Download Garnet in the Earth`s Mantle

Garnet in the Earth’s Mantle
Bernard J. Wood*, Ekaterina S. Kiseeva* and Andrew K. Matzen*
1811-5209/13/0009-421$2.50
DOI: 10.2113/gselements.9.6.421
luminous garnet, (Mg,Fe2+,Ca)3 (Al,Cr) 2Si3O12 , is an important constituent of mantle peridotite (~10%) and of the other abundant upper
mantle rock, eclogite (~50%). Its unusual crystal chemistry means that
it strongly prefers some trace elements and confers a “garnet signature” on
mantle melts. As depth increases from 250 to 600 km, garnet increases in
abundance in mantle rocks, dissolving large fractions of the other silicates
and becoming Si rich (majoritic). These compositional changes are observed
in some garnets found as inclusions in diamond. Garnet disappears from
mantle assemblages at about 700 km depth, where it is replaced by an even
denser silicate, perovskite.
A
roxene in addition to olivine and
orthopyroxene. Harzburgites and
lherzolites are generally thought
to be related to one another,
with harzburgite being the solid
residuum after partial melting of
more fertile “primitive” lherzolite.
This relationship led Ringwood
(1962) to propose that the primitive, or unmelted, mantle has a
composition that is a mixture
of residual peridotite (75%) and
basalt (25%). In this mantle
composition, which Ringwood
KEYWORDS : phase transformations, garnet, mantle peridotite, eclogite, diamond
named “pyrolite,” garnet becomes
inclusions, majorite
stable at the expense of spinel at
pressures above 2.8 GPa (>85 km
INTRODUCTION
depth; FIG. 1) at the solidus. The reaction relationship
Garnet is a common mineral in crystalline rocks of the
between the two aluminous minerals (spinel and garnet)
Earth’s upper mantle and lower crust and occasionally may be represented approximately as:
occurs in volcanic liquids. Garnet is unlike most other
crustal minerals in that Mg and Fe2+ are 8-coordinated by
0.4Ca(Mg,Fe2+)Si2O6 + 3.2(Mg,Fe2+)SiO3 + (Mg,Fe2+)Al2O4 =
oxygen while Al occupies a site with 6-fold coordination
clinopyroxene
orthopyroxene
spinel
(1)
(Mg and Fe2+ are usually 6-fold and Al 4-fold coordinated
(Ca,Mg,Fe2+) 3Al2 Si3O12 + (Mg,Fe2+) 2 SiO4
in other minerals) – garnet crystal chemistry is discussed
garnet
olivine
further in Geiger (2013 this issue). As a result of their
high cation coordination numbers, garnets have relatively
o
high densities, low compressibilites and increasing stability At the solidus temperature of 1460 C, the transition from
spinel lherzolite to garnet lherzolite takes place over a depth
with increasing pressure. In consequence, garnet-bearing
rocks have higher seismic velocities and densities than interval of about 5 km (Robinson and Wood 1998). The
low-pressure, garnet-free rocks of similar composition, a transformation has a positive pressure–temperature slope of
characteristic that has been used in attempts to constrain ~40 bar/degree but is strongly curved because, as temperature decreases, both the solubility of alumina in pyroxene
the compositions of the lower crust, upper mantle and
transition zone. The high coordination numbers also mean and the disorder of Mg and Al in the spinel decrease. These
changes have the effect of flattening the slope of the transithat garnet residual to, or fractionating from, silicate liquids
imparts a characteristic geochemical signature on the tion in pressure–temperature space (FIG. 1). The overall
result is that garnet peridotite is stable only in relatively
trace element pattern of the product igneous rock. These
cool and deep parts of the mantle (conditions that are often
properties make garnet a petrologically and geochemically
important constituent of the upper mantle and transition observed beneath thick cratonic lithosphere); elsewhere,
spinel peridotite is the stable form.
zone. In this article, we explore garnet’s occurrence in
the mantle and its influence on the physical properties of
Compositionally, peridotitic garnets are composed of about
mantle rocks and on the geochemical properties of mantle- 75% pyrope (Mg Al Si O ), 10% grossular (Ca Al Si O )
3
2 3 12
3
2 3 12
derived silicate melts.
