited availability of suitable sized melt
inclusions in Ol phenocrysts of the most magnesian compositions.
Our boninite–harzburgite reaction experiments allow
us to determine the effect of an SC on melting mantle.
Petrological studies of subduction zone magmatic suites
clearly document the influence of SCs on the geochemistry of erupted magmas (e.g. Turner et al., 1997).
Although geochemical tracers can be used to identify the
various potential sources (e.g. sediment) of the SC (Turner
et al., 1997), it is difficult to constrain the composition of
the SC independently. The simple addition of H2O as
an analogue for the SC in peridotite melting experiments
(Green, 1976), although providing data on the effects of
H2O as a pure component on mantle melting, may
not be directly applicable to natural subduction-related
magmatic suites. An advantage of boninite–harzburgite
reaction experiments is that we are adding a melt component closely matching a natural magma type, which
is itself the result of subduction zone melting. Therefore
in our experiments the actual SC involved in HCB
genesis is present in the composition of the added boninite
itself, even though we cannot independently isolate it.
To determine the effect of this SC on mantle melting
we also performed melting experiments on a refractory
lherzolite and harzburgite under anhydrous conditions
with no added SC to compare with the boninite–
harzburgite reaction experiments.
259
JOURNAL OF PETROLOGY
VOLUME 41
NUMBER 2
FEBRUARY 2000
Fig. 1. (a) Plane-polarized transmitted light photograph of ‘anhydrous’ run product T-3572 (1·5 GPa, Table 2). The boninite-melt layer (top)
has mostly quenched to glass. Dark band separating the melt layer from crystals (bottom) is a zone of quench crystallization extending out from
the harzburgite layer. Width of view is >2 mm. (b) and (c) are reflected light and plane-polarized transmitted light photographs, respectively,
of the ‘wet’ run product T-3464 (2 GPa, 1500°C, Table 2). (b) shows that the melt layer has quenched to large areas of glass and coarse quench
pyroxene, and does not reveal the presence of graphite in the melt layer. However, (c) shows that the melt layer contains a broad band of
disseminated graphite derived from the graphite capsule. Width of view >2 mm.
260
FALLOON AND DANYUSHEVSKY
MELTING OF REFRACTORY MANTLE
analytical grade oxides (SiO2, Al2O3 and MgO) and
synthetic fayalite to a hydrous natural glass, also from
the Troodos ophiolite (sample A-46, Table 1, Danyushevsky et al., 1993). This mixture was ground under
acetone in an agate mortar before storage in glass vials
in an oven at 110°C. The resulting starting composition
(BON, Table 1) has H2O content of 2·4 wt %, identical
to that determined for Tongan HCB primary magmas
by melt inclusion studies (Sobolev & Danyushevsky,
1994). The advantage of using a natural glass as a source
for H2O is that it eliminates any uncertainty in the exact
amount of starting H2O and therefore allows us to
monitor changes in H2O contents during our experiments.
Starting compositions T-4347 and TQ-40 were prepared in a similar manner to HZ above. Composition
T-4347 is the glass composition from run T-4347 (Table
2) and was used for reversal experiments at 1·5 GPa (see
text below). Composition TQ-40 is a refractory lherzolite
(minus 40 wt % Fo91·9) composition modelled on the
natural Tinaquillo Lherzolite ( Jaques & Green, 1980).
The rationale of using Ol-depleted compositions in experimental studies has been presented by Falloon &
Green (1987).
For anhydrous experiments, the BON starting composition was dehydrated by firing at 1000°C in an Ar
atmosphere for 24 h.
Table 1: Starting compositions used in the
experimental study
No.:
1
2
3
4
5
6
2PKI-38
HZ
A-46
BON
T-4347
TQ-40
SiO2
44·00
46·62
59·45
51·88
53·81
47·50
TiO2
0·03
0·08
1·21
0·70
1·07
0·13
Al2O3
0·38
0·65
15·16
8·71
5·97
5·35
FeO∗
8·17
4·91
10·61
8·94
7·89
7·51
MnO
0·16
0·15
0·21
0·12
0·27
0·18
MgO
46·41
46·19
2·69
20·69
22·50
32·80
CaO
0·43
0·73
7·43
7·10
6·48
4·97
Na2O
0·06
0·05
2·96
1·70
0·79
0·30
0·28
0·16
K2 O
0·02
0·03
Cr2O3
0·34
0·57
H2O
mg-no.
91·0
94·4
4·0
2·4
31·1
80·5
0·37
0·03
0·86
0·75
83·6
88·6
Compositions: 2PKI-38, natural harzburgite from the Troodos
ophiolite (Sobolev et al., 1993); HZ, synthetic harzburgite
composition modelled on 2PKI-38 (see text); A-46, hydrous
natural glass from the Troodos ophiolite (Danyushevsky et
al., 1993); BON, synthetic boninite composition modelled on
primary HCB magmas determined by Sobolev & Danyushevsky (1994) (see text); T-4347, synthetic composition of
glass in run T-4347 (Table 2); TQ-40, synthetic refractory
lherzolite composition modelled on Tinaquillo Lherzolite
(Jaques & Green, 1980). All compositions are normalized to
100% anhydrous. FeO∗ refers to total Fe as FeO.
EXPERIMENTAL AND ANALYTICAL
TECHNIQUES
Experimental techniques
Starting compositions
All starting compositions used in the experimental study
are presented in Table 1. The harzburgite composition
(HZ, Table 1) was modelled on the composition of a
natural harzburgite from the Troodos ophiolite (sample
2PKI-38, Table 1, Sobolev et al., 1993). The composition
of 2PKI-38 was recalculated to an mg-number of 94
and 40 wt % Ol of Fo93·5 was subtracted. The starting
composition was prepared from a mixture of analytical
grade oxides and carbonates (Ca, Na, K), ground under
acetone in an agate mortar. This mixture was pelletized
and sintered overnight (>16–20 h) at 950°C. An appropriate amount of synthetic fayalite was then added to
the sintered mix and the mixture was again ground under
acetone, before storage in glass vials in an oven at 110°C.
The HCB composition (BON, Table 1), modelled on
the HCB compositions established by the studies of
Sobolev et al. (1993) and Sobolev & Danyushevsky (1994),
was prepared by addition of appropriate amounts of
Run assemblies and temperature control
All experiments were performed using standard pistoncylinder techniques in the High Pressure Laboratory
housed in the School of Earth Sciences, University of
Tasmania (UT). All experiments used talc–Pyrex assemblies (except runs 4–11, Table 2, which used NaCl–
Pyrex assemblies) with graphite heaters and a W97Re3/
W75Re25 (W/Re) thermocouple (calibrated against the
melting point of Au and the e.m.f. of a Type S, Pt/
Pt90Rh10 (Pt) thermocouple, at atmospheric P in an Ar
atmosphere). Temperatures were controlled to within
±1°C of the set point using a Eurotherm type 818
controller. No P correction was applied to the thermocouple calibration. For the majority of experiments
starting compositions were loaded into inner graphite
capsules and then sealed by welding into large capacity
(3·5 mm o.d.) platinum capsules. Boron nitride and
sintered alumina components were used as internal spacers and surrounds for experiments, with experiments
using graphite sealed in platinum capsules. Anhydrous
experiments using graphite capsules only used fired pyrophyllite and alumina spacers, and mullite and alumina
surrounds (runs 4–11, Table 2). The thermocouple
entered the assembly via a pure alumina sheath and was
protected from the graphite or platinum capsule by a
1 mm alumina disc. All experimental components and
261
JOURNAL OF PETROLOGY
VOLUME 41
NUMBER 2
FEBRUARY 2000
Table 2: Experimental run data
No.
Run no.
T (°C)
Time (h)
% BON∗
MLi/MLf
Phase assemblage
1·5 GPa anhydrous
1
T-3569
1480
1·5
62
1·10
ol+opx+l
2
T-3572
1500
1·0
59
0·92
ol+opx+l
3
T-3587
1530
0·2
100
n.a.
l
Harzburgite melting experiments
4
T-4347
1550
23·5
0
n.a.
ol+opx+l
5
T-4348
1600
10·0
0
n.a.
ol+opx+l
Harzburgite reaction experiments
6
T-4350
1510
69·0
19
1·16
ol+opx+l
7
T-4349
1550
25·5
20
0·86
ol+opx+l
8
T-4351
1600
20·3
17
0·54
ol+opx+l
Tinaquillo melting experiments
9
T-4302
1400
47·0
0
n.a.
ol+opx+cpx+sp+l
10
T-4316
1450
48·3
0
n.a.
ol+opx+l
11
T-4312
1500
24·2
0
n.a.
ol+opx+l
1·5 GPa H2O-undersaturated
12
T-3486
1350
4·0
61
1·27
ol+opx+l
13
T-3471
1400
3·0
45
1·00
ol+opx+l
14
T-3462
1450
2·0
80
0·95
ol+l
15
T-3485
1450
2·0
35
0·90
ol+opx+l
16
T-3581
1470
0·2
100
n.a.
l
17
T-3465
1500
3·0
86
0·98
ol+l
18
T-3533
1500
0·2
100
n.a.
l
ol+opx+l
2·0 GPa anhydrous
19
T-3520
1500
1·0
79
1·27
20
T-3521
1530
0·6
80
1·11
ol+opx+l
21
T-3538
1560
0·3
66
0·86
ol+opx+l
22
T-3564
1560
0·3
61
0·91
ol+opx+l
23
T-3565
1600
0·2
33
0·72
ol+opx+l
2·0 GPa H2O-undersaturated
24
T-3490
1400
3·0
77
1·30
ol+opx+l
25
T-3461
1500
1·0
79
0·91
ol+l
26
T-3464
1500
1·0
59
0·76
ol+l
27
T-3494
1500
1·0
29
0·70
ol+opx+l
28
T-3385
1450
0·5
56
1·40
ol+opx+l
29
T-3387
1450
1·0
44
1·33
ol+opx+l
30
T-3488
1450
2·0
80
0·96
ol+opx+l
Time series
2·5 GPa H2O-undersaturated
31
T-3513
1450
2·0
80
n.a.
ol+opx+l∗
32
T-3516
1480
1·4
80
1·14
ol+opx+l
33
T-3514
1500
1·0
78
0·97
ol+opx+l
34
T-3519
1530
0·6
47
n.a.
ol+opx+l∗
% BON, the wt % BON composition (Table 1) in the bulk composition, except in the case of runs 6–8, which used the glass
composition from run 4 (T-4347, Table 1). MLi/MLf, ratio of the initial mass of liquid in the bulk composition (given by %
BON) and the final mass of liquid in the bulk composition (given by mass balance, Table 3). ol, olivine; opx, orthopyroxene;
l, glass. n.a., not applicable.
∗Runs (31 and 34) where the glass composition could not be determined because of poor quenching.
262
FALLOON AND DANYUSHEVSKY
MELTING OF REFRACTORY MANTLE
starting materials were stored in a oven at 110°C. Experiments were performed using the hot piston-out technique ( Johannes et al., 1971). Pressures are accurate to
within ±0·1 GPa.
starting material; (3) the amount of hydrogen infused
into the sample during the experiment.
The first factor above is of particular relevance to our
study, as part of the BON starting material is a natural
glass (A-46, Table 1) and consequently there will be an
initial, but undetermined, amount of Fe2O3 in our BON
starting composition. CO2 is produced via reaction between the graphite capsule and Fe2O3 in the melt according to the reaction
Analytical techniques
FTIR spectroscopy
Infrared spectroscopy was used to monitor any possible
significant changes in H2O contents during the experiments. Measurements were performed using a Bruker
IFS 66 spectrometer with attached optical microscope
(all reflecting optics) and Bruker Opus/IR reduction
software, housed in the Central Science Laboratory
(CSL), UT. Run products were doubly polished
(30–70 lm thick, using Superglue for bonding to a standard thin-section glass slide, during polishing). Diameters
of analysed areas were usually 60–90 lm. During each
analysis, 100 scans were collected with the resolution of
four wavenumbers between 4000 cm−1 and 2400 cm−1.
