Earth and Planetary Science Letters, 89 (1988) 123-145
Elsevier Science Publishers B.V., Amsterdam - Printed in The Netherlands
123
[21
Experimental determination of element partitioning
between silicate perovskites, garnets and liquids:
constraints on early differentiation of the mantle
T. Kato, A.E. Ringwood and T. Irifune *
Research School of Earth Sciences, Australian National University, Canberra, A.C.T. 2601 (Australia)
Received October 23, 1987; revised version accepted February 25, 1988
Distributions of major (Si, A1, Ca, Mg, Fe) and minor elements (Cs, Rb, K, Na, Sr, Ba, Pb, Cr, Sc, Y, Yb, Ho, Sm,
Nd, La, Ti, Zr, Hf, U, Th and Nb) between majorite garnet, MgSiO3-perovskite, CaSiO3-perovskite and coexisting
liquids have been determined experimentally in ultrabasic and basic compositions at pressures of 15-25 GPa and
temperatures of 1400-2200 o C using an MA-8 apparatus. The results demonstrate the capacity of silicate perovskites to
accept a wide range of normally "'incompatible" elements possessing diverse ionic radii and charges into their crystal
structures. All of the minor elements investigated are preferentially partitioned in Mg- or Ca-perovskites as compared
to majorite garnet under subsolidus conditions. In subsolidus assemblages containing garnet (gnt) and Mg-perovskite
(mpv), So, Ti, Zr, Hf and Nb are preferentially partitioned into perovskite with Dmp,,/gm values of 3-14 whereas there
is little fractionation of rare earth elements (REE), all of which have D,.,p,,/gm values of 1-2. In subsolidus assemblages
containing Mg-perovskite and Ca-perovskite, Sc, Nb, Zr, and Hf are preferentially partitioned into Mg-perovskite
whereas all other minor elements are strongly partitioned into Ca-perovskite. Majorite garnet which appears on the
liquidus in ultrabasic compositions at - 1 6 GPa and - 2 1 0 0 ° C is enriched in AIzO 3 compared to the coexisting
ultrabasic liquid by factors of 2-3 and is depleted in CaO and TiO 2 by similar factors. The partition coefficients
D~t/liq which are applicable to liquidus majorite garnet in ultrabasic compositions are Sc (1.7), Yb (1.4), Y (1.3), Hf
(0.8), Zr (0.6), Sm (0.2) and less than 0.1 for K, St, La, Th, U, Ba, Rb and Cs. Partition coefficients (Dmpv/liq) between
Mg-perovskite and ultrabasic liquids were obtained by combining garnet/liquid and subsolidus garnet/Mg-perovskite
partition data obtained at similar temperatures. It was found that So, Hf, Zr, Ti and HREE are enriched in liquidus
Mg-perovskite with Dmpv/li q values ranging between 2 and 14. Partition coefficients between Ca-perovskite (cpv),
garnet and liquid were determined in a basaltic composition at - 20 GPa and - 2000 ° C. Ocpv/li q values for U, Th
and Pb are remarkably high (10-20) and are also high for REE and Sr (2-6) whereas K, Rb, Cs and Ba behave as
incompatible elements with D~pv/liq values less than 0.5.
The results provide strong constraints upon hypotheses which maintain that the mantle experienced extensive
melting during formation of the earth followed by fractional crystallization-differentiation processes involving majorite
garnet a n d / o r MgSiO3-perovskite. The measured partition coefficients show that fractionation of garnet would be
accompanied by sharp decreases of A I / C a and S c / S m ratios in residual liquids. Likewise, fractionation of Mg-perovskite would cause marked variations of L u / H f and especially of S c / S m and H f / S m ratios in residual liquids. The
observation that the present upper mantle possesses near-chondritic relative abundances of Ca, A1, Sc, Yb, Sm, Zr and
Hf categorically excludes models which propose that the bulk mantle once possessed chondritic relative abundances of
Mg, Si, AI, Ca and other lithophile elements, but has differentiated to form a perovskitic lower mantle ( M g / S i < 1) and
a peridotitic upper mantle (Mg/Si > 1). The chemical and isotopic compositions of ancient ( - 4 . 2 Ga) Western
Australian zircons imply that even if the mantle were extensively melted and differentiated around 4.5 Ga, it must
somehow have become effectively rehomogenised by solid state convection by 4.2 Ga. Since this scenario appears
implausible on dynamic grounds, it is concluded that the mantle probably did not experience extensive melting during
the formation of the Earth.
1. Introduction
The bulk composition of the upper mantle has
been estimated by a number of workers, based
* Present address: Department of Geology and Mineralogy,
Hokkaido University, Sapporo 060, Japan.
0012-821X/88/$03.50
© 1988 Elsevier Science Publishers B.V.
upon the complementary geochemical relationships which exist between various classes of
basaltic magmas and their respective peridotitic or
dunitic residues (e.g. [1-4]). A second method o f
estimating mean upper mantle composition has
been based upon the compositions of the most
"primitive" lherzolite xenoliths derived from the
124
upper mantle [1,5-7]. Both methods have yielded
compositions which are in satisfactory agreement.
An important conclusion deriving from the above
investigations is that many lithophile, involatile
elements, e.g. Mg, Ca, A1, Ti, Y, Sc, heavy and
intermediate REE, Zr and Hf are present in approximately chondritic relative abundances in the
upper mantle. However, there is some debate about
the present Ca/A1 ratio of the upper mantle.
Palme and Nickel [8] argue that this value might
be up to 15% higher than the chondritic ratio
whereas other studies of the compositions of
high-temperature peridotites [9] and Iherzolite
xenoliths [7] tend to suggest a chondritic Ca/A1
ratio for the upper mantle. Final resolution of the
issue will require further data. We nevertheless
regard variations in the Ca/A1 ratio (and in the
ratios of other members of the above group of
elements) that are within +15% of chondritic
values as being consistent with our interpretation
of "approximate" chondritic ratios prevailing in
the upper mantle, and more specifically, in those
regions which gave rise to mid-ocean ridge basalts
(MORBs) and high-temperature peridotites.
The geochemical resemblances between modern
MORBs and peridotites and their ancient counterparts indicates that this overall near-chondritic
abundance pattern for lithophile, involatile elements has prevailed over an extensive period of
Earth history extending back at least to 3.8 Ga
[10,11]. The discovery of a suite of ancient
4.10-4.28 Ga zircons in two localities in Western
Australia, Mt. Narryer and Jack Hills [12,13], also
provides evidence of "primitive" ratios of certain
key elements at this early stage. Kinny [14] demonstrated that the Mr. Narryer zircons were derived, ultimately, from a source region possessing
an approximately chondritic L u / H f ratio.
The conclusion that the present upper mantle
has near-chondritic relative abundances of many
involatile lithophile elements and that some of
these ratios have prevailed throughout most of
geological time, provides powerful constraints for
hypotheses of the early development of the Earth.
However, there is an important exception to the
pattern described above. The S i / M g (atomic) ratio
of the upper mantle is close to 0.78 (e.g. [3]) which
is substantially higher than the "primordial"
S i / M g ratio of 0.95 displayed by C1 chondrites.
The implied deficiency of silica in the upper man-
tle has been reconciled with the chondritic earth
model via two types of hypotheses. The first proposes that silicon (as SiO) was preferentially lost
by volatilization in the solar nebula prior to accretion of the Earth [15]. Some support for this
hypothesis is provided by the observation that the
S i / M g ratios of different classes of chondrites
vary substantially.
The second hypothesis maintains that the
S i / M g ratio of the bulk mantle is in fact chondritic
(e.g. [16,17]). Mass balance considerations accordingly imply that the S i / M g ratio of the lower
mantle is substantially higher than that of the
upper mantle. This chemically stratified structure
would require the upper mantle to be dominated
by minerals with orthosilicate (M2SiO4)
stoichiometry such as olivine and silicate spinel,
whereas the lower mantle would possess metasilicate (MSiO3) stoichiometry and would be comprised almost exclusively of MgSiO3-perovskite
(e.g. [18]).
The only process so far proposed which could
cause this inferred gross chemical zonation is some
form of crystallization--differentiation, following
from extensive or complete global melting of the
Earth early in its history. Two classes of gross
differentiation processes have been proposed.