and 15% almandine (Fe 2+ Al Si O ). They have lower
3
GARNET IN UPPER MANTLE PERIDOTITES
Mantle rocks are observed at the surface either as tectonic
fragments (kilometre scale) or as inclusions in explosive
eruptive rocks (centimetre scale). The majority of samples
are peridotites, either harzburgites (~85% olivine, 15%
orthopyroxene) or lherzolites, which contain clinopy-
* Department of Earth Sciences, University of Oxford
South Parks Road, Oxford OX1 3AN, UK
E-mail: [email protected]
E LEMENTS , V OL . 9,
PP.
421–426
2
3
12
magnesium numbers [Mg# = molar Mg/(Mg+Fe2+)] than
coexisting olivines, and the partitioning of Mg and Fe2+
between garnet and olivine provides a useful geothermometer for peridotites (O’Neill and Wood 1979).
At pressures above the spinel stability field, garnet incorporates the chromium and Fe 3+ that are released on
spinel breakdown. These elements substitute for Al3+ in
the octahedral site. However, in contrast to the calcium
end-members, Ca3Cr2Si3O12 (uvarovite) and Ca3Fe3+2Si3O12
(andradite), which are stable at crustal pressures, the Mg–
Cr and Mg–Fe3+ end-member garnets are very unstable.
421
D ECEMBER 2013
100
2.5
Garnet
peridotite
75
2
Spinel
peridotite
1.5
1
800
Solidus
50
of the transition zone (410–660 km depth). In this depth
interval, fertile peridotitic mantle becomes a bimineralic rock consisting of ~65% wadsleyite or ringwoodite
(both high-pressure forms of olivine) and ~35% garnet
(Ringwood 1991). The increase in the modal proportion
of garnet takes place gradually over a wide depth interval
due to pyroxene breakdown and its dissolution into the
garnet structure as Mg and Si are transferred from 6- and
4-coordination in pyroxene to mixtures of 8- and 6-coordination and 6- and 4-coordination, respectively, in garnet.
The continuous reactions result in increasing densities and
seismic velocities of the mantle and may be simplified as
follows:
900 1000 1100 1200 1300 1400 1500
Temperature (oC)
The stability fields of spinel peridotite and garnet
peridotite in pressure–temperature space. The intersection with the solidus is from Robinson and Wood (1998). The
extension of the field boundary to lower temperatures is from a
number of experimental studies.
FIGURE 1
olivine
(2)
orthopyroxene
As mentioned earlier, garnets have higher coordination
numbers and higher densities than olivines or pyroxenes;
“skiagite” garnet, Fe 2+ 3Fe3+ 2 Si3O12 , is no different. This
means that, as pressure increases, the reaction will favour
the reactant, or “skiagite,” side. As a result, at a fi xed oxygen
fugacity relative to a standard buffer, such as fayalitemagnetite-quartz (FMQ), garnet will become richer in
Fe3+ with increasing pressure. Alternatively, with increasing
depth, a garnet of fi xed composition (Fe3+ content) will be
stabilised at progressively lower oxygen fugacity relative
to FMQ. Woodland and Koch (2003) measured garnets
from Kaapvaal craton peridotites and found modest
increases in Fe3+/Fe 2+ with increasing pressure, due, in
part, to transfer of Fe3+ from clinopyroxene to garnet
with increasing temperature. Calculated oxygen fugacities,
however, decline relative to the FMQ buffer in the same P–T
interval (Woodland and Koch 2003) because of the stabilisation of skiagite with increasing pressure. Some of the
interesting side effects of this change are that coexisting
C–H–O fluid phases should change from CO2 rich to CH4
rich with increasing depth and that the mantle may have a
low enough oxygen fugacity to precipitate Ni-rich metallic
alloy at depths greater than 300 km.