H2O contents were estimated using the main OH-stretching peak at >3500 cm−1, following the calibration of
Danyushevsky et al. (1993).
Unfortunately, the above sample preparation technique
results in poor analytical accuracy (mainly because of
difficulties in estimating sample thickness and the necessity to subtract the glue spectrum). Additional problems during analyses of H2O-bearing runs result from:
(1) the unknown density of the quenched material; (2)
the unknown, and variable proportions of glass and
quenched crystals (Fig. 1); (3) changes in the 3500 cm−1
peak shape caused by the presence of quenched crystals
in the analysed area.
The analysed H2O contents in the hydrous experiments
presented in this paper were usually within ±0·5 wt %
from the estimates made on the basis of mass balance
and initial H2O contents. As a result of the above technical
problems, we consider the mass balance estimates to be
more accurate, and we used the IR data only to ensure
that no significant changes in H2O content occurred
during the experiments.
In the case of the nominally anhydrous experiments,
H2O analyses using FTIR spectroscopy were more precise
because of the absence of quench crystallization (Fig.
1a), and we therefore consider the analysed H2O contents
for nominally anhydrous experiments to be reliable.
As our experiments were performed in graphite capsules, we also need to consider the possibility of dissolved
CO2 in our experimental glasses. Holloway et al. (1992)
demonstrated that the dissolved CO2 concentration in
experimental basaltic melts, run in graphite capsules,
depends on three important factors: (1) the concentrations
of ferric and ferrous iron in the starting material; (2) the
amount of any additional CO2 source added to the
Cgraphite(capsule) + 2Fe2O3(melt) + O2−(melt) =
(1)
CO32−(melt) + 4FeO(melt).
If we assume that reaction (1) proceeds to completion
in our experiments, then it is possible for us to calculate
the maximum amount of CO2 we can expect to analyse
by FTIR spectroscopy. We have performed this calculation for runs T-3464 and T-3462 (Table 2). If the
natural glass A-46 equilibrated at oxygen fugacities equivalent to the NNO (nickel–nickel oxide) oxygen buffer
then CO2 contents should vary between 0·15 (T-3464)
and 0·26 (T-3462) wt %. If the natural glass A-46
equilibrated at oxygen fugacities equivalent to the QFM
(quartz–fayalite–magnetite) oxygen buffer then CO2 contents should vary from 0·10 (T-3462) to 0·082 (T-3464)
wt %. For the purposes of analysing the quenched melt
phase in runs T-3464 and T-3462 for dissolved CO2 by
FTIR spectroscopy, the run products were unmounted
from their glass slides by dissolving the Superglue in
acetone. The very small thickness of the samples resulted
in relatively high detection limits for CO2 components
in the melt (>0·07 wt %). However, no CO2 was detected
in either run above the detection limit. This suggests that
no petrologically significant amounts of CO2 was present
in our run products.
Another potential problem of significance for our experiments is hydrogen diffusion out of our experimental
capsules as a result of reaction of H2O and the graphite
capsule. This has the potential to produce significant
amounts of CO2 contents in our experimental glasses via
the reaction
2H2O(melt) + Cgraphite(capsule) =
CO2(melt) + 2H2(diffusive loss).
(2)
For example, if we had 0·5 wt % H2O involved in
reaction (2), then we would expect 0·61 wt % CO2 in
our experimental glasses. As we have checked for CO2
using FTIR spectroscopy with very good detection limits
(see above), we consider that reaction (2) has not produced
any petrologically significant amounts of CO2 in our
experimental glasses. The effectiveness of reaction (2) in
producing CO2 in experimental glasses will depend on
the magnitude of the chemical potential gradient for H2
between the capsule and the furnace assembly. For
example, Green (1973a) and Brey & Green (1975) performed a limited number of H2O-saturated experiments
263
JOURNAL OF PETROLOGY
VOLUME 41
in the presence of graphite using an Ol melilitite and
basanite composition, respectively. These experiments
did not come to equilibrium and significant amounts of
CO2 were found to be dissolved in the melt (D. H. Green,
personal communication, 1999). This result can be simply
explained by the presence of a large amount of added
water (>20 wt %) creating a relatively high chemical
potential for H2 inside the capsule relative to the furnace
assembly. In our experiments, we do not have fluidsaturated conditions, having only 1–2 wt % H2O present,
and therefore the chemical potential gradient for H2
between the capsule and the furnace assembly appears
to be insignificant, as demonstrated by measured CO2
contents below or at detection limits in runs T-3464 and
T-3462 (see above).
Electron microprobe microanalysis
Compositions of glasses were analysed using a Cameca
SX50 electron microprobe, housed in the CSL, UT,
at 15 kV and 20 gA, using international standard
USNM 111240/2 (basaltic glass) from Jarosewich et al.
(1980). Counting times for all elements were 10 s for
the peak and 5 s for the background on both sides
of the peak. Glasses were analysed in scanning mode
with area scans varying from 10 to 50 lm2. Ol and
Opx analyses were obtained using a beam size of
1–2 lm and international standards USNM 122142
(augite) and USNM 111312/444 (Ol) ( Jarosewich et
al., 1980). The mineral phases (except Ol) in run
T-4302 (Table 2) were obtained by energy dispersive
microanalysis using a Cameca MICROBEAM microprobe housed in the Research School of Earth Sciences,
The Australian National University (operating conditions 15 kV, 5 gA).
As glasses in our H2O-bearing experiments quenched
to a mixture of glass and quench pyroxenes (which
grew directly out of the glass, Fig. 1b) it was necessary
to systematically investigate the homogeneity of broad
beam area scans across the entire melt layer. We
systematically performed area scans in >3–4 different
areas, and within each area we analysed progressively
smaller areas (from 50 lm × 50 lm to 10 lm ×
10 lm). In the majority of cases, we found that major
elements remained constant, except for Na2O loss on
smaller area scans. These results combined with good
mass balance (Table 3) indicated that the glass composition before the quenching of our H2O-bearing runs
could be determined with confidence.
However, in a few of our high-P H2O-bearing runs
(T-3513, T-3519, Table 2) analyses of quench melt did
not result in good mass balance, and therefore only Ol
and Opx analyses are presented for these experiments in
Table 3.
NUMBER 2
FEBRUARY 2000
EXPERIMENTAL RESULTS
Attainment of equilibrium
Several lines of evidence can be used to evaluate the
approach to equilibrium of our experiments. These are
as follows:
(1) the maintenance of constant sample bulk composition is essential for equilibration, and has been demonstrated by materials balance (Table 3). The average
square of the sum of the residuals for the 31 experiments
reported here is 0·2 ± 0·2 (range 0·0009–0·6636), which
compares favourably with other peridotite melting studies
(e.g. Kinzler & Grove, 1992, 0·01–3·58; Falloon et al.,
1997, 0·004–0·654; Kinzler, 1997, 0·01–1·15; Gaetani
& Grove, 1998, 0·02–0·13).
(2) The achievement of consistent mineral–melt exchange equilibrium indicates a close approach to equilibrium. The average KD Fe–Mg values for Ol (0·34 ±
0·01) and Opx (0·31 ± 0·01) for our short runs (0·2–4
h) are all within 1r of our longer (10–69 h) experiments
(Ol, 0·32 ± 0·02; Opx 0·31 ± 0·02) and compare
favourably with average values reported by other workers
(e.g. Gaetani & Grove, 1998, 0·34 ± 0·01 and 0·32 ±
0·02; Kinzler, 1997, 0·33 ± 0·02 and 0·33 ± 0·04;
Robinson et al., 1998, 0·32 ± 0·02 and 0·32 ± 0·02;
for Ol and Opx, respectively).
(3) The phase compositions are homogeneous. Unlike
natural mineral mix starting materials, synthetic starting
materials made out of sintered oxides react rapidly to
produce homogeneous run products (Falloon et al.,
1999b). This is demonstrated by the very low standard
deviations on averages of both core and rims of Ol and
Opx analyses presented in Table 3 (in general, each
average Ol and Opx composition present in Table 3
represents equal numbers of core and rim analyses). All
standard deviations for Si, Fe, Mg (Ol ) and Si, Fe, Mg,
Al, Ca (Opx) are within one decimal place (Table 3).
(4) The internal consistency of our experimental data
also indicates a close approach to equilibrium. Ol, Opx
and coexisting melt compositions in our experiments all
systematically change with P and T. For example, the
mg-number of both Ol and Opx increases with T, and the
Al2O3 and CaO contents of Opx decrease with increasing
T. The ratio of the initial mass of added liquid (MLi) to
the final mass of liquid (MLf ) shows an internally consistent and systematic change with T at a given P (Table
2). At T below the liquidus of the BON composition the
mass of liquid decreases (MLi/MLf >1), indicating that
crystallization and re-equilibration of the bulk composition is occurring. At T above the liquidus of the
BON composition, the mass of liquid increases (MLi/
MLf <1), indicating that melting and re-equilibration of
the bulk composition is occurring.
(5) We have also performed time series experiments
at 1·5 GPa at 1450°C demonstrating that equilibrium is
264
Phase
0·64(4)
0·6636
Glass
Residual
Glass
0·09(5)
0·0629
Residual
0·28(2)
0·562(9)
Glass
Orthopyroxene
0·18(1)
Orthopyroxene
Olivine
0·262(7)
Olivine
MB
265
0·67(1)
0·21(1)
0·12(1)
0·0472
Orthopyroxene
Glass
Residual
0·0904
Olivine
0·06(1)
Residual
0·26(2)
Orthopyroxene
Glass
0·67(1)
Olivine
T-4351
T-4349
T-4350
0·149(9)
0·313(8)
0·0218
Orthopyroxene
Glass
Residual
0·0496
Residual
0·536(4)
0·23(1)
Glass
Olivine
0·24(1)
0·0778
Residual
0·53(1)
0·16(1)
Glass
Orthopyroxene
0·30(1)
Orthopyroxene
Olivine
0·54(1)
Olivine
Harzburgite reaction experiments
T-4348
T-4347
Harzburgite melting experiments
T-3587
T-3572
T-3569
1·5 GPa anhydrous
Run no.
51·9(5)
6
41·5(3)
58·4(2)
53·9(6)
7
53·4(3)
2
5
57·7(4)
4
41·3(2)
52·43
Cal
4
57·2(1)
4
41·0(2)
54·47
Sel
5
58·3(1)
5
4
41·7(4)
5
53·82
58·0(7)
5
Sel
41·6(4)
5
51·2(3)
57·6(3)
6
40·9(2)
51·7(5)
10
2
57·0(2)
18
41·1(1)
12
SiO2
18
n
0·88(7)
0·07(3)
1·27(3)
0·07(3)
1·50
0·10(2)
0·72
0·04(3)
1·07
0·06(3)
0·73(3)
0·71(3)
0·04(2)
0·81(4)
0·07(4)
TiO2
4·6(4)
0·37(3)
6·20(9)
0·62(1)
7·26
0·9(1)
4·22
0·29(5)
5·97
0·52(6)
8·5(2)
8·27(6)
0·9(1)
9·5(2)
1·4(2)
Al2O3
Table 3: Compositions of experimental run products
7·4(2)
3·33(9)
4·9(1)
7·5(3)
3·7(2)
5·6(2)
7·16
3·9(2)
5·7(1)
7·86
3·3(1)
5·0(1)
7·89
3·4(1)
5·5(2)
9·3(2)
8·4(1)
4·2(1)
6·7(2)
8·5(2)
4·8(2)
7·4(2)
FeO
0·28(3)
0·15(5)
0·15(2)
0·30(5)
0·15(3)
0·13(2)
0·28
0·16(3)
0·17(6)
0·29
0·11(1)
0·14(1)
0·27
0·13(1)
0·15(2)
0·13(3)
0·12(7)
0·1(1)
0·07(4)
0·15(6)
0·08(4)
0·08(5)
MnO
26·3(7)
36·85(7)
53·0(4)
22·4(4)
35·8(3)
52·1(3)
20·38
35·0(1)
51·5(2)
26·63
36·9(1)
52·6(5)
22·50
36·5(6)
52·2(5)
21·1(2)
21·7(3)
35·7(3)
51·9(2)
19·4(4)
35·3(3)
51·1(2)
MgO
5·1(3)
0·59(5)
0·12(1)
7·2(3)
0·9(1)
0·18(1)
9·11
1·16(5)
0·21(1)
4·25
0·47(1)
0·09(2)
6·48
0·6(1)
0·12(2)
7·0(1)
6·74(7)
0·89(0)
0·18(2)
7·6(3)
1·07(8)
0·22(3)
CaO
0·35(8)
0·03(1)
0·50(4)
0·03(1)
0·48
0·05(1)
0·44
0·03(1)
0·79
0·04(1)
1·6(2)
1·5(3)
0·05(1)
1·7(3)
0·06(1)
Na2O
0·18(5)
0·33(2)
0·40
0·25
0·37
0·30(2)
0·31(4)
0·31(3)
K 2O
0·05(3)
0·07(2)
0·08(3)
P 2O5
0·91(9)
0·54(5)
0·46(3)
0·95(6)
0·67(6)
0·48(6)
0·91
0·8(1)
0·48(4)
0·89
0·57(8)
0·43(6)
0·86
0·64(6)
0·43(7)
0·03(2)
0·18(3)
0·4(1)
0·17(8)
0·20(4)
0·2(1)
0·10(4)
Cr2O3
0·01(2)
0·02(3)
0·05(4)
0·01(2)
0·01(1)
0·01(2)
0·1
0·02(5)
0·03(5)
0·01(2)
0·02(2)
0·01(2)
0·03(4)
NiO
n.d.
n.d.
n.d.
n.d.
n.d.
n.d.