According to one, the Earth experienced extensive
partial melting down to a depth of about 700 km
during, or soon after its formation. Crystallization-differentiation controlled by the separation
of majorite garnet resulted in a silica-enriched
basalt layer (the transition zone a n d / o r lower
mantle) and an overlying silica-depleted upper
mantle of pyrolite composition (e.g. [19]). Experimental support for this hypothesis is based on the
observation that majorite garnet is the high-pressure liquidus phase in both chondrite and pyrolite
model mantle compositions above 14 GPa and
that the solidus-liquidus melting interval is
surprisingly narrow at these pressures [20,21]. The
pyrolite upper mantle is thus interpreted as a
partial melt from a chondritic bulk mantle composition, leading to a silica-enriched transition
zone a n d / o r lower mantle. Ringwood [9] noted
that alternative explanations of the above experimental evidence are possible. For example, the
observations may simply reflect the circumstances
that olivine possesses a high melting point but a
low melting point gradient (dP/dT), whereas this
125
situation is reversed in the cases of pyroxenes and
garnet. Accordingly, melting curves of olivine and
those of garnet/pyroxene will tend to converge at
pressures of a few GPa leading to a narrow melting interval for pyrolite and to majorite being its
liquidus phase.
According to the second model [22,23], the
entire mantle experienced extensive to complete
melting during formation of the Earth. Extensive
crystal-liquid fractionation led to the formation
of a lower mantle predominantly composed of
Mg-perovskite, and an upper mantle of pyrolite
composition. This model is based upon experimental results which show that the liquidus phase
in primitive mantle composition below 700 km is
Mg-perovskite [241.
The objective of the present investigation has
been to test the above hypotheses of gross mantle
differentiation controlled by the separation of perovskite or majorite. We have carried out a series of
high-pressure experiments aimed at determining
the partition behaviour of a wide range of major
(Si, A1, Ca, Mg, Fe) and minor elements (Cs, Rb,
K, Na, Ca, Sr, Pb, Ba, A1, Cr, Sc, Y, Yb, Ho, Sm,
Nd, La, Ti, Zr, Hf, U, Th, Nb) in and between
majorite garnet, Mg-perovskite, Ca-perovskite and
the liquids from which these phases crystallize. It
was hoped that the experiments would provide
strong constraints on the extent to which the
present composition of the upper mantle might
have been influenced by prior differentiation
processes involving crystallization of majorite _+
perovskite(s).
Preliminary descriptions of some of the experimental results and their implications have been
published by Kato et al. [25,26].
2. Experimental procedures
2.1. Starting materials
Several different starting compositions were
prepared for use in the present investigation (Table 1). The compositions were selected in some
cases to enhance the crystallization fields of particular phases near the liquidus and, in other
cases, to promote sub-solidus grain growth and
equilibration. Three of these were based on pyrolite--PA representing a standard pyrolite model
composition whilst PB was derived by subtracting
the olivine component from pyrolite and PC by
removing most CaO from pyrolite in order to
suppress the crystallization field of CaSiO3-perovskite. Another composition, CA, represented a
model mantle composition derived from the silicate phase of C1 chondrites. In addition, two
Mg-rich komatiite compositions KA and KB were
studied, one of which (KA) had previously been
studied by Kato et al. [25]. In addition, experiments were also carried out on a primitive MORB
composition.
The compositions were prepared as homogeneous glasses according to standard procedures
previously used in this laboratory. The glasses
were then finely ground and divided into subgroups. Sets of minor elements were introduced as
oxides in desired proportions (typically at levels of
about 1000 ppm) and thoroughly mixed with the
glass powders. The mixtures were then re-melted
to form homogeneous glasses using rhenium sample holders in an argon atmosphere. The maximum number of additional minor elements added
to any glass was seven, in order to minimise
analytical interferences. The final compositions of
the sets of glasses as analysed by electron microprobe are shown in Table 1.
In most runs carried out above the solidus,
glass starting materials were used. It was found by
experience that equilibrium was quickly achieved
and that crystal sizes were large enough to permit
electron-probe microanalysis. Below the solidus
however, crystals were often too small for satisfactory electronprobe analysis (and difficulties in
achieving equilibrium were sometimes encountered). Accordingly, the glasses used in these runs
were converted to amphibolitic starting materials
containing 1-2% H 2 0 following procedures
described by Irifune and Ringwood [27]. The presence of water was found to enhance subsolidus
grain growth and to facilitate equilibration.
2.2. High-pressure and temperature experiments
High-pressure experiments were carried out
using a modified split-sphere multi-anvil apparatus installed in a 1200-ton uniaxial press [28].
Tungsten carbide cubic anvils with truncated
corners and edge lengths of 3.5 mm were used at
pressures up to 20 GPa whilst the truncated edgelengths were reduced to 2.0 mm in runs carried
out at higher pressures. Pressure calibrations of
the system at room temperature and at elevated
temperatures have been described by Ohtani et al.
126
TABLE 1
Compositions of starting materials
(a) Major element compositions (wt.%)
Komatiite
Pyrolite
KA
PA
SiO 2
TiO 2
AI203
Cr203
FeO
MgO
CaO
Na 2°
46.8
0.35
7.3
0.69
6.1
32.4
5.5
0.63
Sum
99.8
KB
45.9
0.38
5.8
0.42
10.6
31.1
5.5
0.34
100.0
PB
43.9
0.22
4.7
0.41
8.3
38.1
3.4
0.42
99.5
51.8
0.46
11.4
0.92
3.1
23.1
9.5
0.94
101.2
PC
44.5
0.24
4.9
0.43
8,6
38.2
0.50
0.42
97.9
Chondrite
Basalt
CA
BA
49.9
0.22
3.7
0.21
7.2
36.0
2.8
0.23
100.3
50.4
0.57
16.1
0.13
7.7
10.5
13.1
1.9
100.4
(b) Minor element concentrations
Major element
set
Minor element
set
Element, concentration (ppm)
KA
KB
KB
(1)
(2)
(3)
Sc, 520; La, 510; Sm, 430; Yb, 700
K, 2000; Sr, 640; Y, 820; Hf, 840; Zr, 830; Nb, 850
K, 2000; Sr, 1900; Th, 1400; U, 600; Ba, 1500; Cs, 160; Rb, 520
PA
PA
(4)
(5)
K, 900; Sc, 1350; Sm, 930; Yb, 1350; Y, 1250; Hf, 1750; Zr, 1800
K, 1300; Sc, 1700; Sm, 1800; Yb, 1700; Hf, 2300; Nb, 1700; St, 1800
PB
PB
PC
(1)
(2)
(4)
K, 1000; Sc, 2000; La, 1600; Nd, 1700; Sm, 1700; Ho, 2000; Yb, 1700
Sr, 2100; Y, 1400; Hf, 1800; Zr, 1900; Nb, 1800
K, 1000; Sc, 1400; Sm, 950; Yb, 1350; Y, 1300; Hf, 1850; Zr, 1950
CA
(4)
Sc, 1000; Sm, 600; Yb, 950; Y, 750; Hf, 1300; Zr, 1500
BA
BA
BA
(1)
(2)
(3)
Sc, 1800; La, 1500; Nd, 1500; Sm, 1800; Ho, 1800; Yb, 1500
Sr, 1100; Y, 1300; Hf, 820; Zr, 1300; Nb, 1200; Pb, 1000
St, 1400; K, 300; Th, 1600; U, 1000; Ba, 1300; Cs, 300; Rb, 300
KA = komatiitic composition, pyrolite minus 40% olivine [25]; KB = average chondritic komatiite [40]; PA = pyrolite composition
[3]; P B = pyrolite minus all olivine composition [41]; PC = pyrolite composition depleted in CaO [3]; CA = model mantle
composition derived from CI chondrites; BA = primitive MORB composition [27].
[28]. At the very high temperatures ( > 2000 ° C)
used in many of the present experiments, pressure
errors are probably larger than those obtained by
Ohtani et al., and may be in the vicinity of + 10%
of nominal pressure.
The pressure medium used in the experiments
consisted of a segmented octahedron composed of
(Mg t zCox)O (x = 0.05 - 0.10) solid solutions.
The presence of CoO reduces the radiative thermal conductivity of the pressure medium and improves the stability and efficiency of the heater.
The latter consists of two strips of a mixture of
powdered tungsten carbide and diamond. The
sample is encapsulated in a platinum cylinder
welded at both ends to prevent loss of mobile
elements during the experiment. In runs carried
out using 3.5-mm truncated anvils, the heater provides a uniform temperature distribution, and the
temperature of the sample can be measured quite
accurately by means of Pt-Ptl0Rh or W3ReW25Re thermocouples placed on either side of it
[29]. However, short circuits were found to occur
frequently when this configuration was used in the
smaller cells in conjunction with 2.0-mm truncated
anvils at pressures above 20 GPa owing to narrow
gaps ( < 0.1 mm) between the anvils. Accordingly,
in these runs, the heater was calibrated by determining the relationship between power input
and temperature at lower pressures, the latter being
measured by a thermocouple placed at the hot
127
spot. In subsequent runs, the temperatures of samples placed at hot spots were estimated from the
power input calibration curves. In runs carried out
above 2000 ° C, this procedure may yield substantial errors in temperature, by as much as _+200 o C.