With increasing pressure, mantle peridotite undergoes
several important phase transitions in which garnet is a
key player (Ringwood 1991). FIGURE 2 shows the proportions
of different phases in peridotitic mantle as a function of
depth. The modal amount of garnet increases from about
10% at the spinel–garnet transition to 35% in the middle
E LEMENTS
2Ca(Mg,Fe2+)Si2O6 = Ca 2 (Mg,Fe2+)(MgSi)VISi3O12
clinopyroxene
majoritic garnet
(4)
The result of dissolution of “excess” Mg (and other divalent
cations) and Si into garnet (the “majorite” substitution)
is that its Al 2O3 content declines gradually from ~22 to
8 wt% while its CaO content increases from ~4 to 7 wt%
as pressure increases from 3 to 19 GPa and pyroxenes are
absorbed into its structure (Wood 2000).
Beyond 19 GPa (550 km depth), CaSiO3 perovskite becomes
stable in peridotite (FIG. 2) and increasing depth leads to
progressive extraction of Ca from garnet; for example,
at ~23 GPa, which corresponds to the appearance of
(Mg,Fe2+)SiO3 perovskite, the CaO contents of garnet are
lower than 4 wt% (Wood 2000). Perovskite initially forms
at the expense of ringwoodite, but as pressure increases,
the majoritic [(Mg,Fe2+) 3 (MgSi)VISi3O12 ] component of the
garnet begins to dissolve in the (Mg,Fe)SiO3 perovskite,
increasing the alumina content of the remaining garnet
(Wood 2000). Eventually the (Mg,Fe 2+)SiO3 perovskite
dissolves the remaining aluminous garnet as the components Al 2O3, Fe3+ AlO3 and (Mg,Fe 2+)(Al,Fe3+)O2.5 (Frost
and Langenhorst 2002), and garnet disappears from the
peridotite assemblage.
200
Depth
km
The 6-fold-coordinated sites in garnet can incorporate
enough Fe3+ that their contents can be easily measured
by Mössbauer spectroscopy and calibrated as an oxygen
barometer for garnet peridotites (Gudmundsson and Wood
1995; Stagno et al. 2013). There are several possible ways to
represent the equilibrium between the three iron-bearing
phases, the simplest being:
garnet (skiagite)
(3)
NaAlSi2O6 + Ca(Mg,Fe2+)Si2O6 = NaCa(Mg,Fe2+)(AlSi) (5)
VI Si O
3 12
clinopyroxene
majoritic garnet
The substitution of Cr and Fe3+ for Al is facilitated by the
addition of Ca to the garnet, an example of the “reciprocal
solution” effect (Wood and Nicholls 1978).
2Fe2+ 3Fe3+2 Si3O12 = 4Fe2+2 SiO4 + 2Fe2+ SiO3 + O2
4(Mg,Fe2+)SiO3 = (Mg,Fe2+) 3 (MgSi)VISi3O12
orthopyroxene
majoritic garnet
300
Olivine
(Mg,Fe) SiO
2
Pyroxene
4
400
13
Wadsleyite
Garnet
500
600
700
800
Ringwoodite
21
(Mg,Fe)SiO3 perovskite
with up to 5% Al O
2
CaSiO3
perovskite
3
+ (Mg,Fe)O
900
0.4
0.5
0.6
0.7
0.8
0.9
Volume fraction
FIGURE 2 Mineral proportions in fertile peridotite (pyrolite) as
a function of depth in the mantle (Ringwood 1991).
Note the increasing modal amount of garnet with depth from
200 to 450 km due to the dissolution of pyroxene into garnet.
Garnet becomes more majoritic in this depth interval (see text).