0·5
0·7
H2O∗
100·3(8)
100·4(2)
100·3(7)
99·7(6)
99·6(8)
100·0(5)
99·92
99·2(2)
99·3(3)
99·4
99·9(3)
99·3(9)
97·4
100(1)
99·8(8)
101·2(5)
100·7(6)
101·3(4)
100·6(7)
100·1(6)
101·5(3)
99·4(4)
Probe
FALLOON AND DANYUSHEVSKY
MELTING OF REFRACTORY MANTLE
Phase
MB
0·37(1)
0·13(1)
0·50(1)
0·0700
Orthopyroxene
Glass
Residual
0·1132
Olivine
0·37(1)
Residual
0·0009
Residual
Glass
0·247(2)
Glass
0·25(2)
0·004(1)
Spinel
0·37(1)
0·080(3)
Clinopyroxene
Orthopyroxene
0·295(2)
Orthopyroxene
Olivine
0·370(1)
Olivine
266
0·2554
Residual
Glass
0·88(1)
Glass
T-3533
0·126(1)
Olivine
0·0379
Residual
Glass
0·395(8)
Glass
T-3465
0·17(1)
Orthopyroxene
0·5601
0·440(5)
Residual
Olivine
0·84(1)
T-3581
T-3485
0·165(9)
Glass
0·0834
Residual
Olivine
0·45(1)
Glass
51·9(4)
12
51·3(2)
40·7(2)
8
51·1(2)
53·1(4)
15
8
57·2(4)
18
40·8(2)
14
52·1(2)
13
41·1(1)
53·4(4)
10
8
57·3(3)
15
41·30(7)
52·5(5)
12
5
56·7(3)
10
41·3(2)
12
50·3(2)
10
13
56·7(1)
4
0·77(2)
0·73(3)
0·72(5)
0·74(4)
0·07(2)
0·69(3)
0·77(3)
0·08(3)
0·95(5)
0·10(2)
0·31(4)
0·03(2)
0·46(2)
0·05(2)
0·63(3)
0·15(4)
TiO2
8·8(1)
8·8(3)
8·79(8)
8·3(2)
0·9(4)
8·60(6)
9·5(2)
1·0(3)
11·0(4)
1·6(1)
9·9(1)
1·74(3)
12·4(1)
2·8(4)
15·50(5)
48(2)
6·3(3)
2·8(2)
Al2O3
9·6(1)
8·9(6)
7·5(1)
9·4(1)
7·6(1)
4·0(4)
6·3(3)
8·8(1)
7·2(2)
7·5(2)
4·2(5)
6·66(4)
8·1(2)
5·6(4)
8·1(1)
8·4(1)
4·54(8)
7·14(9)
8·12(7)
5·0(2)
8·0(2)
8·4(1)
9·2(2)
4·61(6)
5·4(4)
9·30(9)
FeO
0·10(5)
0·08(4)
0·08(5)
0·08(4)
0·14(6)
0·09(5)
0·12(4)
0·10(5)
0·07(4)
0·13(5)
0·09(2)
0·08(4)
0·14(5)
0·08(5)
0·12(3)
0·20(5)
0·16(5)
0·15(3)
0·20(4)
0·15(3)
0·14(2)
0·22(2)
0·18(3)
MnO
20·5(3)
20·4(6)
50·9(2)
20·4(2)
21·2(4)
36·1(6)
52·3(2)
21·0(1)
51·4(2)
19·2(3)
35·8(6)
51·62(4)
16·1(9)
34·2(5)
50·1(2)
19·8(4)
34·1(1)
50·4(3)
16·3(1)
32·9(5)
49·9(3)
13·46(7)
20·1(3)
21·6(3)
32·2(2)
49·1(6)
MgO
7·2(1)
7·1(2)
0·16(1)
7·1(1)
6·6(1)
0·8(2)
0·15(3)
7·01(7)
0·17(2)
7·0(1)
0·94(2)
0·15(1)
8·6(3)
1·4(1)
0·20(2)
9·7(2)
1·36(4)
0·30(2)
12·0(1)
2·23(7)
0·33(2)
12·5(1)
0·21(8)
13·8(4)
2·20(2)
0·33(4)
CaO
1·4(2)
1·59(5)
1·65(3)
1·5(3)
0·05(1)
1·09(2)
1·7(3)
0·06(2)
1·8(3)
0·05(2)
0·13(5)
0·56(7)
0·04(0)
0·74(3)
0·05(1)
1·15(3)
0·3(1)
Na2O
0·31(4)
0·34(7)
0·29(2)
0·32(3)
0·38(3)
0·33(2)
0·46(4)
0·08(1)
0·13(2)
0·21(1)
K 2O
0·06(4)
0·05(2)
0·07(3)
0·06(3)
0·03(2)
0·05(3)
0·06(2)
0·05(3)
0·01(1)
0·02(2)
P 2O5
0·01(2)
0·10(2)
0·02(3)
0·04(3)
0·38(4)
0·7(2)
0·3(1)
0·10(4)
0·07(2)
0·32(4)
0·5(2)
0·19(2)
0·19(4)
0·2(1)
0·13(5)
0·81(7)
1·03(6)
0·43(6)
0·6(1)
1·2(2)
0·4(2)
0·21(6)
22(2)
1·50(4)
1·50(8)
0·16(5)
Cr2O3
0·02(3)
0·01(2)
0·02(2)
0·10(4)
0·19(3)
0·40(5)
0·12(6)
0·21(6)
0·44(6)
0·09(3)
0·45(5)
NiO
2·4
2·3
2·4
2·1
2·3
2·4
3·0
n.d.
n.d.
n.d.
H2O∗
98·0(4)
97(1)
99·3(4)
99·5(3)
97·8(6)
100·3(4)
99·7(3)
98·1(4)
99·7(7)
97·4(7)
100·8(4)
99·3(1)
94(1)
100·7(5)
98·7(4)
98·5(3)
99·9(2)
99·4(5)
100·9(5)
100·5(5)
100·3(4)
99·5(3)
102·7(5)
102·0(8)
100(1)
100·2(9)
Probe
NUMBER 2
T-3462
0·14(1)
0·1383
Residual
0·41(1)
0·48(1)
Glass
Orthopyroxene
0·18(2)
Orthopyroxene
Olivine
0·34(1)
Olivine
40·7(1)
9
49·0(1)
3
3
55·8(6)
0·3(3)
47·8(2)
4
41·0(4)
51·9(3)
4
6
55·9(4)
7
40·6(3)
4
SiO2
7
n
VOLUME 41
T-3471
T-3486
1·5 GPa H2O-undersaturated
T-4312
T-4316
T-4302
Tinaquillo melting experiments
Run no.
Table 3: continued
JOURNAL OF PETROLOGY
FEBRUARY 2000
Phase
267
0·407(5)
0·14(1)
0·46(1)
0·0467
Olivine
Orthopyroxene
Glass
Residual
0·1337
0·67(1)
Glass
Residual
0·11(2)
0·
Residual
Orthopyroxene
0·764(8)
Glass
0·22(1)
0·100(9)
Olivine
0·135(4)
0·1053
Residual
Orthopyroxene
0·72(1)
Glass
Olivine
0·17(2)
0·1272
Residual
0·11(1)
0·62(1)
Glass
Orthopyroxene
0·27(2)
Orthopyroxene
Olivine
0·11(1)
Olivine
MB
T-3494
T-3464
T-3461
T-3490
0·18(3)
0·41(3)
0·3662
Orthopyroxene
Glass
Residual
0·0274
0·41(1)
Residual
Olivine
0·774(2)
Glass
0·3535
0·229(2)
Residual
Olivine
0·868(8)
0·1801
Residual
0·141(7)
0·59(1)
Glass
Glass
0·24(2)
Orthopyroxene
Olivine
0·16(1)
Olivine
2·0 GPa H2O-undersaturated
T-3565
T-3564
T-3538
T-3521
T-3520
2·0 GPa anhydrous
Run no.
41·2(1)
57·3(2)
51·3(7)
6
8
51·9(4)
15
41·4(2)
8
51·1(3)
18
41·3(1)
51·1(5)
12
8
56·7(2)
24
40·7(2)
51·4(4)
9
14
57·8(2)
11
41·5(3)
51·4(2)
10
5
57·2(2)
5
41·4(2)
50·7(4)
10
3
57·4(3)
14
41·2(1)
50·9(4)
8
5
56·7(3)
14
40·9(2)
50·2(2)
9
17
56·4(2)
12
41·0(2)
14
SiO2
18
n
0·59(3)
0·05(3)
0·63(3)
0·70(3)
0·88(5)
0·09(3)
0·61(6)
0·06(5)
0·65(6)
0·10(4)
0·66(3)
0·05(2)
0·81(6)
0·10(3)
0·87(5)
0·10(2)
TiO2
6·1(3)
1·1(2)
7·1(2)
8·3(3)
11·1(2)
1·7(2)
7·0(2)
0·6(2)
7·8(2)
1·1(1)
7·6(2)
1·1(2)
9·5(2)
2·2(3)
9·9(2)
2·5(3)
Al2O3
8·1(1)
3·7(3)
6·0(1)
7·7(3)
5·6(2)
8·6(2)
6·6(2)
8·9(3)
5·6(4)
8·5(2)
7·9(3)
3·8(4)
5·4(1)
8·7(3)
3·9(3)
6·2(2)
8·4(2)
3·8(4)
6·1(1)
8·8(1)
5·0(3)
8·1(1)
9·5(2)
5·2(6)
8·5(1)
FeO
0·15(6)
0·10(5)
0·1(1)
0·12(6)
0·07(6)
0·09(6)
0·06(5)
0·12(6)
0·08(5)
0·09(6)
0·15(8)
0·10(4)
0·07(4)
0·09(5)
0·08(4)
0·08(5)
0·11(5)
0·08(5)
0·07(5)
0·12(6)
0·05(3)
0·06(5)
0·10(4)
0·06(3)
0·09(4)
MnO
26·8(7)
36·2(3)
52·1(1)
24·9(6)
52·6(2)
22·6(6)
51·8(2)
17·2(5)
34·0(3)
50·4(2)
25·4(4)
36·4(6)
52·3(2)
23·1(3)
35·9(3)
52·0(3)
24·6(7)
36·1(2)
52·3(2)
20·2(6)
34·1(2)
50·6(2)
18·7(3)
33·7(5)
50·04(2)
MgO
5·0(4)
0·84(5)
0·16(2)
5·7(4)
0·14(3)
6·6(1)
0·16(2)
8·4(2)
1·4(1)
0·22(2)
5·4(2)
0·8(2)
0·15(3)
6·4(2)
1·0(2)
0·17(2)
6·2(2)
0·9(1)
0·18(20
7·7(1)
1·4(1)
0·24(3)
8·6(2)
1·6(2)
0·24(2)
CaO
1·2(2)
1·35(4)
1·5(1)
1·7(2)
0·07(3)
1·5(4)
0·08(2)
1·40
0·12(2)
1·3(5)
0·09(2)
1·5(3)
0·11(2)
1·7(3)
0·12(2)
Na2O
0·22(3)
0·25(3)
0·31(3)
0·39(5)
0·30(6)
0·29(2)
0·24(7)
0·33(5)
0·32(3)
K 2O
0·04(3)
0·02(3)
0·06(3)
0·06(2)
0·03(2)
0·05(3)
0·02(2)
0·07(3)
0·07(3)
0·12(4)
P 2O5
0·48(3)
0·6(2)
0·4(1)
0·25(3)
0·15(2)
0·14(4)
0·07(2)
0·11(4)
0·2(2)
0·09(4)
0·25(7)
0·5(2)
0·5(5)
0·16(6)
0·52(4)
0·18(2)
0·15(4)
0·4(2)
0·18(6)
0·11(4)
0·2(2)
0·10(3)
0·10(5)
0·3(3)
0·01(2)
Cr2O3
0·02(2)
0·01(2)
0·01(2)
0·01(2)
0·03(3)
0·01(1)
0·01(1)
0·02(2)
NiO
1·8
1·8
2·1
3·1
0·3
0·3
0·2
n.d.