However, relative uncertainties in sets of closely
spaced runs are believed to be substantially
smaller. Sometimes, in runs carried out above
2000 ° C, the heater displayed unstable behaviour
with one heater strip becoming much hotter than
the other. This was usually recognisable from the
resistance change of the heater and from inspection of the heater after the run. Results from these
runs were discarded.
In the course of an experiment, pressure was
applied first by loading to the tonnage required
for the desired pressure, and the sample was then
heated over a period of about 10 minutes to the
desired temperature. Heating durations were typically about 10 minutes for subsolidus runs and
1-3 minutes for above-solidus runs. The experiments have not been reversed. However, we believe that (local) equilibrium was closely approached in the great majority of runs. This opinion is based upon considerable experience in the
use of the apparatus on different systems and the
consistency obtained between the results of runs
carried out over a wide range of times and using
differently prepared starting materials. After completion of the run, the sample was quenched by
shutting o f f the power supply and pressure was
then released slowly, over 12 hours, in order to
minimize the frequency of gasket blowouts. The
recovered specimens were prepared as polished
thin sections for optical and electronprobe examination. Small samples were crushed for examination by powder X-ray diffraction.
3. Methodology
Phase identifications in the run products were
made by combining textural observations under
optical and electron microscopes with the results
of powder X-ray diffraction and electron microprobe analyses. The latter were made using a
Camebax Microbeam instrument in the wavelength dispersive mode. The electronprobe was
operated at an acceleration voltage of 25 kV and a
beam current of 30-50 nA. Eight major elements
(Si, Ti, A1, Cr, Fe, Mg, Ca and Na) were measured
simultaneously with up to 7 dopant minor elements under a fixed beamspot position. Potassium
and Sc abundances were measured using their K a 1
peaks whilst La, Nd, Sm, Ho, Yb, Nb, Zr, Y, Sr,
Hf, Th, Ba, Rb, Cs and Pb were determined from
L a 1 peaks. Uranium was determined from the
M a a peak. Corrections for peak overlap onto
background were made for Ca on Sc, Si on Sr, Ti
on Ba, Ba on Ti. Counting times corresponding to
required detection limits (typically 100-300 ppm)
are 200-600 seconds. Uncertainties associated with
statistical scatter in peak and background counts
are satisfactorily small for Sc, La, Sm, Yb, Hf, K
and Ba ( < 100 ppm), but are slightly larger for Y,
Zr, Sr, Nb, Rb, Cs, Th, U and Pb (100-300 ppm).
Most runs carried out under subsolidus conditions produced assemblages consisting of wellcrystallized garnets a n d / o r Mg-perovskites with
crystal sizes of 5-20 /zm. In runs carried out
between solidus and liquidus, garnet usually occurred as well formed crystals, typically 10-100
/~m across. Ca-perovskite retrogressively transformed to glass on release of pressure as reported
earlier by Ringwood and Major [30] but was readily recognizable and could be analysed in most
cases. Ca-perovskite mostly occurred in areas with
smaller dimensions (3-5/~m) and sometimes posed
analytical difficulties with total oxides typically
summing to only 80-95%. It is believed that the
low totals are due to microcracks and microporosity developed during quenching and retrogressive
transformation. In these above-solidus runs,
former liquid was recognisable as areas of fibrous
quench crystals. In cases where sufficient segregation of liquid had occurred, it was possible to
determine the bulk composition of the quenched
material from broad beam analyses of multiple 25
/~m x 25 /zm areas. However, this was possible in
only a limited number of runs because of the
presence of intergrown primary crystals.
The results of many runs carried out in the
melting region required careful interpretation. In
the compositions investigated, the melting interval
between solidus and liquidus was at most about
200 o C, which is comparable with the uncertainty
of temperature measurement using the 2.0-mm
truncated anvil system. Consequently, a considerable number of runs were replicated with the
objective of obtaining a reasonable sample of experimental points between the solidus and liqui-
128
dus. This technique worked satisfactorily in the
cases of runs aimed at establishing garnet-liquid
relationships which utilized 3.5-mm anvils and a
pressure cell which generally behaved in a stable
manner. However, it was less successful when
2-ram anvils were used, generally at 25 GPa and
above 2000 ° C which were the conditions necessary in order to study equilibria between fiquids
and Mg- and Ca-perovskites. One problem was
that in many of these runs, a temperature gradient
was established across the charge because of slight
variations in the performance of the two strip
heaters. In a number of runs, this resulted in one
side of the charge being essentially subsolidus and
the other side being completely melted. Alternatively, the textural relationships were sufficiently
ambiguous so that the results could not be confidently interpreted. In either event, the results of
such runs were discarded. Altogether, thirty runs
were attempted with the aim of determining Caand Mg-perovskite/liquid partition equilibria. A
number of these failed because of blowouts, with
consequent anvil destruction, so that the exercise
was both time-consuming and expensive. Of these
runs, only one was completely successful, whilst
useful results were obtained from additional runs.
All three of these runs were on Ca-perovskite/
liquid systems in MORB bulk composition BA.
Because of these experimental difficulties, the distributions of minor elements between Mg-perovskite and ultrabasic liquids were determined
indirectly. A satisfactory set of partition coefficients Dgnt/liq for majorite~ garnet/liquid equilibria was first determined. Runs were then carried
out under subsolidus conditions in compositions
where Mg-perovskite coexisted with majorite
garnet possessing a composition similar to those
of garnets which had equilibrated with ultrabasic
liquid, thereby providing appropriate partition
coefficients between Mg-perovskite and garnet
(Dmpv/gnt). The partition coefficient (D) for a
given element between Mg-perovskite and ultrabasic liquid was then obtained from the relationship:
Ornpv/liq = Ompv/gnt X Ognt/liq
4. Results
4.1. Ultrabasic compositions
Melting and subsolidus phase relationships ob-
served in ultrabasic compositions KA, KB, PA
and CA are given in Table 2a and subsolidus
phase relationships in modified pyrolite compositions PB and PC are given in Table 2b.
Under subsolidus conditions at 16 GPa, the
ultrabasic compositions KA, KB, PA and CA
crystallized completely to assemblages of majorite
garnet and olivine. The solidus temperatures for
both komatiite compositions KA and KB lie between 1900 and 2000°C in the pressure interval
12-18 GPa. (Loss of iron to the platinum capsules
causes the final compositions of KA and KB to be
quite similar.) Majorite becomes the liquidus phase
above 16 GPa in these compositions and the melting interval is smaller than about 100 o C. Similar
results have been found by other workers in a
range of ultrabasic compositions [20,21,24,25,34].
At 16 GPa, the melting intervals of the pyrolite
and chondrite compositions PA and CA occur at
significantly higher temperatures ( - 2200 o C) than
were observed in the komatiites.
Successful analyses of coexisting majorite garnet
and liquid require relatively large degrees of partial melting. The results on twelve such runs (six
for KA-KB, 5 for PA and 1 for CA) are shown in
Table 3. These runs displayed relatively homogeneous distributions of primary majorite crystals
throughout the sample demonstrating that fairly
uniform temperature gradients had been achieved
(Fig. la).
Subsolidus phase relationships between coexisting majorite garnet, Mg-perovskite and Ca-perovskite were determined in four runs on modified
pyrolite compositions PB and BC and are recorded in Table 2b. Runs conducted at 25 GPa
and 1400 ° C consisted of the above three phases.
Crystal sizes of majorite and Mg-perovskite were
5-10 /xm, permitting accurate electronprobe
analyses which are presented in Table 4. Ca-perovskite grains (retrogressively transformed to glass)
were smaller, ranging up to a few microns, and
analyses may have been slightly biased by minor
beam overlap. Run 259 at higher temperatures
(1900 o C, 25 GPa) yielded a well-crystallized subsolidus assemblage of garnet grains ( - 50 ~m) +
Ca-perovskite ( - 1 0 /~m) but Mg-perovskite was
not found. Because of their crystal size, accurate
analyses of coexisting garnet and Ca-perovskite
were readily obtained. A further run (295-2) on
pyrolite depleted in CaO (PC) at 1900 ° C 26 GPa
129
yielded well-crystallized Mg-perovskite + garnet,
as shown in Fig. lb, and Ompv/gnt values for
several key elements were obtained.