422
Pressure
GPa
3
Depth
km
GPa
3.5
D ECEMBER 2013
29
1
GARNET AND SEISMIC VELOCITIES
IN THE MANTLE
The densities and elastic properties of garnet and the
other silicates of the mantle may, in principle, be used
to investigate radial and lateral variations in mantle
bulk composition. For example, using predominantly
low-pressure–low-temperature elastic data, Anderson and
Bass (1986) suggested that the transition zone (410–660
km) is composed of a relatively garnet-rich composition,
which they called “piclogite.” Piclogite is similar to pyrolite
in that it is a mixture of depleted peridotite and basalt, but
piclogite has much more basalt. The result is that, under
upper mantle conditions, a piclogitic mantle would have
abundant pyroxene and garnet and only ~30% olivine
(compared to ~65% olivine for pyrolite). A transition zone
of piclogite would move the mantle Mg/Si ratio closer to
that of the Sun, something that is cosmochemically attractive. Piclogite’s lower olivine content generates smaller 410
and 660 km p-wave discontinuities (olivine–wadsleyite and
ringwoodite–perovskite transitions; FIG. 2) than pyrolite,
but a steeper velocity gradient between the 410 and 660 km
discontinuities because of the greater extent of transformation of pyroxene to garnet and higher modal proportion
of the latter (Anderson and Bass 1986). P-wave velocities
of piclogite are reasonably consistent with those of the
Preliminary Reference Earth Model (PREM; Dziewonski
and Anderson 1981), which has discontinuities of 2.6%
and 4.8% in p-wave velocity at 400 and 670 km, respectively. PREM is based largely on the free oscillations of the
Earth, however, and a model based purely on p- and s-wave
velocities, such as ak135 (Kennett et al. 1995), gives larger
discontinuities of 4.0% at 410 km and 5.8% at 660 km.
The difficulty of arriving at a defi nitive compositional
model based solely on wave speeds is compounded by
lateral variations in seismic velocities and uncertainties
in the elastic properties of the minerals of interest at high
pressures and temperatures. Most recent work suggests a
mantle that is broadly peridotitic (pyrolitic) but with an
increasing proportion of basaltic component in the transition zone and lower mantle (Cobden et al. 2009).
ECLOGITES AND RELATED ROCKS
IN THE MANTLE
Although most upper mantle xenoliths are peridotites
(FIG. 3B), eclogites are a second major lithology in kimberlite and dominate the xenolith lithologies in five welldocumented diamondiferous kimberlite pipes: Roberts
Victor and Bobbejaan (South Africa), Zagadochnaya
(Russia), Orapa (Botswana) and Koidu (Sierra Leone).
Eclogites are rocks composed mostly (≥75%) of subequal
amounts of garnet and omphacitic clinopyroxene (FIG. 3C).
Unlike garnets from peridotite lithologies, garnets from
A
FIGURE 3
mantle eclogite xenoliths are Cr-poor, Ca-rich pyrope–
grossular–almandine mixtures with wide variations in the
grossular content (Jacob 2004). Arguably, the most important feature of eclogites is that their bulk compositions
often closely resemble the compositions of basalts. The
amount of eclogite in the mantle, its origins, and the part
it plays in petrogenesis are subjects of considerable debate.
Noting that basalts and eclogites are compositionally
similar but that garnet-rich eclogites are the denser, highpressure form, it was suggested on a number of occasions
in the early to mid 20th century that the Moho corresponds
to the phase transformation of basalt to eclogite and that
the upper mantle is eclogitic rather than peridotitic.
This model, which would explain basalt genesis as near
total melting of the upper mantle, was discredited by the
experimental work of D. H. Green and A. E. Ringwood on
the basalt–eclogite transformation (Green and Ringwood
1967). These authors found that the conditions of the
transformation are such that eclogite is stable throughout
most of the continental crust and hence could not be
responsible for the Moho discontinuity under the continents. Furthermore, the transformation is spread out over a
considerable pressure interval, unlike the seismically sharp
Moho. With increasing pressure, the pyroxene–plagioclase
assemblage of basalt is joined by garnet at about 1 GPa,
which grows in modal amount at the expense of plagioclase and pyroxenes. Eventually plagioclase is eliminated
when the pressure is sufficiently high for clinopyroxene
to accommodate all the sodium in the rock as NaAlSi2O6
and NaFe3+ Si2O6 components (~2.5 GPa).