n.d.
H2O∗
98·8(7)
100·3(3)
100·5(5)
97·0(7)
99·5(3)
97·5(4)
99·2(5)
97·9(9)
100·7(5)
100·0(4)
99·7(4)
100·9(4)
98·9(6)
99·4(4)
99·6(7)
98·2(4)
98·7(4)
99·9(4)
100·0(2)
100·0(5)
100·7(4)
99·6(6)
99·2(5)
101·3(5)
99·1(3)
Probe
FALLOON AND DANYUSHEVSKY
MELTING OF REFRACTORY MANTLE
Phase
0·106(4)
0·066(9)
0·828(7)
0·319
Glass
Residual
0·6360
Residual
Orthopyroxene
0·33(3)
Glass
Olivine
0·29(4)
0·4475
Residual
0·38(2)
0·40(3)
Glass
Orthopyroxene
0·32(4)
Orthopyroxene
Olivine
0·28(2)
Olivine
MB
268
57·2(4)
n.d.
Glass†
0·06(2)
0·68(4)
0·05(2)
0·77(3)
0·07(1)
1·3(3)
8·7(5)
1·4(2)
9·8(2)
2·1(3)
2·4(2)
8·4(2)
1·0(1)
10·3(9)
2·0(3)
10·5(8)
2·6(4)
Al2O3
3·9(4)
5·8(1)
8·9(3)
4·3(1)
6·7(1)
9·2(1)
4·9(2)
7·8(2)
5·3(4)
8·5(2)
8·7(2)
4·4(2)
7·1(2)
8·2(2)
5·2(8)
8·2(2)
8·5(4)
5·9(3)
8·6(4)
FeO
0·05(4)
0·05(2)
0·12(4)
0·03(2)
0·06(5)
0·13(5)
0·06(4)
0·06(5)
0·09(3)
0·08(6)
0·09(4)
0·07(4)
0·07(5)
0·19(5)
0·09(7)
0·04(8)
0·12(2)
0·11(4)
0·10(4)
MnO
36·1(6)
52·5(2)
23·0(5)
35·7(4)
51·8(2)
20·6(3)
34·6(5)
50·8(2)
34·2(4)
50·2(4)
21·8(6)
35·7(3)
51·5(3)
19(2)
33·9(7)
49·8(6)
19(2)
32·6(8)
49·9(6)
MgO
0·9(3)
0·14(2)
6·6(1)
0·87(9)
0·15(2)
7·58(8)
1·2(2)
0·17(2)
1·38(6)
0·19(3)
6·8(2)
0·9(1)
0·18(3)
7·8(5)
1·6(3)
0·3(1)
8·3(5)
2·1(6)
0·28(9)
CaO
0·08(4)
1·40(6)
0·07(1)
0·15(2)
1·57(9)
0·11(2)
0·14(2)
1·6(2)
0·05(2)
1·6(3)
0·12(1)
1·60(5)
0·15(4)
Na2O
0·26(3)
0·31(4)
0·27(3)
0·43(8)
0·01(1)
0·43(6)
K2 O
0·06(4)
0·06(3)
0·04(3)
0·09(2)
0·07(3)
P 2O5
0·3(1)
0·18(2)
0·12(4)
0·14(4)
0·09(5)
0·12(4)
0·2(1)
0·06(3)
0·3(2)
0·11(4)
0·13(4)
0·3(1)
0·09(3)
0·18(3)
0·5(3)
0·2(7)
0·11(4)
0·3(2)
0·2(1)
Cr2O3
0·03(2)
0·01(1)
0·02(2)
0·02(1)
0·01(2)
0·02(2)
0·03(6)
0·02(2)
NiO
2·3
2·7
2·3
3·2
3·4
H2O∗
100·6(3)
99·2(2)
98·1(9)
100·7(3)
99·6(4)
98·1(4)
100·2(5)
99·2(5)
101·2(4)
100·1(6)
98·0(7)
100·8(4)
99·8(5)
97(1)
100·1(5)
100·1(7)
98·5(8)
102·3(6)
99·2(5)
Probe
Numbers in parentheses next to each analysis or mass balance are 1r in terms of the last units cited; e.g. 0·540(9) refers to 0·540±0·009. Mass balance ( MB) in
weight fraction was performed using least-squares linear regression using the software PETMIX and ‘Residual’ refers to the square of the sum of the residuals.
All analyses have been normalized to 100 wt % before averages have been calculated (number of analyses used in each average is given under n), the average
unnormalized totals are given under ‘Probe’. n.d., not determined; Sel, a selected analysis; Cal, a calculated analysis (see text for discussion).
∗H2O content as determined by FTIR spectroscopy in the case of nominally anhydrous runs and calculated content from mass balance, assuming no H2O loss
during the course of the experiment.
†No reliable analysis possible because of quench modification.
Olivine
41·2(1)
50·1(7)
12
6
0·1868
Residual
57·5(2)
12
0·80(2)
Glass
41·2(2)
9
49·8(2)
8
17
56·7(4)
22
Orthopyroxene
0·15(2)
41·1(3)
12
0·08(2)
0·73(4)
0·06(2)
0·88(8)
0·10(4)
0·88(9)
0·12(3)
TiO2
NUMBER 2
T-3519
0·05(1)
0·1892
Residual
Orthopyroxene
0·70(2)
Glass
Olivine
0·23(2)
Orthopyroxene
n.d.
Glass†
0·07(1)
56·1(2)
13
Olivine
40·9(3)
18
51·5(5)
16
Orthopyroxene
57·5(3)
Olivine
41·1(2)
51·0(9)
8
9
56·4(5)
17
41·4(4)
50·6(8)
7
8
56·1(6)
6
40·9(3)
20
SiO2
18
n
VOLUME 41
T-3514
T-3516
T-3513
2·5 GPa H2O-undersaturated
T-3488
T-3387
T-3385
Time series
Run no.
Table 3: continued
JOURNAL OF PETROLOGY
FEBRUARY 2000
FALLOON AND DANYUSHEVSKY
MELTING OF REFRACTORY MANTLE
achieved within 2 h (Table 2). Runs T-3385, T-3387
and T-3488 show systematic changes in melt and mineral
compositions, and MLi/MLf with time (Tables 2 and 3).
Most importantly, the experiments show that within 2 h
the Ol KD Fe–Mg decreases from 0·38 to an equilibrium
value of 0·34.
In summary, we believe the experimental data presented here closely approximate equilibrium phase assemblages.
In Fig. 1a–c we present photographs of our run products to demonstrate the difference in quenching between
‘anhydrous’ (Fig. 1a) and hydrous experiments (Fig. 1b
and c), and to illustrate that our hydrous runs contain a
significant amount of disseminated graphite throughout
the quenched melt layer (Fig. 1b and c). The layer of
graphite was apparent only after the experimental run
products had been prepared as double-sided polished
thin-sections for FTIR analysis (compare Fig. 1b with
Fig. 1c). The broad bands of disseminated graphite
appear to correlate with electron microprobe area scans
that showed anomalously low Na2O contents. Although
we do not have an explanation for the observed phenomenon, the graphite is unlikely to have quenched from
the melt, as we do not observe any CO32− peaks in our
FTIR spectrum. We therefore believe that the graphite is
mechanically derived from the enclosing graphite capsule,
the process being linked to, or enhanced by, the presence
of H2O.
‘Anhydrous’ peridotite reaction
experiments
Melting of refractory lherzolite and harzburgite at 1·5 GPa
The results of our melting experiments at 1·5 GPa on
the refractory compositions TQ-40 and HZ are presented
in Fig. 2a and b, and compared with melting cotectics
for more fertile lherzolite compositions MPY (Robinson
et al., 1998) and MM-3 (Falloon et al., 1999b). Direct
melting experiments on TQ-40 at 1400, 1450 and 1500°C
(runs 9–11, Table 2) resulted in well-equilibrated run
products in which there was a significant melt fraction
(F = 25–50 wt %, Table 3). It was possible to obtain
consistent glass analyses which appear to be free from
the effects of quench modification, as a result of the
presence of relatively large glass pools. We were able to
obtain very good mass balances with very low residual
sums (0·0009–0·1132, Table 3). Direct melting experiments on HZ at 1550 and 1600°C (runs 4 and 5,
Table 2) resulted in well-equilibrated run products in
which there was a low F (>6–12 wt %, Table 3).
Although it was only possible to analyse glass in small
pools, selected area scans gave good mass balance and
equilibrium KD Fe–Mg for Ol and Opx (Table 3). To
confirm these compositions as equilibrium melts, reversal
experiments were performed by reacting the composition
of the melt phase in run T-4347 with HZ at 1550 and
1600°C (runs 7 and 8, Table 2). Both runs T-4349 and
T-4351 resulted in well-equilibrated run products in
which there was a significant F (23–31 wt %, Table 3).
It was possible to obtain consistent glass analyses that
appear to be free from the effects of quench modification,
as a result of the presence of relatively large glass pools.
We were able to obtain very good mass balances with
very low residual sums (0·02–0·05, Table 3). The glasses
in the reversal experiments are very close in composition
to those obtained from the direct melting experiments
on the HZ composition at 1550 and 1600°C (Fig. 2a
and b) and we therefore believe that they represent
equilibrium melts of the HZ composition at 1·5 GPa
under ‘nominally’ anhydrous conditions.
Although we have not analysed our dry melting experiments on TQ-40 and HZ for H2O by FTIR spectroscopy, we believe our experiments have negligible
H2O contents for the following reasons: (1) we have used
exactly the same experimental techniques as in the study
of Falloon et al. (1999b), who determined that H2O
contents were between 0·07 and 0·17 wt % for the
experimental glasses chosen for FTIR analysis; (2) calculated anhydrous Ol liquidus T for our experiments
using the Ol geothermometer of Ford et al. (1983) are all
with 15°C of nominal experimental run T (see Fig. 7b,
below). This suggests that our H2O contents are <>0·1
wt % (see discussion below).