4.2. MORB composition BA
Phase relationships displayed by the primitive
MORB basalt composition (BA) are given in Ta-
TABLE 2
Conditions and results of high-pressure runs
Run
No.
Pressure
(GPa)
Temperature
( ° C)
Heating time
(minutes)
Results
(a) Assemblages containing majorite garnet, + o l i v i n e + m o d i f i e d spinel+liquid in komatiite compositions K A and KB, pyrolite
composition PA, and chondrite composition CA
Composition KA-(1)
213-1
201-1
199-1
204-1
182 a
231-1
12
12
12
12
16
16
1900
2010
2100
2150
2090
2100
5.0
1.0
1.0
1.5
2.0
1.0
O1, Gt,
O1, Gt,
O1, Gt,
O1, L
Gt, L
Gt, O1,
CPx
L
L
12
12
12
12
15
16
1900
2010
2100
2150
2100
2100
5.0
1.0
1.0
1.5
1.0
1.0
O1, Gt,
O1, Gt,
O1, Gt,
O1, L
Gt, O1,
Gt, L
15
18
2100
2100
1.0
4.0
Gt, L
Gt, L
16
16
16
16
16
16
16
16
2000
2050
2100
2100
2200
2230
2260
2280
3.0
2.0
1.0
1.0
1.0
1.0
1.0
1.0
Gt,
Gt,
Gt,
Gt,
Gt,
Gt,
Gt,
L
24.5
2200
1.0
Gt, MS, L
16
16
16
16
16
16
16
2100
2130
2170
2200
2230
2260
2280
1.0
1.0
1.0
1.0
1.0
1.0
1.0
Gt,
Gt,
Gt,
Gt,
Gt,
Gt,
L
L
Composition KB-(2)
213-2
201-2
199-2
204-2
264-1
231-2
CPx
L
L
L
Composition KB-(3)
264-2
257
Composition PA-(4)
14 b
11 b
5b
19 b
27-2
29-2
31-2
30-2
O1
O1
L
L
O, L (trace)
O1, L
L
Composition PA-(5)
293 b
Composition CA-(4)
23
25
26
27-1
29-1
31-1
30-1
O1
O1
O1
O1
O1, L
L
(b) Assemblages containing Mg-perovskite, majorite garnet + Ca perovskite _+spinel in modified pyrolite compositions PB and PC
Composition PB-(I)
215
259
25.5
25.0
1400
1900
10
3.5
Gt, MgPv, CaPv
Gt, CaPv
25.0
1400
10
Gt, MgPv, CaPv
24.9
1900
3
Composition PB-(2)
251
Composition PC-(4)
295-2
Gt, MgPv, Sp
130
TABLE 2 (continued)
Run
No.
Pressure
(GPa)
Temperature
( o C)
Heating time
(minutes)
Results
(c) Assemblages containing majorite garnet, Ca perovskite + liquid in primitive MORB composition BA
Composition BA-(1)
268
269
270
271
282 c
260
262
20.0
20.0
20.0
20.0
22.0
24.5
24.0
2000
2100
2150
2200
> 2100
1400
1900
2.0
3.0
1.0
1.0
1.0
10
1.0
Gt,
Gt,
Gt,
Gt,
Gt,
Gt,
Gt,
St
St, L (trace)
St, L (trace)
L
CaPv, L
CaPv, St
CaPv, St
20.0
22.0
24.0
25.0
2200
> 2100
1800
1400
1.0
1.0
2.0
12
Gt,
Gt,
Gt,
Gt,
L
CaPv, L
CaPv, St
CaPv, St
20.0
24.0
24.0
2200
1400
> 2100
1.0
10
1.0
Gt, L
Gt, CaPv, St
Gt, CaPv, St, L
Composition BA-(2)
273
279 c
27-6
228
Composition BA-(3)
272
253
266
Abbreviations: O1 = olivine; Gt = majorite garnet; CPx = clinopyroxene; L = liquid; MgPv = Mg-perovskite; CaPv = Ca-perovskite;
MS = modified spinel; Sp = spinel; St = stishovite.
a The same run reported by Kato et al. [25].
b "Amphibolite" starting materials were used.
c Pt capsules leaked owing to melting at one side.
b l e 2c a n d a n a l y s e s o f c o e x i s t i n g p h a s e s are g i v e n
o f t h i s c o m p o s i t i o n t o lie a t a b o u t 2100 ° C. H o w -
i n T a b l e 5. U n d e r s u b s o l i d u s c o n d i t i o n s a b o v e 15
ever, t h e d e g r e e o f m e l t i n g is s o s m a l l at 2100 a n d
2150°C that liquid compositions could not be
G P a , this c o m p o s i t i o n c r y s t a l l i z e s t o a n a s s e m b l a g e o f m a j o r i t e g a r n e t + m i n o r s t i s h o v i t e [31]. A
s e r i e s o f e x p e r i m e n t s at 20 G P a , s h o w s t h e s o l i d u s
d e t e r m i n e d i n t h e s e s a m p l e s . S t i s h o v i t e is t h e first
p h a s e t o d i s a p p e a r as m e l t i n g p r o c e e d s w i t h i n -
Fig. 1. SEM images of the run products in high-magnesian composiuons. (a) Phase assemblage of majorite garnet and liquid in run
231-2 at 16 GPa and 2100 ° C. Scale bar is 100/zm. (b) Phase assemblage of garnet, Mg-perovskite and spinel in run 295-2 and 24.9
GPa and 1900 o C. High concentrations of Zr and Hf cause Mg-perovskite grains to be bright in this image. Scale bar is 10 /zm.
131
creasing temperatures. Three runs at 2200 ° C produced assemblages consisting of large crystals of
garnet and interstitial quench liquid as shown in
Fig. 2a. Limited zoning is observed with compositional variation in C a / M g ratios being smaller
than +3%. Analyses of coexisting garnets and
quench liquids are given in Table 5a.
At pressures between 20 and 25 GPa, Ca-perovskite occurs in association with majorite garnet
and stishovite in subsolidus runs, as previously
found by Irifune and Ringwood [27]. As temperature increases above the solidus, first stishovite,
and then Ca-perovskite are consumed, leaving a
magnesian garnet as the liquidus phase. Analyses
of coexisting garnet and Ca-perovskite pairs below
the solidus are given in Table 5b.
Despite m a n y attempts (section 3) only one run
was successful in producing coexisting Ca-perovskite, garnet and liquid. This run, No. 266, is
shown in Fig. 2b. The central part consists of
garnet, Ca-perovskite and stishovite and the liquid
phase has segregated to one side of the platinum
container. This phase assemblage is considered to
result from a small degree of partial melting just
above the solidus. Compositions of perovskite,
garnet and liquid are given in Table 5c.
Partial melting occurred in two other runs at 22
G P a and > 2100 ° C. However, the platinum capsules themselves melted on one side and the liquids
were found to be significantly contaminated by
the pressure medium material. Analyses of coexisting Ca-perovskite and garnet in these runs are
given in Table 5c. These results remain useful
since estimates of Ca-perovskite-liquid partition
coefficients can be obtained from mass balance
considerations (section 5.4).
5. Discussion
5.1. Subsolidus partition relationships
This investigation has demonstrated the capacity of silicate perovskites to accept a wide range of
normally "incompatible" elements possessing diverse ionic radii and charges into their crystal
structures at concentrations exceeding 1000 p p m
and ranging up to several percent (Tables 4 and
5). Non-sillicate perovskites are also known to
display this characteristic. Because of the high
solubilities of normally incompatible elements in
silicate perovskites in relation to their geochemical
abundances, these elements should be located in
perovskite phases rather than in accessory minerals in the lower mantle.
All of the minor elements investigated are preferentially partitioned into Ca- or Mg-perovskite as
compared to majorite garnet. In mineral assemblages containing garnet and Mg-perovskite, Sc,
Hf, Zr, N b and Ti are preferentially partitioned
into perovskite with Dmpv/gnt values of 3-18,
whereas there is little fractionation of rare earths,
Fig. 2. SEM images of the run products in DSDP basalt composition (E). Scale bars are 100 ~tm. (a) Phase assemblage of majorite
garnet and liquid in run 273 at 20.0 GPa and 2200 ° C. Zoning in large garnet grains is caused by small compositional variations in
Ca/Mg. (b) Phase assemblageof garnet, Ca-perovskite,stishovite and liquid in run 266 at 24.0 GPa and > 1900 o C. Liquid phase is
observed to accumulate near the walls of the platinum capsule (at the bottom of the photograph).