Although not responsible for the Moho, the basalt–eclogite
transformation must occur at the tops of subducting slabs
where mid-ocean ridge basalt (MORB) is converted to the
eclogite assemblage of garnet plus clinopyroxene. At still
higher pressures, eclogite undergoes further phase transformations analogous to those in peridotite, principally, the
conversion of pyroxene components into garnet. FIGURE 4
(Irifune et al. 1986; Ricolleau et al. 2010) illustrates the
phase relations of eclogite as a function of depth in the
Earth. The dissolution of pyroxene components into garnet
leads to a rock that is ~95% garnet at depths of about
500 km, the remainder being small amounts of excess
SiO2 which crystallises as stishovite at any pressure above
9 GPa. As pressure increases, as for peridotitic compositions, CaSiO3 perovskite becomes stable at about 18 GPa.
As the modal proportion of CaSiO3 perovskite increases,
the CaO content and modal proportion of garnet gradually decline until (Mg,Fe)SiO3 perovskite becomes stable
at about 23 GPa (~660 km). Thereafter, the fraction of
garnet rapidly declines such that beyond 750 km depth,
(Mg,Fe)SiO3 perovskite, CaSiO3 perovskite, stishovite and a
new aluminous (NAL) and/or a calcium ferrite phase make
B
C
(A) Garnet inclusion in diamond. (B) Garnet peridotite xenolith. (C) Eclogite xenolith.
IMAGES COURTESY OF Z. SPETSIUS (A) AND K. MACDONALD (B)
E LEMENTS
423
D ECEMBER 2013
of other high-pressure minerals trapped as inclusions in
diamonds (Harte 2010) reinforce the hypothesis that some
of the diamonds have originated well below the cratons.
300
Pressure
GPa
Depth
km
200
Pyroxene
400
GARNET AND GEOCHEMISTRY
OF VOLCANIC ROCKS
13
Garnet
500
600
Considerable effort in both experimental and observational
geochemistry has been expended in trying to understand
the controls on the major and trace element compositions of melts that are erupted at the surface of the Earth.
Garnet does not crystallise at low pressures in most igneous
systems but, in a fertile peridotite, it is stable below ~85 km
depth (2.8 GPa) at the mantle solidus (see FIG. 1). Thus, if
we can identify a “garnet signature” in eruptive products
we have good reason to believe that there was, at least,
some melting taking place at depths greater than 85 km.
Seismological evidence of melt at the lithosphere–asthenosphere boundary to at least 75 km depth (Qi et al. 1997)
suggests that garnet is likely to be present in the source
region of at least some basalts. In this section we review
how the presence, or absence, of garnet in the source region
affects the trace element geochemistry of melts that we
observe at the surface.
21
700
(Mg,Fe)SiO3
perovskite
800
900
0
0.2
SiO2
NAL
stish
phase
0.4
0.6
Volume fraction
CaSiO3
perovskite
0.8
29
1.0
Composite diagram (Irifune et al. 1986; Ricolleau et al.
2010) showing phase transformations and proportions
for an eclogite composition as a function of depth in the Earth.
Note that eclogite becomes essentially a garnetite in the deeper
parts of the transition zone (410–660 km). Stish = stishovite.
NAL = new aluminous phase, Na-rich
FIGURE 4
up the modal mineralogy. Ringwood (1991) argued that
subducted oceanic crust would be dominantly composed
of garnet at 650 km depth and that this “garnetite” should
be less dense than the surrounding peridotite and have the
potential to form a perched basaltic layer in the mantle.
More recent density data indicate, however, that eclogite is
denser than ambient mantle throughout the lower mantle
(Ricolleau et al. 2010) and that recycled basalt can subduct
to the core–mantle boundary.
INCLUSIONS IN DIAMONDS
Although experiments provide constraints on the majorite
substitution mechanisms (reactions 3–5) in mantle garnet,
natural occurrences of majoritic garnet are rare because
most xenoliths have apparent equilibration pressures of
<7 GPa (<220 km). Under these conditions, as can be seen
from FIGURE 2, the dissolution of pyroxene into garnet has
hardly begun. Therefore, in order to fi nd natural examples
we need to turn to mineral inclusions in diamond (FIG. 3A).