Assuming a linear relationship between F and T for
the HZ composition at 1·5 GPa, the solidus (>0% F )
lies between >1490 and 1495°C, and at 1510°C the
HZ composition has >2% F. In run T-4350 (Table 2)
we attempted to determine the composition of a nearsolidus melt for the HZ composition at 1510°C by a
reaction experiment using the T-4347 glass composition
and HZ. Run T-4350 resulted in a well-equilibrated run
product with very small melt pools. Glass area scan
analyses on these pools varied systematically in major
elements because of differing proportions of quench
crystals and glass. Using the Ol geothermometer of Ford
et al. (1983) we calculated an anhydrous Ol liquidus T
for each individual glass scan. Calculated Ol liquidus T
varied from 1445 to 1565°C. The composition of the
melt fraction in run T-4350 was then calculated by linear
regression, in similar manner to Falloon & Green (1988),
using the calculated Ol liquidus T as the independent
variable. The calculated composition of the melt in run
T-4350 produced a very good mass balance with a very
low residual sum (>0·08, Table 3). The position of the
calculated melt composition for run T-4350 is consistent
with the higher-T melt compositions of HZ in Fig. 2a
and b. Using the calculated glass composition of run T4350 and the coexisting residual phases, mass balance
indicates that the glass in run T-4350 represents >2·2%
269
JOURNAL OF PETROLOGY
VOLUME 41
Fig. 2. Comparison of glass compositions from anhydrous 1·5 GPa
melting experiments on refractory lherzolite TQ-40 and refractory
harzburgite (HZ) with glass compositions from melting experiments on
fertile lherzolite compositions MPY (Robinson et al., 1998) and MM-3
(Falloon et al., 1999b) in the molecular normative projection from Ol
(a) onto the face [ Jd + CaTs + Lc]–Di–Qz and from Di (b) onto the
base [ Jd + CaTs + Lc]–Qz–Ol of the ‘basalt tetrahedron’ (Falloon &
Green, 1988; see insert). Ο, glass compositions in direct melting
experiments on HZ; Η, glass compositions in reversal experiments of
HZ melt compositions; Φ, glass in direct melting experiments on TQ40 in equilibrium with Ol + Opx; Ε, glass composition in direct
melting experiment on TQ-40 in equilibrium with Ol + Opx + Cpx;
half-filled diamond in (b) is the composition of the eutectic melt in the
system Fo–En–Qz from the study of Taylor (1973); continuous lines
delineate Ol + Opx ± Sp + L cotectics; dashed lines delineate Ol +
Opx + Cpx ± Sp + L cotectics; dot–dash lines delineate Ol + L
cotectics, with arrows in (b) pointing towards the respective bulk
compositions; dotted lines labelled 1 and 2 in (b) delineate the inferred
movement of the Ol + Opx + L cotectic as sources become more
refractory (see text for discussion); labelled symbols in (a) refer to the
respective peridotite compositions MPY (crossed diamond), MM-3
(crossed circle), TQ-40 (square with diagonal cross) and HZ (crossed
square). Cotectics for MPY and MM-3 are based on the experimental
data of Robinson et al. (1998) and Falloon et al. (1999b), respectively.
F (residual sum of squares is 0·2104). We therefore believe
that the calculated glass composition in run T-4350 is a
close approximation to an equilibrium near-solidus melt
of the HZ composition at 1·5 GPa.
In Fig. 2a and b the position of equilibrium melting
cotectics for the HZ is compared with those for fertile to
refractory lherzolite compositions within the molecular
normative basalt tetrahedron. The model lherzolites
MPY, MM-3 and TQ-40 vary systematically in their
CaO/Al2O3 (0·776, 0·897 and 0·929, respectively) and
NUMBER 2
FEBRUARY 2000
CaO/Na2O (7·22, 11·52 and 16·56, respectively) values,
and therefore the three compositions can be used to
examine the changes in phase relationships for fertile
mantle undergoing a process of isobaric dynamic melting.
Higher CaO/Na2O values of a composition are reflected
in higher Qz/[ Jd + CaTs + Lc] values in Fig. 2b,
whereas higher CaO/Al2O3 values are reflected in higher
Di/[ Jd + CaTs + Lc] values in Fig. 2a. As can be seen
from Fig. 2b, the mantle melting cotectics for the three
lherzolite compositions are almost coincident. However,
the following differences between the melting cotectics
of the three lherzolite compositions can be seen in an
examination of Fig. 2a and b:
(1) as the position of near-solidus initial melts, in the
basalt tetrahedron (Fig. 2a and b), is dependent on the
CaO/Na2O value of the lherzolite composition (Falloon
et al., 1997b), there are significant differences between
the initial melts of the three lherzolite compositions.
Initial melts from the most fertile composition (MPY)
are strongly nepheline (Ne) normative (Fig. 2b), have
relatively low normative diopside (Di) contents in Fig. 2a,
and are relatively high in SiO2 (52·7 wt %) and Na2O
(7·27 wt %) (Robinson et al., 1998). Initial melts from the
slightly more refractory MM-3 composition are still Ne
normative (Fig. 2b), have higher normative Di contents
in Fig. 2a, and have intermediate SiO2 (49·21 wt %) and
Na2O (3·81 wt %) contents (Falloon et al., 1999b). Initial
melts from the refractory TQ-40 composition are inferred
to have lower SiO2 and Na2O contents, to be Ol and
hypersthene (Hy) normative in Fig. 2b, and to have the
highest normative Di content in Fig. 2a.
(2) The point at which Cpx is eliminated from the
residue, as seen in Fig. 2b, moves progressively towards
the Ol–Qz boundary of the tetrahedron, as lherzolite
sources become more refractory. In Fig. 2a the Cpx-out
point moves to higher normative Di as lherzolite sources
become refractory and increase in CaO/Al2O3 value.
(3) The point at which Opx is eliminated from the
residue also moves progressively towards the Ol–Qz
boundary of the tetrahedron in Fig. 2b, as lherzolite
sources become more refractory. In Fig. 2b, it can be
seen that the Ol + Opx + L cotectics for the three
lherzolite compositions are almost coincident; however,
in Fig. 2a, these cotectics are clearly separate, because
of the differing CaO/Al2O3 values of the lherzolite compositions. The more refractory compositions define Ol
+ Opx + L cotectics at progressively higher CaO/Al2O3
values.
Compared with the mantle melting cotectics for the
lherzolite compositions, the Ol + Opx + L cotectic for
the HZ composition plots in a significantly different
position in the basalt tetrahedron, as can be seen from
Fig. 2b. The progressive melting of sources more refractory than TQ-40 may have been expected to produce
liquids in equilibrium with Ol + Opx lying on an extension
270
FALLOON AND DANYUSHEVSKY
MELTING OF REFRACTORY MANTLE
of the Ol + Opx cotectics for lherzolite compositions
towards the simple system eutectic forsterite ( Fo) +
enstatite (En) + L at 1·5 GPa (dotted curve, labelled 1
in Fig. 2b). However, as can be seen from Fig. 2b, melts
from the HZ composition are displaced towards the Qz
apex of the basalt tetrahedron and are Qz normative at
T <1600°C. As well as this displacement, the orientation
of the Ol + Opx + L cotectic within the basalt tetrahedron, as seen in Fig. 2b, is significantly different
compared with the Ol + Opx + L cotectics for lherzolite
sources. These differences in normative composition are
also reflected in significantly different major element
compositions as can be seen in Fig. 3. Melt compositions
from HZ have significantly higher SiO2, lower Al2O3 and
slightly lower CaO at a given MgO content compared
with melts from lherzolite sources (Fig. 3). Therefore
melt compositions from lherzolite sources cannot be
used to infer the compositions of melts from refractory
harzburgite compositions, because, as can be seen in Figs
2 and 3, there is a discontinuity in composition between
melts from lherzolite and harzburgite sources. Liquids in
equilibrium with Ol + Opx from more refractory sources
than HZ are inferred to lie on an extension of the Ol +
Opx + L cotectic for HZ towards the Fo + En + L
eutectic at 1·5 GPa (dotted curve labelled 2 in Fig. 2b).
In Fig. 2a, melt compositions from HZ have high
CaO/Al2O3 values and achieve lower normative Di contents than melts in equilibrium with Ol + Opx from
lherzolite compositions.
Boninite–harzburgite reaction experiments
We performed two anhydrous reaction experiments at
1·5 GPa and five experiments at 2 GPa where the
resulting run product was a melt in equilibrium with a
refractory harzburgite residue (Tables 2 and 3). Four of
these nominally ‘anhydrous’ experiments were analysed
by FTIR spectroscopy for H2O contents, which vary
from 0·2 to 0·7 wt % (Table 3). Probably some H2O
remained in the starting composition BON as a result of
incomplete devolatilization during firing at 1000°C (see
experimental techniques section).
In Fig. 4a and b we compare the results of our reaction
experiments with the mantle melting cotectics for the
refractory lherzolite composition TQ-40 at 1·5 and 2
GPa and for HZ at 1·5 GPa. The glass compositions in
equilibrium with Ol + Opx from our reaction experiments
are consistent with the position of the anhydrous Ol +
Opx ± Cpx cotectics determined for TQ-40 and HZ in
the following two respects:
(1) as for the anhydrous cotectics for TQ-40 at 1·5
and 2 GPa, the 1·5 and 2 GPa Ol + Opx + L cotectics
defined by the reaction experiments also show a consistent
shift in normative Ol with P in the projection from Di
(Fig. 4b);
Fig. 3. Comparison of SiO2 wt % (a), Al2O3 wt % (b) and CaO wt %
(c) against MgO wt % between experimental glass compositions from
lherzolite vs harzburgite sources. Data sources are listed in the caption
to Fig. 2 and this study (Table 3).
(2) the Ol + Opx + L cotectics defined by the reaction
experiments maintain the same orientation to those of
HZ in Fig. 4a and b but are displaced towards the Jd
+ CaTs + Lc apex of the normative tetrahedron.
As can be seen from both Figs 2b and 4b, glass
compositions from progressive melting of fertile mantle
define smooth approximately curvilinear trends in the
normative projection from Di, and glass compositions
from the melting of refractory mantle (HZ and boninite
reaction experiments) do not lie on the extrapolation of
these trends towards the Ol–Hy join. Thus the assumption
of Sobolev et al. (1993) and Sobolev & Danyushevsky
(1994) that the linear extrapolation of cotectics defined
by fertile mantle compositions could be used to estimate
the P–T conditions of boninite formation is invalid, and
results in a significant overestimation of the P of primary
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JOURNAL OF PETROLOGY
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Fig. 4. Comparison of glass compositions from ‘anhydrous’ boninite–
harzburgite reaction experiments with glass compositions from anhydrous 1·5 and 2 GPa melting experiments (this study; Falloon et al.,
1988, 1999a) on refractory lherzolite TQ-40 and refractory harzburgite
(HZ) at 1·5 GPa in the molecular normative projection from Ol (a)
onto the face [ Jd + CaTs + Lc]–Di–Qz and from Di (b) onto the base
[ Jd + CaTs + Lc]–Qz–Ol of the ‘basalt tetrahedron’ (Falloon & Green,
1988; see insert). Β, anhydrous harzburgite reaction experiments at 2
GPa; Χ, anhydrous harzburgite reaction experiments at 1·5 GPa; thick
continuous lines delineate Ol + Opx + L cotectics defined by the
harzburgite reaction experiments; cotectics for HZ as for Fig. 2.
Cotectics for TQ-40 at 1·5 and 2 GPa: continuous lines delineate Ol
+ Opx ± Cpx ± Sp + L cotectics; dot–dash line delineates an Ol + L
cotectic, with arrow in (b) pointing towards the TQ-40 bulk composition.
Other symbols as for Fig. 2.
HCB generation. Our results also demonstrate the important control of mantle bulk composition on the position of cotectics in the normative projection.