132
The partitions of the large quadrivalent cations
Zr 4+, Hf 4+, U 4+ and Th4÷ are of particular
crystal-chemical and geochemical interest. It is
unlikely that these large cations are able to substitute for Si 4+ in the octahedral sites in the perovskite lattice. Kesson et al. [32] and Solomah et
al. [33] showed that in titanate perovskites and
all of which have Dmpv/gnt values around 1-2. In
subsolidus assemblages containing coexisting Mgand Ca-perovskites, Sc, Nb, Zr and Hf are preferentially partitioned into Mg- perovskite whereas
Na, K, Rb, Cs, Ca, Sr, Pb, Ba, Sc, Y, Yb, Ho, Sm,
Nd, La, Th and U are strongly partitioned into
Ca-perovskite.
TABLE 3
Compositions of coexisting garnets and liquids in ultrabasic compositions KA, KB, PA and CA (maior elements values in wt.%, trace
element values in ppm)
Composition:
KA
KB
KB
Run No.:
Conditions:
182
16 GPa, 2090 ° C
231-1
16 GPa, 2100 ° C
264-1
15 GPa, 2100 ° C
Phase:
Gt
L
Gt
L
Gt
L'
42.0
0.28
7.6
0.61
3.4
40.0
5.6
0.13
48.6
0.24
12.1
0.96
4.7
28.9
4.7
0.31
43.0
0.41
8.4
0.75
5.1
34.9
6.9
0.58
47.5
0.29
16.8
0.25
0.30
30.3
3.2
0.06
43.3
0.94
6.7
0.10
0.45
40.6
6.3
0.67
100.3
99.6
99.9
98.9
98.7
99.2
1200 (50) a
< 60
<100
1000 (50)
700
200
350
700
1150 (100)
< 100
330 (50)
1000 (100)
850
580
530
800
< 100
1000 (100)
1000 (100)
1000 (IO0)
< 100
< 100
2400 (100)
SiOz
47.1
0.15
17.3
0.93
1.6
30.0
3.2
0.06
TiO 2
A1203
Cr203
FeO
MgO
CaO
Na20
Sum
Sc
La
Sm
Yb
(100)
(100)
(100)
(100)
Sc
La
Sm
Yb
(100)
(100)
(100)
(150)
Sr
Y
Hf
Zr
Nb
K
50O (2O0)
1100 (100)
2300(200)
550 (100)
1600 (200)
Composition:
KA
KB
KB
Run No.:
Conditions:
Phase:
231-2
16 GPa, 2100 ° C
264-2
15 GPa, 2100° C
257
18 GPa, 2100 ° C
L
Gt
L
Gt
L
SiO 2
TiO 2
A1203
Cr20 3
FeO
MgO
CaO
Na 2°
48.0
0.20
12.8
0.46
4.3
29.5
4.0
0.16
46.5
0.38
7.5
0.30
6.8
30.1
6.3
0.53
46.5
0,16
18,1
0,35
0,82
29,5
2.5
0.06
49.8
0.46
8.8
0.29
2.3
32.1
5.5
0.23
48.5
0.17
15.3
0.30
0.77
31.1
3.3
0.09
50.6
0.37
7.0
0.24
2.4
31.6
6.2
0.27
Sum
99.4
98.4
98.6
99.2
99.6
98.7
300 (100)
100 (100)
< 100
< 100
< 700
< 200
< 60
1900
1500
520
170
1300
500
1400
Gt
Sr
Y
Hf
Zr
Nb
K
< 100
850 (100)
700 (100)
450 (100)
< 100
100 (100)
1500
620
820
1200
700
(200)
(150)
(150)
(300)
(200)
1100 (200)
Sr
Ba
Rb
Cs
Th
U
K
(150)
(100)
(50)
(100)
(300)
(100)
(200)
Sr
Ba
Rb
Cs
Th
U
K
<
<
<
<
<
260 (50)
200 (50)
100
100
700
200
100
1600 (200)
1800 (200)
360 (100)
21o (lOO)
lOOO(3oo)
600 (200)
1300 (200)
133
Table 3 (continued)
Composition:
PA b
PA b
PA b
Run No.:
Conditions:
5
16 GPa, 2100 o C
9
16 GPa, 2100 ° C
293
24 GPa, 2200 o C
Phase:
Gt
L
Gt
L
Gt
46.0
0.23
4.6
0.35
2.7
46.7
3.6
0.42
48.8
0.06
14.0
0.97
2.3
31.5
1.8
0.06
45.5
0.24
5.7
0.46
6.1
38.0
3.6
0.45
99.5
99.6
50.1
0.11
12.0
0.80
2.0
33.9
2.0
0.05
SiO 2
TiO 2
A1203
Cr203
FeO
MgO
CaO
Na20
Sum
101.0
Zr
Hf
Y
Sc
Sm
Yb
K
1800
1400
1350
2600
300
2200
< 100
98.7
(300)
(300)
(200)
(200)
(100)
(200)
3000
1950
1200
1450
1000
1700
700
(200)
(200)
(100)
(100)
(100)
(200)
(100)
1400 (100)
1700 (100)
1400 (100)
gr
Hf
Y
Sc
Sm
Yb
K
2400 (300)
2400 (200)
1200 (100)
26o0 (200) 1600(2oo)
250 (50) 1100(100)
22OO (2OO) 1700(2OO)
< 100
800 (100)
L
51.3
0.15
8.5
0.68
4.5
30.9
4.3
0.35
42.0
1.2
3.5
0.40
7.5
30.8
7.0
0.75
100.8
93.5
Sr
Nb
Hf
200 (200)
300 (200)
2500 (300)
Sc
Sm
Yb
3100 (3o0)
55O (5O)
26OO(2OO)
18000 (8OOO)
20000 (10000)
5000 (1000)
(500)
2300
7000 (2000)
(300)
2300
7500 (1000)
K
600 (200)
Composition:
PA
PA
CA
Run No.:
Conditions:
29-2
16 GPa, 2230 ° C
31-2
16 GPa, 2260 ° C
31-1
24GPa, 2260 ° C
Phase:
Gt
L
Gt
L
Gt
L
SiO 2
48.4
0.12
12.4
0.80
3.2
31.4
2.4
< 0.10
45.0
0.23
5.0
0.37
6.0
38.5
3.4
0.30
48.1
0.09
14.0
0.99
2.2
30.0
2.1
< 0.010
46.7
0.22
4.9
0.40
3.6
39.0
3.2
0.35
52.8
0.10
10.8
0.53
2.7
33.1
1.3
< 0.10
50.5
0.22
4.2
0.23
4.5
36.8
2.6
0.20
98.7
98.8
98.4
101.3
99.2
ZiO 2
A1203
Cr203
FeO
MgO
CaO
Na20
Sum
Zr
Hf
Y
Sc
Sm
Yb
K
1900(+300)
1600 (200)
2000 (200)
2600 (200)
300(+100)
2100
(200)
< 100
2600(+400)
2400 (100)
1400 (100)
1500 (100)
1100 (100)
1600 (100)
680
(50)
99.8
Zr
Hf
Y
Sc
Sm
Yb
K
2000
1400
1900
2700
200
2500
< 100
(300)
(100)
(300)
(150)
(100)
(200)
2800
1700
1400
1400
900
1600
700
(300)
(300)
(100)
(50)
(100)
(100)
(200)
Zr
Hf
Y
Sc
Sm
Yb
1100
700
800
1500
< 150
1200
(200)
(200)
(100)
(100)
(200)
1600
1100
700
900
500
900
(200)
(100)
(100)
(50)
(100)
(100)
a Figures in brackets indicate analytical uncertainties.
b "Amphibolite" starting materials are used.
zirconolite, U 4+ and Th 4+ mainly replace Ca 2+ in
8-12 fold sites and that electroneutrality is
achieved by substitution of Mg 2+ or 2(A13 +, Ti3 +)
for Ti 4+ in the octahedral sites. Thus, these elements are believed to be incorporated in the crystal
structure as inverse perovskite components, e.g.
U M g O 3. It seems likely that in the present experi-
ments, Zr 4+ and Hf 4+ preferentially replace Mg 2+
and Ca 2+ in Mg- and Ca-perovskites and that
charge compensation is achieved via replacement
of either 2A13+ or Mg 2+ cations for Si 4+. Likewise, U 4+ and Th 4+ are believed to enter C a 2+
sites in Ca-perovskite.