As the most common silicate inclusion in diamonds, garnet
has the potential to provide unique information about old
continental cratons, whose keels extend to depths of more
than 200 km.
FIGURE 6 shows the chondrite-normalised rare earth element
(REE) contents of clinopyroxene and garnet from peridotite
xenoliths found in the Vitim volcanic field (Ionov et al.
1993). Although the xenoliths all contain olivine, orthopyroxene, clinopyroxene and either garnet or spinel, clinopyroxene and garnet are the only minerals that contain
significant concentrations of REEs. In spinel peridotites
clinopyroxene is the dominant reservoir of REEs and in
garnet peridotite both clinopyroxene and garnet contain
significant amounts of REEs. Garnet has a strong preference
for the heavy REEs (HREEs), which means that clinopyroxene coexisting with garnet is depleted in these elements
(FIG. 6). This leads to the question: if the residue of partial
melting contains garnet, can its presence in the source
region be inferred by examining the REE pattern of the
eruptive products?
Garnets trapped as inclusions in diamonds are generally
attributed to either eclogitic or peridotitic paragenesis
(E-type or P-type, respectively), although there are rare
websteritic varieties. E-type garnet is Cr poor (<0.4 wt%
Cr2O3) and Ca rich, and has lower and more variable Mg#
(45–70). P-type garnet is Cr rich and Ca poor, and has
higher and more consistent Mg# (~80–92) (Stachel and
Harris 2009).
FIGURE 5 shows the compositions of high-pressure garnets
from diamond inclusions (Harte 2010) plotted so as to
highlight the majorite substitutions. As can be seen,
increasing majorite content leads to increasing Si4+, Ti4+
and Na + at the expense of Al3+, Cr3+ and divalent cations.
Although most of the world’s diamonds are interpreted to
have formed within the cratonic lithosphere at depths to
~250 km (Stachel and Harris 2009), comparison of majoritic
garnet inclusions with those of garnet generated in highpressure experiments (Kiseeva et al. 2013) indicates that
these particular inclusions come from deeper in the mantle
at transition zone depths (410–660 km). Moreover, findings
E LEMENTS
Compilation of majoritic garnet compositions
from inclusions in diamonds.
Note that end-member majorite would plot at x = 4, y = 4.
M = divalent cations; pfu = per formula unit.
424
FIGURE 5
D ECEMBER 2013
U4+ is closer to ro than Th4+. Therefore, in the presence of
garnet, the higher U bulk partition coefficient leads to a
melt enriched in Th compared to U. This means that the
excesses of 230 Th relative to its parent, 238U, observed in
many MORB samples are most readily explained by melting
in the garnet peridotite stability field (Beattie 1993).
Chondrite-normalized rare earth element concentrations in clinopyroxene (Cpx) and garnet (Grt) from
spinel (Sp) peridotite and garnet peridotite of Vitim Plateau
(Ionov et al. 1993)
FIGURE 6
FIGURE 7 shows garnet–liquid and clinopyroxene–liquid
partition coefficients determined in one experiment at
3.0 GPa (Pertermann et al. 2004). The partition coefficients
are given by:
[i] xtal
(Di =
), where [i] xtal and [i] melt
[i] melt
refer to weight concentrations in the crystal and melt
phases. As anticipated, the garnet–liquid partition coefficients for the heavy REEs, (those with the smallest ionic
radii) are significantly higher than the clinopyroxene–
liquid values. While the clinopyroxene–liquid partition
coefficients for all REEs are below 1, the garnet–liquid partition coefficients for the smaller HREEs are close to 10,
with the light REEs (LREEs) being progressively excluded
as the ionic radius increases. The result means that melting
in the presence of significant amounts of residual garnet
will generate liquids exhibiting depletions in the HREEs
relative to the LREEs.
The roughly parabolic relationship between the partition coefficient and ionic radius was fi rst observed in
the 1960s (Onuma et al. 1968) and quantified in terms
of strain energy by Blundy and Wood (1994). The latter
authors showed that the elastic strain energy introduced
into the crystal lattice when an ion of radius r i is partitioned into a site of radius ro can be systematized and
quantified. If the ion has radius ro then there is no strain
energy of substitution and the partition coefficient, Do, is a
maximum for ions of that charge (Blundy and Wood 1994).