H2O-undersaturated peridotite reaction
experiments
We performed three H2O-undersaturated reaction experiments at both 1·5 and 2 GPa and four experiments
at 2·5 GPa where the resulting run product was a
melt in equilibrium with a refractory harzburgite residue
(Tables 2 and 3). Although runs T-3462 and T-3465 at
1·5 GPa and T-3461 and T-3464 at 2 GPa produced
dunite residues, the compositions of these experimental
melts are not significantly displaced from an Ol + Opx
+ L cotectic. We therefore infer that the T values of the
respective experiments are close to the T at which Opx
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FEBRUARY 2000
Fig. 5. Comparison of glass compositions from H2O-undersaturated
boninite–harzburgite reaction experiments at 1·5, 2 and 2·5 GPa with
glass compositions from anhydrous 1·5 and 2 GPa melting and reaction
experiments on refractory harzburgite (HZ) at 1·5 GPa in the molecular
normative projection from Ol (a) onto the face [ Jd + CaTs + Lc]–Di–Qz
and from Di (b) onto the base [ Jd + CaTs + Lc]–Qz–Ol of the
‘basalt tetrahedron’ (Falloon & Green, 1988; see insert). Η, hydrous
harzburgite reaction experiments at 2 GPa in equilibrium with Ol +
Opx (open diamonds with cross, in equilibrium with Ol only); Ο, hydrous
harzburgite reaction experiments at 1·5 GPa (half-filled diamond, in
equilibrium with Ol only); half-filled squares, hydrous harzburgite
reaction experiments at 2·5 GPa; thick continuous lines [labelled ‘1·5
dry’ and ‘2·0 dry’ in (b)] delineates the position of the 1·5 and 2·0 GPa
Ol + Opx + L cotectics defined by ‘anhydrous’ harzburgite reactions
(see Fig. 4b); dashed (labelled ‘1·5’), dot–dash (labelled ‘2·0’), and dotted
(labelled ‘2·5’) lines in (b) delineate the positions of Ol + Opx +
L cotectics defined by the H2O-undersaturated harzburgite reaction
experiments. Melts in equilibrium with refractory lherzolite TQ-40 are
shown in (a) for comparison. Other lines and symbols as for Figs 2
and 4.
is eliminated from the residue for the bulk composition
used. Accordingly, we use T-3462, T-3461 and T-3464
to help define the position of Ol + Opx + L cotectics
at 1·5 and 2 GPa in normative projections for H2Oundersaturated conditions (Fig. 5a and b).
In Fig. 5a and b, we compare the glass compositions
from the H2O-undersaturated reaction experiments with
the Ol + Opx ± Cpx + L cotectics defined by the
anhydrous experiments presented in Fig. 4a and b. As
can be seen from Fig. 5b, H2O has a significant effect at
both 1·5 and 2 GPa on the position of the Ol + Opx +
L cotectics for refractory mantle in the projection from
Di. The presence of H2O at a given T causes the Ol +
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FALLOON AND DANYUSHEVSKY
MELTING OF REFRACTORY MANTLE
Opx + L cotectics to move towards the Qz apex, but
maintaining their orientation relative to the anhydrous
cotectics. It can also be seen from Fig. 5b that the
displacement in the Ol + Opx + L cotectics towards the
Qz apex, at similar H2O contents, is greater at 1·5 GPa
than at 2 GPa.
Calculated H2O contents obtained from our mass
balance calculations (Table 3) suggest that H2O contents
vary from >2 to 3 wt % in the range 1450–1350°C at
1·5 GPa, 1500–1400°C at 2 GPa and 1500–1480°C at
2·5 GPa. As discussed in the analytical technique section,
the analysed H2O contents of hydrous experiments were
usually within 0·5 wt % of the calculated values, but the
latter are considered to be more accurate.
The shift in the Ol + Opx + L cotectics, defined by
the boninite–harzburgite reaction experiments, at 1·5
and 2 GPa is consistent with a change in the nature
of the melting reaction, resulting from an increased
proportion of Opx contributing to the melt (Green, 1976;
Gaetani & Grove, 1998). This change in the melting
reaction probably reflects the depolymerizing effect of
H2O on melt structure. As Opx is relatively rich in SiO2
and poorer in FeO compared with Ol, the increase in
Opx to the melting reaction is reflected in higher SiO2,
lower FeO and higher MgO contents compared with
anhydrous melts (compared on an anhydrous basis) at
the same T, as can be seen from Fig. 6a–c.
DISCUSSION
Olivine liquidus depression as a result of
dissolved H2O
To calculate the effect of dissolved H2O on the Ol
liquidus, we need to accurately calculate the anhydrous
Ol liquidus of water-bearing glasses. In this study, we
have checked and then employed the Ol geothermometer
of Ford et al. (1983) to calculate the anhydrous liquidus
of H2O-bearing glasses in equilibrium with Ol. The Ford
et al. (1983) geothermometer is empirically based on a
data set of 648 experiments at atmospheric P and 99
high-P experiments on dry mafic and ultramafic magmas
and related synthetic systems (Ford et al., 1983). The
empirical model of Ford et al. (1983) reproduces the Ol
liquidus T of their data set to ±1°C. Danyushevsky et
al. (1996) have demonstrated that the Ford et al. (1983)
geothermometer can also reproduce the experimental
T of atmospheric-P experiments conducted on MORB
compositions to an accuracy of ±10°C. In Fig. 7a–c,
we compare the calculated Ol liquidus T for 229 glass
compositions in equilibrium with Ol from published and
unpublished experimental studies performed at UT from
atmospheric P to 3 GPa and T between 1100 and 1650°C
(references in the caption to Fig. 7).
Fig. 6. SiO2 (a), FeO (b) and MgO (c) vs T (°C) for anhydrous and
H2O-undersaturated harzburgite reaction experiments at 1·5 and 2
GPa. The arrows show the shift in the Ol + Opx + L cotectics with
dissolved H2O in the melt. All melt compositions are recalculated to
100 wt % anhydrous. Β, anhydrous harzburgite reaction experiments
at 2 GPa; Χ, anhydrous harzburgite reaction experiments at 1·5 GPa;
Η, hydrous harzburgite reaction experiments at 2 GPa in equilibrium
with Ol + Opx; Ο, hydrous harzburgite reaction experiments at 1·5
GPa.
The atmospheric-P experiments using the Pt thermocouple (Fig. 7a) closely conform to a 1:1 line with an
accuracy of ±10°C. High-P experiments using the W/
Re thermocouple (Fig. 4b) are also normally distributed
around the 1:1 line, but display a slightly larger scatter
than the atmospheric-P experiments. This scatter can be
explained by a number of factors, including analytical
errors, quench modification of experimental glasses, presence of small amounts of H2O or other volatiles, and
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JOURNAL OF PETROLOGY
VOLUME 41
Fig. 7. Calculated Ol liquidus T using the Ol geothermometer of Ford
et al. (1983) vs run T for 229 nominally anhydrous glasses, with
<>3 wt % total alkalis, in equilibrium with Ol from published and
unpublished studies performed at Hobart (see text for discussion). Φ
(a), experiments using the Pt thermocouple at atmospheric P; Η (b),
high-P experiments using the W/Re thermocouple; Ο (b), melting
experiments on TQ-40 and HZ presented in this study; Β (c), high-P
experiments using the Pt thermocouple; continuous line is the 1:1 line
and dotted lines represent ±15°C from the 1:1 line. Data sources:
Green et al. (1979), n = 14; Duncan & Green (1987), n = 22; Eggins
(1992), n = 13; Kuehner (1992), n = 24; Falloon et al. (1997b), n =
2; Falloon et al. (1988), n = 51; Falloon et al. (1999a), n = 43; Falloon
et al. (1999b), n = 22; T. J. Falloon, unpublished data,1997, n = 38.
non-equilibrium. High-P experiments using the Pt thermocouple are shown in Fig. 7c. The generally higher
calculated Ol liquidus T values of these experiments are
consistent with the well-known drift of the Pt thermocouple in the relatively reducing environment of the
piston-cylinder apparatus (Holloway & Wood, 1988).
This shift was not avoided even by using short run
times (average >2·5 h for experiments using the Pt
NUMBER 2
FEBRUARY 2000
thermocouple, average >41 h for experiments using the
W/Re thermocouple). An alternative explanation, which
we consider less likely, is that the majority of high-P Pt
experiments were not entirely anhydrous and had small
amounts of dissolved H2O (0·2 wt % on average).
We are confident that the atmospheric-P experiments
and the high-P W/Re experiments are nominally anhydrous and these data clearly demonstrate that the Ford
et al. (1983) geothermometer can be used with confidence
to calculate the anhydrous Ol liquidus T for basaltic
magma compositions. Our analysis of the Ford et al.
(1983) geothermometer demonstrates that positive and
progressively larger errors start to occur in the calculated
liquidus T of basaltic liquids using the Ford et al. (1983)
geothermometer if the alkalis are >3 wt %. Our data set
has an average total alkalis of 1·8 ± 1 wt % (range
0·37–5·88).
Using the Ford et al. (1983) geothermometer, we have
calculated the Ol liquidus depression caused by dissolved
H2O for experimental glasses from the literature (and
this study; references in the caption to Fig. 8) in equilibrium with Ol where the H2O content has been analysed
or can be reasonably estimated, and these data are
presented in Fig. 8 (n = 106, H2O contents 0·2–21
wt %, alkalis <>3 wt %). In addition to the factors
contributing to errors in calculating the dry liquidus T
(see above), additional errors in determining the exact
H2O contents of the experimental glasses are also a cause
of scatter in Fig. 8.
Our approach is empirical and ignores the saturation
state of the system (under- or oversaturated in H2O).
Although there are a number of factors that control the
activity of H2O in silicate melts and hence the Ol liquidus
depression, our empirical approach demonstrates that
the most important factor in controlling the Ol liquidus
depression is the concentration of dissolved H2O in the
melt. A full thermodynamic discussion and analysis of
the data set is beyond the scope of this paper.
The empirical relationship between the calculated Ol
liquidus depression and dissolved H2O contents defined
by the data set presented in Fig. 8 (r = 0·927) is given
by the equation
olivine liquidus depression (°C) =
74·403 × (H2O wt %)0·352.
(3)
The equation that describes this empirical relationship
is non-linear with an error >9%. It can be seen from
Fig. 8 and equation (3) that the effect of H2O on the Ol
liquidus depression is significant even for small amounts
of H2O (e.g. >74°C at 1 wt % H2O; >59°C at 0·5 wt
%), thus even for nominally ‘anhydrous’ magmatic suites
with low H2O contents (e.g. MORB, back-arc basin
tholeiites, island arc tholeiites) the effects of small amounts
of water will have petrogenetic significance. The role of
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FALLOON AND DANYUSHEVSKY
MELTING OF REFRACTORY MANTLE
Fig. 8. The Ol liquidus depression caused by increasing amounts of
dissolved H2O in basaltic magmas with <>3 wt % total alkalis. The
data are from those studies where the glass composition in equilibrium
with Ol is given and the H2O content is also given or can be reasonably
estimated (n = 106; H2O contents vary from 0·2 to 21 wt %). Data:
this study; Jenner (1982); Ulmer (1989); Umino & Kushiro (1989);
Kushiro (1990); Foden & Green (1992; one outlier excluded); Sisson
& Grove (1993); Baker et al. (1994); Gaetani et al. (1994); Hirose &
Kawamoto (1995); Panjasawatwong et al. (1995, one outlier excluded);
Kawamoto (1996); Hirose (1997); Gaetani & Grove (1998, one outlier
excluded); T. J. Falloon, unpublished data. Studies that are inconsistent
with the data in Fig. 8, and are therefore not plotted, are Luhr (1990)
[Tanhydrous < Thydrous; rhyolitic compositions used by Luhr (1990) are
beyond the compositional range of Ford et al. (1983)] and Baker &
Eggler (1987) (mixed volatiles and significant Fe loss). The best fit line
through the data [equation (3), see text] is shown (R = 0·927).
H2O on the petrogenesis of MORB has been more fully
discussed by Danyushevsky et al. (2000).
As an independent test of the validity of our approach
we have used equation (3) to calculate the Ol liquidus T
of four hydrous Ol-bearing experiments from the recent
study of Moore & Carmichael (1998). As can be seen
from Fig. 9, our empirical equation satisfactorily predicts
the actual experimental T.