The distributions of Zr and Hf between majorite
5000
(500)
100
(50)
150
(50)
300
(30)
250
(50)
250
(50)
850
(100)
99.8
1300
(100)
< 80
SiO 2
TiO 2
A1203
Cr=O3
FeO
MgO
CaO
Na20
Sum
K
Yb
Ho
Sm
Nd
140
(50)
200
(50)
330
(50)
200
(50)
500
(50)
100.0
46.6
0.06
19.7
1.6
2.8
24.6
3.7
0.66
Phase:
La
56.3
1.2
4.9
0.5
2.8
32.8
1.4
0.13
Gt
R u n No.:
Conditions:
Sc
215
25.5 GPa, 1400 ° C
MgPv
PB
Composition:
1400
(100)
6800
(800)
7500
(1000)
7500
(1000)
8500
(1000)
7500
(1000)
2200
(2OO)
100.0
54.8
2.0
3.6
0.4
0.4
3.6
34.7
0.47
CaPv
< 100
250
(100)
600
(100)
150
(50)
Sm
Ho
K
Yb
< 80
1700
(100)
< 80
Nd
La
Sc
100.0
47.4
0.03
16.6
1.1
1.8
24.5
5.9
0.62
Gt
2100
(100)
9000
(1000)
9000
(2000)
13000
(3000)
13000
(3000)
10000
(2000)
4500
(50O)
93.7
51.7
1.2
4.2
0.38
0.66
3.1
31.8
0.61
CaPv
259
25.0 GPa, 1900 ° C
PB
Nb
Zr
Hf
Y
Sr
1500
(500)
600
(200)
880
(200)
850
(100)
500
(100)
100.2
48.2
0.19
16.6
1.3
2.1
24.9
5.5
1.4
Gt
500
(300)
300
(200)
5300
(800)
4300
(500)
5000
(500)
99.7
52.5
1.3
5.7
0.72
3.9
33.3
2.1
0.16
MgPv
251
25.0 GPa, 1400 ° C
PB
3500
(500)
8500
(1000)
2200
(500)
3500
(1000)
1200
(300)
100.0
53.9
1.9
4.6
0.51
0.92
4.7
32.6
0.84
CaPv
K
Zr
Hf
Y
Yb
Sm
Sc
5900
(300)
1000
(200)
3800
(500)
3900
(1000)
22000
(5000)
25000
(5000)
350
(IO0)
95.7
48.5
1.3
2.1
0.24
8.0
34.8
0.55
0.13
MgPv
2000
(300)
700
(300)
2300
(500)
1800
(500)
1200
(300)
1700
(300)
600
(2OO)
100.2
45.0
0.20
9.8
0.74
8.0
34.7
1.2
0.64
Gt
295-2
24.9 GPa, 1900 ° C
PC
Compositions of coexisting garnets, Mg-perovskites and Ca-perovskites, in modified pyrolite compositions PB and PC (major element values in wt.%, trace element values in
ppm)
TABLE 4
135
TABLE 5
Compositions of coexisting garnets, Ca-perovskites and fiquids in primitive M O R B system BA (major element values in wt.%, trace
element values in ppm)
(a) Compositions of coexisting garnets and liquids
R u n No:
Conditions:
271
20 GPa, 2200 ° C
Phase:
Gt
L
Gt
SiO 2
TiO 2
A1203
Cr203
FeO
MgO
CaO
Na20
45.4
0.16
20.2
0.09
0.82
25.0
7.4
0.27
54.2 (2.0)
1.9 (0.5)
8.4 (1.5)
0.08 (0.03)
2.0 (0.5)
13.7 (3.0)
16.8 (2.0)
2.6 (1.0)
44.8
0.17
19.9
0.05
0.51
25.5
7.3
0.25
Sum
99.4
99.8
98.5
100.0
< 100
320 (70)
700 (100)
420 (50)
< 100
2200
1000
2000
2000
Sc
La
Nd
Sm
Ho
Yb
2000
< 100
< 100
320
1300
1800
(200)
(50)
(100)
(200)
273
20 GPa, 2200 ° C
1400
5000
6000
5000
3500
2500
(100)
(1000)
(1000)
(1000)
(1000)
(500)
Nb
Zr
Y
Hf
Sr
272
20 GPa, 2200 ° C
L
Gt
51.9 (3.0)
1.3 (0.5)
9.3 (3.0)
0.03 (0.03)
2.6 (1.5)
16.6 (4.0)
15.6 (2.0)
2.6 (1.0)
45.4
0.14
19.9
0.07
0.98
24.8
7.7
0.21
99.2
1500 (50O)
(500)
(300)
(100)
(400)
Sr
Ba
K
Rb
Th
U
100 (100)
<
<
<
<
100 (100)
100
100
700
200
(b) Compositions of coexisting garnets and Ca-perovskites under subsolidus conditions
R u n No.:
Conditions:
260
24.5 GPa, 1400 ° C
262
24.0 GPa, 1900 o C
228
25.0 GPa, 1400 o C
253
24.0 GPa, 1800 ° C
Phase:
Gt
CaPv
Gt
Gt
CaPv
Gt
CaPv
SiO 2
TiO 2
A1203
Cr203
FeO
MgO
CaO
Na20
50.6
0.26
19.0
0.14
2.4
13.2
11.7
1.5
48.1
1.5
8.2
0.04
2.0
4.3
31.9
1.0
49.5
3.3
7.8
0.06
1.7
2.7
34.5
0.80
48.6
0.27
18.2
0.13
6.6
12.8
11.4
1.8
47.8
3.5
5.2
0.03
1.5
3.0
33.2
1.6
Sum
98.8
97.0
100.2
98.4
99.8
95.6
2000
(200)
130
(50)
130
(50)
630
(lOO)
720
(lOO)
650
(100)
1000
(3O0)
6000
(1000)
7000
(1000)
3800
(8oo)
5700
(lOO3)
5300
(400)
350
(70)
500
(100)
350
(20)
900
(200)
450
(20)
3200
(4OO)
7100
(300)
2000
(100)
6100
(400)
3400
(300)
Sc
La
Nd
Sm
Ho
Yb
49.9
0.31
18.8
0.08
3.2
14.2
13.2
1.9
Sc
La
Nd
Sm
Ho
Yb
CaPv
44.6
4.0
8.2
0.09
1.5
2.5
28,1
1.4
101.6
90,4
1800
(IO0)
200
(100)
250
(100)
420
(100)
1300
(lOO)
1500
(100)
1300
(200)
19000
(2000)
20000
(2000)
20000
(2000)
13000
(100o)
6200
(400)
48,4
0,41
18,8
0.13
6.8
13.1
10.6
1.9
Sr
Y
Hf
Zr
Nb
K
< 100
Th
< 1000
U
< 150
Ba
Cs
100
(100)
< 100
Rb
< 140
Sr
200
(200)
500
(IO0)
14000
(2000)
10000
(1000)
8000
(300)
1000
(500)
300
(200)
8500
(1000)
136
TABLE 5 (continued)
(c) Compositions of garnets and Ca-perovskites coexisting with liquids
Run No.:
Conditions:
279
22.0 GPa, > 2100 ° C
282
22.0 GPa, > 2100 ° C
266
24.0 GPa, > 2100 ° C
Phase:
Gt
CaPv
Gt
CaPv
Gt
CaPv
L
SiO2
TiO2
A1203
Cr203
FeO
MgO
CaO
Na20
45.0
0.24
19.7
0.10
3.9
14.9
13.2
1.4
42.0
5.0
5.0
0.06
5.1
5.1
26.8
1.7
45.1
0.33
18.7
0.06
2.6
16,7
13,0
1.3
42.5
6.0
6.0
0.08
1.0
3.5
28.5
2.0
45.2
0.22
20.9
0.06
0.81
16.6
14.2
1.5
38.2
2.5
10.2
0.03
0.40
1.0
36.8
0.77
50.6 (1.0)
0.91 (0.05)
16.7 (0.5)
0.07 (0.02)
1.1 (0.1)
12.4 (0.5)
15.1 (1.0)
1.8 (0.2)
Sum
98.4
90.8
97.8
89.6
99.5
89.9
98.7
Nb
< 100
500
(100)
20 000
(2000)
22 000
(2000)
23 000
(2000)
9000
(1000)
3000
(500)
< 100
< 100
< 100
1600
(100)
< 100
K
Sr
8000
(1000)
20000
(3000)
20 000
(2000)
13 000
(2000)
6800
(500)
20000
(5000)
Sr
< 100
Ba
< 100
Rb
< 100
9500
(500)
300
(150)
< 100
Cs
< 100
Th
< 700
U
< 200
300
(100)
4000
(500)
3000
(400)
750
(100)
750
(100)
2500
(500)
1000
(300)
Zr
Hf
Y
Pb
100
(100)
250
(100)
300
< 200
Sc
La
Nd
Sm
Ho
Yb
400
(50)
500
(80)
1200
(100)
1600
(100)
garnet and M g - p e r o v s k i t e are r e m a r k a b l y sensitive
to temperature. A t 1 4 0 0 ° C , Dmpv/gnt values for
Z r a n d H f are 5 a n d 6 respectively. These increase
to 15 a n d 18 at 1 9 0 0 ° C (Table 7). Even m o r e
d r a m a t i c b e h a v i o u r is shown b y g a r n e t / C a perovskite (cpv). A t 1400 ° C, Dcpv/gnt values for
Z r a n d H f are 4 a n d 2.5 respectively. These increase to 200 a n d 50 at 2 1 0 0 ° C (Table 5c)! These
r e m a r k a b l e variations m a y be c o n n e c t e d with a
strong t e n d e n c y for Z r a n d H f to favour substitution as inverse perovskite c o m p o n e n t s at elevated
temperatures.