The strain energy of substitution increases in an approximately parabolic way with the difference between r i and ro ,
yielding a parabolic decrease of D i with radius difference,
as observed (FIG. 7). The ro value of the garnet X-site is
smaller than the ro value of the M2 site in clinopyroxene,
and the garnet site is much “stiffer,” with a higher Young’s
modulus than the clinopyroxene site (Van Westrenen et al.
2001). This means that smaller ions prefer the garnet site
and that garnet discriminates much more strongly than
does clinopyroxene against ions that are different from ro
in radius. Thus, garnets preferentially retain the smaller
HREEs (3 + ions) and are effective at separating, or fractionating, the middle rare earths (MREEs) from one another.
Similar observations can be made for ions of other charge,
such as Th4+ and U4+, substituting into garnet and clinopyroxene. Th4+ is larger than U4+ and, in peridotitic garnet,
E LEMENTS
Application of the principles discussed above to identifying and quantifying the presence of garnet in the source
regions of igneous rocks is not always unequivocaI. If garnet
was present as a residual phase, we would expect fractionation of MREEs from HREEs, and this has frequently been
reported. For example, Shen and Forsyth (1995) showed
that Sm/Yb ratios in MORB correlate with other geochemical parameters that are believed to indicate melting in
the deeper parts of the asthenosphere. After correcting
for variable enrichment of incompatible elements, they
concluded that the areas with the strongest “garnet signature” are the areas with the thinnest crust. In these regions,
the relatively cold lithosphere inhibits shallow melting;
thus a greater fraction of the observed MORB was, they
argued, produced deep, in the garnet stability field (Shen
and Forsyth 1995), and melting ceased relatively deep. In
contrast to this result, Bourdon et al. (1996) showed that
some MORBs have 230 Th/238U activity ratios above unity,
implying melting in the presence of residual garnet. More
importantly, these high 230 Th/ 238U activity ratios were
observed at ridges with thick crust (Bourdon et al. 1996).
This identification of the isotopic “garnet signature” in
areas with thick crust is directly opposite to the inferences
of Shen and Forsyth (1995), who identified the “garnet
signature” in areas of thin crust. One possible way to reconcile these confl icting results is through the melting of a
heterogeneous source that is a mixture of spinel peridotite
and garnet pyroxenite (Hirschmann and Stolper 1996), but
it is apparent that there is no clear unequivocal interpretation of the “garnet signature” in mid-ocean ridge basalts.
Experimentally measured crystal–melt partition coeffi cients for garnet (Grt) and clinopyroxene (Cpx) from a
single experiment at 3 GPa plotted as a function of ionic radius
(Pertermann et al. 2004). Solid lines are lattice strain fits to the
data using the model of Blundy and Wood (1994). The filled circles
indicate the “best-fit” radius ro.
425
FIGURE 7
D ECEMBER 2013
CONCLUDING REMARKS
We have shown that garnet is an extremely important phase
at high pressures and temperatures in the mantle. Garnet
is denser than other low-pressure ferromagnesian silicates
because of the relatively high coordination numbers of Mg2+
and Fe2+ (both 8-coordination) and Al3+ (6-coordination).
This means that once garnet becomes stable in peridotite
(>85 km depth), it increases in modal amount, from ~10%
to ~35%, by dissolving the low-pressure pyroxenes (which
have the same number of cations per 12 oxygens) into its
structure. A similar but more pronounced change in modal
proportions is found in eclogite compositions in which the
garnet–clinopyroxene assemblage becomes stable at about
70 km depth (Green and Ringwood 1967) and progressively changes from ~30% to ~95% garnet as clinopyroxene
dissolves into the garnet and the latter becomes increasingly majoritic in composition.
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The relatively high density and cation coordination
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We thank E. Baxter, W. Griffi n and J. Van Orman for
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for the use of their images. We also acknowledge support
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D ECEMBER 2013