The non-linear effect of H2O in depressing the Ol
liquidus T is consistent with the speciation model for
water in silicate melts proposed by Stolper (1982). In
this model, water is dissolved in silicate melts as both
molecular water and as hydroxyl groups, and the proportions of species are controlled by the equilibrium
reaction
H2Omolecular(melt) + O°(melt) = 2OH(melt)
(4)
where O° refers to a bridging oxygen and OH represents
an OH group attached to a silicate polymer. Stolper
(1982) demonstrated that at low H2O contents, water is
dissolved predominantly as hydroxyl groups [see Figs 1
and 2 of Stolper (1982)] and dominates over dissolved
molecular H2O. Molecular water contents increase with
increasing water contents until at high water contents
molecular H2O is the dominant H-bearing species (Stolper, 1982). As Kushiro (1975) demonstrated that the
Fig. 9. Comparison of calculated Ol liquidus T using equation (3) (see
text) and the Ol geothermometer of Ford et al. (1983) vs experimental
T for Ol-saturated liquids from the study of Moore & Carmichael
(1998).
addition of components that lead to melt depolymerization (i.e. OH−, Mysen, 1977) enlarges the
primary liquidus volumes of less polymerized phases
relative to more polymerized phases (e.g. Ol vs Opx),
Stolper (1982) predicted that the most dramatic changes
as a result of dissolved H2O contents are effected by the
first several percent of dissolved H2O. This is precisely
the effect we observe in the decrease of Ol liquidus T
with increasing amounts of dissolved H2O in the melt.
The non-linearity of the effect of H2O on the liquidus
of silicate phases was also suggested by Hamilton et al.
(1964) and Kushiro & Yoder (1967).
Petrogenesis of HCB from Northern Tonga
and the Troodos ophiolite
Northern Tonga HCB
The experimental results from our harzburgite reaction
experiments enable us to establish a petrogenetic grid
for H2O-undersaturated melting of depleted mantle appropriate for HCB petrogenesis at 1·5, 2 and 2·5 GPa.
We are now in a position to more accurately determine
the P–T conditions of primary Tongan HCB formation.
As can be seen from Fig. 10a, the two primary Tonga
HCB compositions (Table 4) determined by the melt
inclusion study of Sobolev & Danyushevsky (1994) overlie
an >1·5 GPa Ol + Opx + L cotectic with >2–3 wt %
dissolved H2O in the melt for melting of refractory mantle
affected by an appropriate SC. Also, there is a close match
in composition between the 1·5 GPa H2O-undersaturated
melts (2–3 wt % H2O) and the HCB primary magmas
(1–2 wt % H2O) (Table 4). This result strongly suggests
that the Tongan HCB primary magma compositions
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JOURNAL OF PETROLOGY
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Fig. 10. Molecular normative projection from Di onto the base [ Jd
+ CaTs + Lc]–Qz–Ol of the ‘basalt tetrahedron’ (see insert) comparing
the primary or parental boninite compositions from Tonga (a) and
Troodos (b) with the experimentally determined mantle melting cotectics (this study) for refractory peridotite. T1 and T2 are the Tongan
primary boninite compositions (Table 4), and DG, GpI and GpIII are
the Troodos boninite compositions (Table 4). Thin dashed line in (b)
is an Ol + Opx + L cotectic at 0·8 GPa defined by the experimental
study of Duncan & Green (1987). Half-filled square is the recalculated
GpIII Upper Pillow Lava composition of Sobolev et al. (1993) (Table
4 and see text for discussion). Thick continuous and dotted lines are Ol
+ Opx + L cotectics for refractory mantle under H2O-undersaturated
conditions at 1·5 and 2 GPa, respectively, determined by boninite–
harzburgite reaction experiments.
determined by the melt inclusion technique of Sobolev
& Danyushevsky (1994) are indeed in equilibrium with
the mantle.
In Table 5 we present a comparison of the estimated T
of HCB melt generation determined by two independent
techniques, i.e. the melt inclusion technique and the
peridotite reaction experiments. Table 5 compares the
calculated Ol liquidus T at 1·5 GPa determined for
Tongan HCB compositions by Sobolev & Danyushevsky
(1994) based on the results of homogenization experiments of Ol-hosted melt inclusions at atmospheric P
with T determined by the results of this study at 1·5
GPa. MgO (anhydrous) and H2O contents for the Tongan
compositions listed in Table 5 are from Table 4. The
liquidus T values at atmospheric P are calculated using
the Ol geothermometer of Ford et al. (1983) and an
Ol liquidus depression of 95°C for 2 wt % H2O, as
experimentally determined by Sobolev & Danyushevsky
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FEBRUARY 2000
(1994) for the Tonga 1 composition and using equation
(3) for the Tonga 2 composition. The liquidus temperatures at 1·5 GPa are based on the T dependence of
the Ol liquidus slope with P based on the Ol geothermometer of Ford et al. (1983). The experimental
harzburgite equilibration temperatures are based on the
1·5 GPa ‘wet’ experiments (Tables 2 and 3) assuming
that the linear relationship between MgO and T can be
extrapolated to MgO contents appropriate for the Tonga
1 composition. Both the melt inclusion technique of
Sobolev & Danyushevsky (1994) and the peridotite reaction experiments give consistent results, and this
strongly suggests that HCB petrogenesis requires T between >1406 and 1512°C at depths of >45 km in the
mantle wedge (Table 5).
Although both techniques produced similar results
(within ±35°C), we consider that the T of HCB petrogenesis is best determined by using the Ford et al.
(1983) Ol geothermometer in combination with equation
(3). This suggests that the Tonga 1 primary melt composition is a primary magma at T >1480°C at 1·5 GPa.
However, many other workers (e.g. Fisk, 1986; Kelemen, 1995) have proposed that boninitic magmas in
general are the result of wall-rock reaction processes
between normal tholeiitic magmas and refractory harzburgite at relatively shallow levels (<1 GPa). The data
presented in this study and the melt inclusion study of
Sobolev & Danyushvesky (1994) constrain such reaction
and re-equilibration to occur at >1·5 GPa and >1480°C
in the case of HCB, if the postulated residue or reacting
wall-rock is harzburgite and the reacting melt reaches
chemical equilibrium with the harzburgite wall rock.
However, if it is postulated that reaction with harzburgite
at lower P has eliminated Opx from the wall-rock so that
the Tonga boninite composition attains its observed
composition within a dunite channel through harzburgite,
then the precursor (pre-reaction) magma for the Tongan
boninite should lie on a vector from the Qz apex through
the Tongan boninite as shown in Fig. 10a. This is because
reaction between a higher-P magma and harzburgite
at lower P will cause the reactant magma to change
composition towards the Qz apex, as a result of the
following generalized reaction:
Meltprecursor reactant + Opxwallrock → Meltnew + Olivwallrock/melt
Meltnew = Meltprecursor reactant + SiO2.
(5)
In the projection from Di (Fig. 10a), such precursor
melts could lie on higher-P Ol + Opx + L residue trends
and would be less Hy rich than the Tongan boninite
composition, but would lie at significantly higher T
(>1480°C). Therefore in terms of wall-rock reaction
models of magma genesis, the composition of the Tongan
boninite represents the minimum P (>1·5 GPa) and T
(>1480°C) of formation. However, we believe that higher
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FALLOON AND DANYUSHEVSKY
MELTING OF REFRACTORY MANTLE
Table 4: High-Ca boninite compositions from North Tonga and Troodos
No.:
1
2
3
4
5
6
7
8
Tonga
Tonga
T-3485
T-3486
Troodos
Troodos
Troodos
Troodos
1
2
1·5 GPa
1·5 GPa
DG
Gp I
Gp III
Gp III
SiO2
51·92
52·68
53·10
52·50
52·40
51·51
50·63
TiO2
0·10
0·40
0·74
0·95
0·30
0·33
0·20
0·23
Al2O3
7·07
9·52
8·30
11·00
11·70
11·36
10·27
11·81
FeO∗
9·29
8·63
7·60
8·10
8·40
8·30
8·84
9·14
MnO
0·20
0·20
0·14
0·14
MgO
24·24
19·05
21·20
16·10
15·80
18·35
20·94
16·33
CaO
6·46
7·64
6·60
8·60
10·70
8·61
8·54
9·82
Na2O
0·61
1·29
1·50
1·80
0·70
1·18
0·51
0·59
K 2O
0·10
0·60
0·32
0·46
0·10
0·09
0·07
0·08
H 2O
mg-no.
2·0
1·1
1·7
1·5
82·3
79·7
83·3
78·0
52·00
0·16
—
77·0
2
2
2
79·7
80·8
76·1
Temperature (°C) (1·5 GPa)
Dry
1572
1510
1545
1465
1432
1487
1521
1441
Wet
1477
1433
—
—
—
1392
1426
1346
Exp.
—
—
1450
1350
—
—
—
—
Tonga 1 and Tonga 2 are the primary western and eastern group boninite melts respectively from Sobolev & Danyushevsky
(1994). Experimental glass compositions (nos 3 and 4) are from this study (Tables 2 and 3). Troodos DG is the calculated
parental composition to the Upper Pillow Lavas of the Troodos ophiolite of Duncan & Green (1980); Troodos Gp I is the
most magnesian homogenized olivine (Fo93·2) hosted melt inclusion from the Gp I Upper Pillow Lavas (Portnyagin, 1997)
and Gp III (no. 7) is the primary boninite magma composition for the Gp III Upper Pillow Lavas from Sobolev et al. (1993).
Troodos Gp III (no. 8) is the recalculated Sobolev et al. (1993) composition (see text for discussion). All compositions have
been renormalized to 100 wt % anhydrous. FeO∗refers to total Fe as FeO. ‘Dry’ refers to calculated anhydrous olivine liquidus
temperatures at 1·5 GPa using the olivine geothermometer of Ford et al. (1983); ‘Wet’ refers to calculated hydrous olivine
liquidus temperatures using equation (3) (see text) and the known water contents listed. ‘Exp’ refers to the run temperatures
listed in Table 2.
T values are unrealistic, and therefore reaction models
for HCB are not viable.
Troodos HCB
The Upper Pillow Lavas of the Troodos ophiolite can
be divided into three geochemical groups (Gp I, II
and III) based on Al2O3/TiO2 values and incompatible
element content (Cameron, 1985; Duncan & Green,
1987; Sobolev et al., 1993). Sobolev et al. (1993) and
Portnyagin (1997) defined the composition of the most
magnesian Ol phenocrysts for each lava group: >Fo94
in Gp I and II lavas; and >Fo92 in Gp III. As the Gp
III lavas are the most refractory in terms of Al2O3/TiO2
and REE abundances, Sobolev et al. (1993) assumed that
the composition of the most magnesian Ol for the Gp
III lavas should also be at least Fo94. In Table 4, we
list the compositions of calculated primary or parental
magmas for the end-member Gp I. As the primary or
parental magmas for Gp I and II are very similar, and
no suitable melt inclusions were found in Gp II lavas,
we assume that the established Gp I parental composition
is representative of the Gp II lavas as well. In Fig. 10b
we compare calculated Troodos primary or parental
magma compositions with the experimental melts for
refractory mantle, and in Fig. 11a–c we compare the
Troodos primary or parental magma compositions with
our experimental data on oxide vs MgO plots.
The Gp I primary composition (Table 4, Figs 10b and
11) is based on homogenized melt inclusions in magnesian
Ol (Fo93·2, Portnyagin, 1997). The Gp I primary composition, although having lower SiO2, and slightly higher
Al2O3 and CaO contents than the experimental data
(Fig. 11), is consistent with melting of refractory mantle
at >1·5 GPa under H2O-undersaturated conditions (Fig.
10b). The compositional differences between Tongan
and Troodos Gp I HCB primary compositions probably
reflect differences in the degree of depletion of the mantle
source (Tonga more refractory).