I n c o n t r a s t to this behaviour, the p a r t i t i o n coefficients for Sc, Y, Y b a n d Sm b e t w e e n m a j o r i t e
garnet a n d M g - p e r o v s k i t e show little v a r i a t i o n
between 1 4 0 0 ° C a n d 1 9 0 0 ° C (Table 4b). Likewise, p a r t i t i o n coefficients of Sr, Yb, Ho, Sm, N d
a n d L a between m a j o r i t e garnet a n d C a - p e r o v s k i t e
show only small changes between 1 4 0 0 ° C a n d
1900 ° C ( T a b l e 5b).
5.2. Garnet-liquid partition relationships
D a t a on the d i s t r i b u t i o n of m a j o r a n d m i n o r
150
(100)
50000
(5000)
25 000
(3000)
elements b e t w e e n g a r n e t a n d liquid were o b t a i n e d
in twelve u l t r a b a s i c c o m p o s i t i o n s , K A , KB, P A
a n d C A b e t w e e n 15 a n d 18 G P a a n d b e t w e e n
2000 a n d 2280 o C. E x a m i n a t i o n s of the p r o p o r tions of garnet, olivine a n d liquid in the charge
show t h a t the runs c o r r e s p o n d to varying degrees
of p a r t i a l m e l t i n g b e t w e e n the solidus a n d liquidus. N e a r - l i q u i d u s runs were p r e f e r r e d for the
accurate d e t e r m i n a t i o n of p a r t i t i o n coefficients
since these c o n t a i n e d relatively large garnet
crystals a n d i n t e r v e n i n g regions of q u e n c h e d liquid
which c o u l d be r e a d i l y a n a l y z e d w i t h o u t b e a m
overlap.
A n a l y t i c a l results for the d i s t r i b u t i o n s of m a j o r
and minor elements between garnet and ultrabasic
liquids are given in T a b l e 3. D a t a for the d i s t r i b u tions of A1203, T i O 2 a n d C a O b e t w e e n garnet a n d
liquid are p l o t t e d in Fig. 3 which also d i s p l a y s
c o r r e s p o n d i n g results f r o m o t h e r l a b o r a t o r i e s
[20,21,24,34]. T h e r e is little v a r i a t i o n in g a r n e t /
liquid p a r t i t i o n coefficients for C a O a n d T i O 2
t h r o u g h o u t the wide range of u l t r a b a s i c c o m p o s i tions. D~t/u q for A120 3 seems to show a small
137
0.50
20
(b)
[]
(a)
TI02
Al2Oa
ip
c
~ 0 .fi
_
0.25
15
c
,,2
I,-
m
c
~ 10
0
/
0.25
T I 0 2 In l i q u i d , wt%
0.50
10
/
i
(c) C a O
0.1
I
5
A1203 In l i q u i d , wt%
I
10
o
=
m 5
c
0
0
Komatllte
0
Pyrollte
}
This study
Q
Chondrlte
[]
Chondrlte
[20,34]
[]
Perldotlte
[21,24]
[]
I
5
C l I O In l i q u i d , w1%
I
10
Fig. 3. Concentrationsof (a) A1203, (b) TiO2, and (c) CaO in majorite garnet and coexistingultrabasic liquids in komatiite, pyrolite
and chondrite compositions.Circles denote data from this work. Squares denote data from other laboratories [20,21,24,34].
dependence on the bulk composition in the sense
that the D values increase with MgO content. Our
data yield a well-determined value D~t/li q of 2.5 +
0.5 for A1203. Only two of our 12 measurements
fall significantly outside this range. It is possible
that the discrepancies were caused by contamination of the quench liquid by alumina used in
polishing.
The data in Table 3 do not disclose any systematic dependence of the garnet/liquid partition
coefficients for trace elements upon bulk composition among the ultrabasic compositions investigated. A preferred set of representative g a r n e t /
liquid coefficients for major and minor elements
in ultrabasic compositions has been assembled
from the data of Table 3 and is given in Table 6.
It is possible only to provide upper limits for
partition coefficients of incompatible elements
where these are much smaller than unity because
of uncertainties introduced by counting statistics.
We find values of Dg~t/,q for K, Sr and Ba to be
smaller than 0.1. D values for La, Cs, Rb, U, Th
and Nb are less well constrained but are probably
smaller than 0.2. Based upon their observed in-
compatible behaviour in magmas, the D values of
this second group would generally be expected to
be smaller than those of K, Sr and Ba, i.e. they are
probably smaller than 0.1.
Preferred partition coefficients for garnet-liquid in the basaltic system are also given in Table
6. These were estimated from data obtained in
three runs at 20 G P a and 2200°C. The liquid
composition is relatively silicious (51-54% SiO2)
and is also calcium-rich (15-17% CaO). G a r n e t /
liquid partition coefficients for Sc, Yb, Y, Hf and
Zr are significantly smaller than in the ultrabasic
compositions, probably due to compositional differences between the liquids. In both systems, Sc,
Y, Yb, H f and Zr behave essentially as compatible
elements, whereas U, Th, La, Nd, Sm, Ba, Sr, K,
Rb and Cs are incompatible with Dgnt/liq < 0.]. A
significant feature in both systems is the strong
exclusion of intermediate REE (Sm) from garnet.
5.3. Mg-perovskite-liquid partition relationships
The bulk compositions of the majorite garnet
phase in equilibrium with Mg-perovskite (Table 7)
are close to those of garnets on the liquidus in
138
TABLE 6
Partition coefficients, Dgnt/liq between garnets and liquids in
ultrabasic (K, P, C) and basic (BA) compositions
Dgnt/liq
ultrabasic a
basic
Major elements (wt.% ratios)
TiO 2
AI203
FeO
CaO
Na20
0.40 (0.10)
2.5 (0.5)
0.6 (0.2)
0.60 (0.10)
0.1 (0.1)
0.1
2.2
0.3
0.4
0.1
1.7
1.4
1.3
0.8
0.6
0.2
1.4
0.7
0.7
0.2
0.15
Trace elements (ppm ratios)
Sc
Yb
Y
Hf
Zr
Sm
K, Sr, Ba, (Cs, Rb, La,
Th, U, Nb)
< 0.1
< 0.1
a Partition coefficients Dgnt/Liq in ultrabasic compositions are
estimated from the results of runs 182, 264-1, 264-2, 257, 5,
9, 29-2, 31-2 and 31-1.
u l t r a b a s i c c o m p o s i t i o n s . A c c o r d i n g l y , we h a v e
e s t i m a t e d the c o m p o s i t i o n of t h e p e r o v s k i t e p h a s e
w h i c h w o u l d crystallize o n the l i q u i d u s o f u l t r a -
b a s i c l i q u i d s at p r e s s u r e s e x c e e d i n g 25 G P a f r o m
the r e l a t i o n s h i p :
Ompv/li q = Ompv/gnt X Ognt/liq
T h e v a l u e s o f Ognt/liq a r e t a k e n f r o m T a b l e 6 a n d
w e r e o b t a i n e d at a t e m p e r a t u r e of a b o u t 2100 o C.
T h e s e h a v e b e e n c o m b i n e d w i t h t w o sets o f
Dmpv/gnt v a l u e s o b t a i n e d at 1400 ° C a n d 1900 ° C
( T a b l e 4) a n d t h e results are s h o w n in T a b l e 7.