As can be seen in Fig. 10b, the plotted position of the
Gp III composition appears to be consistent with melting
of refractory mantle at >1·5 GPa. However, as can be
seen from Fig. 11a–c there are significant differences
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Table 5: Temperatures of
primary HCB melt
generation
Tonga 1
Tonga 2
MgO wt %
24·24
19·05
H2O wt %
2·0
1·1
Melt inclusion approach
Liquidus temperatures (°C)
at 0·1 GPa (dry)
1497
1436
1402
1359
1477
1433
1512
1406
Liquidus temperatures (°C)
at 0·1 GPa (wet)
Liquidus temperatures (°C)
at 1·5 GPa (wet)
Experimental approach
Harzburgite equilibration
temperatures (°C) at 1·5 GPa
(H2O 2–3 wt %)
Table 5 compares olivine liquidus temperatures at 1·5 GPa
experimentally determined for Tongan high-Ca boninite compositions from homogenization experiments of olivine-hosted melt inclusions by Sobolev & Danyushevsky (1994) with
temperatures determined by the results of this study at 1·5
GPa (see text for discussion).
between the Gp III primary composition of Sobolev et
al. (1993) and the experimentally determined, H2Oundersaturated melts at 1·5 GPa for refractory mantle.
The Gp III boninite primary composition of Sobolev et
al. (1993) has significantly higher Al2O3 and CaO, and
lower SiO2, for it to be in equilibrium with a refractory
harzburgite mantle at 1·5 GPa with >2 wt % H2O in
the melt. These compositional characteristics are in the
opposite direction to those expected for a source more
refractory than the Gp I and II primary or parental
magmas as inferred from the geochemistry of Gp III
(Sobolev et al., 1993). Possibly the assumption made by
Sobolev et al. (1993) that Fo94 was the most magnesian
Ol for the Gp III lavas was in error. Assuming the Gp
III composition is in equilibrium with the most magnesian
Ol found for this suite (Fo92), we have listed this composition in Table 4 and it is plotted in Figs 10b and 11.
The recalculated Gp III composition is only slightly more
magnesian (16·33 vs 15·80 wt %) than the parental picrite
composition calculated by Duncan & Green (1980) (Table
4). The parental composition of Duncan & Green (1980)
was based on the major element variations of the entire
Upper Pillow Lava suite (using mainly Gp III samples)
and the most magnesian Ol known at that time (Fo91·7).
Duncan & Green (1987) determined that the calculated
parental composition of Duncan & Green (1980) is in
Fig. 11. SiO2 (a), Al2O3 (b) and CaO (c) vs MgO wt % for the Troodos
parental or primary boninite compositions (Table 4) compared with
the anhydrous mantle melting cotectics (Ol ± Opx ± Cpx ± Sp + L)
for fertile mantle (MM-3) at 1·0 (dash–dot line) and at 1·5 GPa
(continuous line) and for refractory mantle (HZ, dashed line) at 1·5
GPa. Dashed line labelled ‘0·8’ in (a) is the Ol + Opx + L cotectic
defined by the experiments of Duncan & Green (1987) at 0·8 GPa.
Also plotted are the compositions of glasses (labelled fields) from H2Oundersaturated boninite–harzburgite reaction experiments at 1·5 and
2·0 GPa. The field labelled ‘1·5 ‘Boninite’’ also includes the Tongan
primary or parental composition T2 (Table 4). Filled arrow is an Ol
control.
equilibrium with a mantle harzburgite residue at 0·8
GPa at 1360°C under anhydrous conditions. In Fig. 10b,
the Duncan & Green (1980) parental composition is used
to establish the position of a dry 0·8 GPa mantle melting
Ol + Opx + L cotectic. By comparing compositional
differences of coexisting Ol, Opx and Cpx produced experimentally with natural occurrences, Duncan & Green
(1987) concluded that their calculated parental composition contained >1 wt % H2O and was in equilibrium
with a harzburgite wall-rock assemblage at P only slightly
278
FALLOON AND DANYUSHEVSKY
MELTING OF REFRACTORY MANTLE
above 0·8 GPa. The new data available on the Gp III
lavas (Sobolev et al., 1993; Portnyagin, 1997) and the
results of this study fully support the results of Duncan
& Green (1987). The recalculated Gp III primary magma
of Sobolev et al. (1993) (Table 4) contains >2 wt %
H2O and is in equilibrium with a harzburgite wall-rock
assemblage at >1–1·2 GPa.
Previous experimental work on boninite
petrogenesis
The results of this study suggest primary HCB magma
genesis at >1·5 GPa, with >1–2 wt % H2O in primary
magmas at >1430–1480°C, in substantial disagreement
with previous experimental studies suggesting T between
1130 and 1260°C, P ranging from 0·3 to 1·8 GPa, and
water contents ranging from 1 to 20 wt % (Umino &
Kushiro, 1989; Van der Laan et al., 1989). Our results
differ from these previous studies for the following
reasons: (1) the experimental HCB compositions were
chosen by Umino & Kushiro (1989) and Van der Laan
et al. (1989) on the basis that whole-rock compositions
had mg-numbers appropriate to be in equilibrium with
a mantle Ol of >Fo88, and not by Ol compositions
actually occurring in the studied suites; (2) the experiments
were performed under a wide range of H2O-bearing
conditions, and, in some studies (Umino & Kushiro,
1989), no attempt was made to constrain actual H2O
contents in the natural magmas. These previous experimental studies, although useful in defining crystallization conditions and potential liquid lines of descent,
are not appropriate for defining conditions of primary
boninite magma genesis.
The results of this study are also in disagreement with
previous suggestions that melting of fertile to depleted
mantle at low P can produce boninite compositions under
anhydrous conditions (e.g. Takahashi & Kushiro, 1983;
Takahashi et al., 1993; Klingenberg & Kushiro, 1996).
At low P and under anhydrous conditions, Opx melts
incongruently and consequently mantle melts have higher
SiO2, and relatively lower FeO at a given MgO than
higher-P melts. Several workers have noted this effect,
and have suggested that these low-P melts are in fact
‘boninites’. The error is that only two characteristics of
boninites (relatively high MgO at intermediate SiO2) have
been used without consideration for the compositional
characteristics of boninite suites themselves. Anhydrous
experiments are not relevant to H2O-rich boninite suites.
The effect of subduction components on
mantle melting
In Fig. 12a to f, the SiO2, Al2O3, FeO, MgO, CaO and
Na2O contents of glass compositions from our 1·5 GPa
boninite reaction and HZ melting experiments are plotted
against their calculated anhydrous Ol liquidus T using
the Ford et al. (1983) geothermometer. In Table 6 the
compositions of runs T-3472 (‘anhydrous’ reaction experiment, representing a mantle melt produced in the
presence of a dry SC), T-3485 (H2O-undersaturated
reaction experiment, representing a mantle melt produced in the presence of a wet SC) and T-4349 (anhydrous HZ reversal experiment, representing a mantle
melt produced in the absence of an SC) are compared
on an anhydrous basis. The glass compositions from
these three experiments have almost identical calculated
anhydrous Ol liquidus T using the Ford et al. (1983)
geothermometer (Table 6).
Our discussion on the effect of the SC on mantle
melting at 1·5 GPa is based on the following two important assumptions: (1) the added boninite component
in our reaction experiments is representative of a primary
melt produced by the melting of a refractory harzburgite
in the presence of an SC; (2) the HZ composition is
representative of the refractory harzburgite composition
involved in HCB boninite petrogenesis. As neither of
these assumptions can be proven the results of our
comparison are only of a preliminary nature.
As can be seen from Fig. 12 and Table 6, the effect
of the ‘dry’ SC is to produce mantle melts with relatively
lower SiO2 and MgO, and relatively higher FeO, Al2O3
and Na2O contents. The effect of H2O in the ‘wet’ SC
is to produce mantle melts with higher SiO2 and lower
FeO contents relative to mantle melts produced in the
presence of the ‘dry’ SC (see Fig. 6). This results in
mantle melts with similar SiO2 and FeO contents to
mantle melts produced without an added SC. The effect
of the SC on mantle melting cotectics is dramatic, as
illustrated in Fig. 5b. The mantle melting cotectic at 1·5
GPa for the refractory HZ composition moves towards
the Ol apex of the tetrahedron (towards relatively higher
P) with the addition of the SC component. A very
significant result from our experimental study is that the
effect of the SC component on the movement (towards
Ol) of mantle melting cotectics as seen in the basalt
tetrahedron is in the opposite direction to the effect of
H2O as a pure component (towards Qz). Therefore the
results of experimental studies that seek to model mantle
melting in the subduction zone environment based on
the simple addition of H2O as a proxy for the SC should
be interpreted with caution. Our experimental results
suggest that it is important to constrain the nature of
components in addition to H2O so as to use experimental
mantle melting data to constrain the P and T of subduction zone primary magmas.
In the case of HCB petrogenesis, the addition of the
SC has resulted in significantly lower CaO/Al2O3 and
CaO/Na2O values compared with melting in the absence
of a subduction component (Table 6). This suggests that
279
JOURNAL OF PETROLOGY
VOLUME 41
NUMBER 2
FEBRUARY 2000
Fig. 12. SiO2 (a), Al2O3 (b), FeO (c), MgO (d), CaO (e) and Na2O (f ) vs calculated anhydrous Ol liquidus T using the Ol geothermometer of
Ford et al. (1983) for experimental glasses at 1·5 GPa. Η, direct melting and reversal experiments on HZ; Χ, anhydrous boninite–harzburgite
reaction experiments; Ο, H2O-undersaturated boninite–harzburgite reaction experiments.
the SC has high Al2O3 and Na2O in addition to H2O.
This agrees with the work of Pearce et al. (1992), who
suggested that the SC in the case of the Bonin–Mariana
Forearc boninites is a hydrous melt of subducted amphibolitized ocean crust that is high in Al2O3 and Na2O.
Implications for melt generation in
subduction zones
The results of this study confirm and strengthen the
conclusions of Sobolev & Danyushevsky (1994) that HCB
genesis involves significantly high T (>1480°C) at relatively shallow depths in the mantle wedge above a
subducting slab. Any geodynamic model of boninite
formation must take into account this well-constrained
petrogenetic information. As discussed by Danyushevsky
et al. (1995), in the case of the northern Tongan HCB,
models involving melting of cold metasomatized mantle
wedge by contact melting from normal tholeiitic magmas
[with mantle potential T (Tp) >1280°C, McKenzie &
Bickle, 1988] can be excluded. Instead, the Tongan
HCB can be simply explained by H2O-fluxed melting of
280
FALLOON AND DANYUSHEVSKY
MELTING OF REFRACTORY MANTLE
general cannot involve processes involving the interaction
of the mantle wedge with shallow upwelling asthenosphere (Tp >1280°C).
Table 6: Glass compositions from runs
T-3472
T-3485
T-4349
SC ‘dry’
SC ‘wet’
no SC
SiO2
51·90
53·10
53·40
TiO2
0·71
0·74
1·27
Al2O3
8·27
8·30
6·20
FeO
8·4
7·60
7·50
MnO
0·12
0·14
0·30
21·20
22·40
MgO
21·7
CaO
6·74
6·60
7·20
Na2O
1·5
1·50
0·50
K2 O
0·31
0·32
0·33
Cr2O3
0·18
0·38
0·95
H2O
0·5
2·1
CaO/Al2O3
0·81
0·79
CaO/Na2O
4·5
4·4
14·4
82·15
83·25
84·18
mg-no.
ACKNOWLEDGEMENTS
We acknowledge the technical assistance of David Steele,
Nick Ware, Wieslav Jablonski and Graeme Rowbottom.
We also thank Professor David Green for helpful discussions on mantle melting and comments on this manuscript. T.J.F. and L.V.D. acknowledge research support
from the Australian Research Council (ARC). T.J.F.
acknowledges support from a Royal Society–ARC Endeavour Research Fellowship, and L.V.D. acknowledges
support from ARC Postdoctoral and Queen Elizabeth II
Research Fellowships. We acknowledge support of the
Museum of Natural History, Washington, DC, which
provided electron microprobe standards. This manuscript
was improved by the constructive reviews of Simon
Turner, Richard Arculus and an anonymous referee.
—
1·16
T°Crun
1500
1450
1550
T°CFord
1552
1545
1540
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