T h e d a t a set o b t a i n e d at 1900 ° C is p r e f e r r e d for
c a l c u l a t i o n s o f Dmpv/liq v a l u e s for use in m a n t l e
d i f f e r e n t i a t i o n c a l c u l a t i o n s b e c a u s e o f closer p r o x i m i t y to t h e t e m p e r a t u r e at w h i c h Dgnt/, q v a l u e s
w e r e o b t a i n e d . T h e d a t a set b a s e d o n Dmpv/li q
(1400 o C) is useful b e c a u s e it p r o v i d e s a g u i d e to
the s e n s i t i v i t y of c a l c u l a t e d Dmpv/,q v a l u e s to the
t e m p e r a t u r e at w h i c h Dmpv/gnt v a l u e s w e r e det e r m i n e d . T h e d a t a i n d i c a t e t h a t Dmpv/gnt v a l u e s
at 2100 ° C a r e likely to b e s i m i l a r to t h o s e m e a s u r e d at 1 9 0 0 ° C f o r Sc, Y b , Sm, b u t m a y b e
s i g n i f i c a n t l y h i g h e r for Z r a n d Hr.
I t o a n d T a k a h a s h i [24] r e c e n t l y d e t e r m i n e d the
m a j o r e l e m e n t c o m p o s i t i o n of M g - p e r o v s k i t e o n
t h e l i q u i d u s o f a p e r i d o t i t i c c o m p o s i t i o n in a
single e x p e r i m e n t at 25 G P a a n d 2500 ° C. T h e i r
results are c o n s i s t e n t w i t h o u r s w i t h i n c o m b i n e d
e x p e r i m e n t a l errors, e x c e p t for A 1 2 0 3 p a r t i t i o n
c o e f f i c i e n t s . W e o b t a i n Dmpv/li q f o r A1203 of 0.5
as c o m p a r e d to 1.0 o b t a i n e d b y I t o a n d T a k a h a s h i .
T h e l a t t e r result s h o u l d b e c o r r e c t e d to a l l o w for
TABLE 7
Partition coefficients between Mg-perovskite and garnet at 1400 and 1900°C and estimated partition coefficients between
Mg-perovskite and ultrabasic liquid
TiO 2
A1203
CaO
Na20
Sc
La
Sm
Yb
Y
Hf
Zr
Nb
Dmpv/gnt
1400 o C
from Table 4
1900 o C
from Table 4
Dgm/liq
2100 o C
from Table 6
Dmpv/liq
estimated a
( - 2000 o C, 25 GPa)
10
0.3
0.3
0.2
4
(1.5) b
1.5
1.3
(0.5) b
6
5
10
7
0.2
0.4
0.2
3
_
1.4
1.7
2.2
18
15
-
0.40
2.5
0.60
0.1
1.7
< 0.1
0.2
1.4
1.3
0.8
0.6
< 0.1
3
0.5
0.2
0.02
5
< 0.1
0.3
2
3
14
9
- 1
a Estimated by Dmpv/qiq = Dmpv/gnt (at 1900 o C) × Dgnt/liq.
b Large uncertainties involved due to low concentration.
139
the effects of contamination by chromium (from
the heaters) encountered in their experiments. A
corrected partition coefficient based on total (A1
+ Cr) 3+ yields a slightly lower Dmpv/Uq value of
0.8.
It is possible that the remaining discrepancy is
caused by an experimental artefact. The charge
characterized by Ito and Takahashi had been completely molten on one side, but was crystalline on
the other. This latter region consisted of Mg-perovskite plus a few percent of irregularly dispersed,
small Ca-perovskites. The latter may have been
stabilized by La203 contamination also derived
from the heater. This texture strongly suggests the
previous existence of a marked temperature gradient across the sample. Moreover, the proportions,
distributions and sizes of Ca-perovskite inclusions
within the major mass of Mg-perovskite suggest
that this region of the charge had experienced a
modest degree of partial melting sufficient to
eliminate magnesiowiistite, and that the liquid
phase had migrated to the opposite side of the
capsule where the charge was completely melted
because of this higher temperature. Thus the composition of Mg-perovskite as measured would correspond to that of a near-solidus phase and would
not represent that of a phase which had equilibrated with the bulk liquid. This situation would
have caused values for equilibrium partition coefficients (Dmpv/liq) which are greater than unity to
be underestimated and values smaller than unity
to be overestimated. Accordingly, it seems possible that Ito and Takahashi have overestimated
Drnpv/fiq for A1203 and under-estimated the corresponding value for TiO 2. This situation highlights
the importance of minimizing temperature gradients across experimental charges and of selecting
them for analysis on the basis of textural evidence
which indicates relatively uniform temperature
gradients.
5.4. Ca-perovskite-liquid partition relationships
The most useful data set for estimating Ca-perovskite/liquid partition behaviour was obtained
in Run 266 on the MORB composition at 24 GPa
and above 1900°C (Table 8). This sample has
experienced only a small degree of melting and
contains residual Ca-perovskite, garnet and
stishovite. The perovskite phase demonstrates
spectacular enrichments in U and Th, and a mod-
TABLE 8
P a r t i t i o n c o e f f i c i e n t s (Dcpv/aiq) b e t w e e n C a - p e r o v s k i t e s a n d
l i q u i d s in b a s a l t i c c o m p o s i t i o n (BA)
Directly measured
in r u n 266
Th
U
Pb
Hf
Zr
TiO 2
Sr
La
Nd
Sm
Ho
Yb
Y
Nb
Sc
Na20
K, Rb,
Cs, B a
Calculated from mass balance
in r u n s 279, 282, a n d 266 a
observed
preferred
Depv./li q
Dcpv/liq b
20
25
2.7
2.4
(5.0) c
0.4
<0.1
20
20
10
6
5
4
3
3.8
4.6
5.5
2.7
1.2
2.5
1.4
0.2
0.4
<0.1
range
7 -100
6 - 70
3 - 14
2 - 11
2 9
1 . 5 - 25
1 . 8 - 7.7
1 . 4 - 6.0
1 . 6 - 6.9
1 . 9 - 8.2
1 . 3 - 3.5
0.81.4
1 . 1 - 3.8
0.12.3
0 . 1 - 0.2
0 . 2 - 0.6
<0.2
a Calculated from mass balance assuming 80±10% garnet,
< 5% C a - p e r o v s k i t e , a n d 15 + 10% liquid.
b O b t a i n e d a t 80% g a r n e t , 2.5% C a - p e r o v s k i t e , a n d 17.5%
liquid.
c O b s e r v e d b y I t o a n d T a k a h a s h i [24] in lherzolite system.
est enrichment in Sr. Barium, Cs, Rb and K
behave essentially as incompatible elements and
are strongly partitioned into the liquid. The texture indicates the presence of a temperature gradient across this charge (Fig. 2b). Accordingly, as
discussed earlier, the Ocpv/li q values for U, Th and
Sr (Table 8) may be underestimated whilst Ba, Cs,
Rb and K may be even more incompatible than
indicated in Tables 5c and 8.
Partial melting in the basaltic composition was
also observed in runs 279 and 281 at 22-24 GPa
and - 2200 ° C. As noted earlier, it was not possible to obtain liquid compositions directly by electronprobe microanalysis in these runs. Nevertheless, approximate limits on the crystal/liquid partition coefficients for many important elements
can be obtained by mass balance calculations,
based on the initial bulk composition and the
observed compositions of garnet and Ca-per-
140
ovskite, treating the composition of the liquid as
an unknown:
Cbulk = (egn, X Cgnt ) q- (ecpv X Ccpv)
+ ( P|iq X eli q )
where C denotes the concentration of trace elements in a given phase (or in the bulk composition) and P denotes the mass proportion of the
phases in question. The proportion by mass of the
major phase, garnet, can be estimated from observed textures and from constraints imposed by
mass balances for the major elements, Si, Mg and
Ti. The proportion of garnet egnt is thus found to
be 80 + 10%. The very high concentration of some
minor elements (e.g. U, Th, La) in Ca-perovskite
requires that the proportion of this phase be
smaller than 5%. Accordingly, the degree of melting experienced by this sample is estimated as
5-30%. These limits have been used to calculate a
permitted range of partition coefficients for many
elements between Ca-perovskite and liquid. This
range, together with preferred values is given in
Table 8. Where comparisons are possible, the preferred values so obtained for an assemblage of
80% garnet, 2.5% Ca-perovskite and 17.5% liquid
are consistent with those measured directly in run
266 and with a single direct measurement of the
Ocpv/liq partition coefficient for La determined by
Ito and Takahashi [24] on a peridotitic composition.
The remarkably high partition coefficients of
Th, U and Pb are noteworthy. Many other normally incompatible elements have partition coefficients between 1 and 6, whereas K, Rb, Cs and Ba
continue to behave as incompatible elements with
D values less than 0.1.
6. Implications for gross mantle differentiation
If the present peridotitic upper mantle is a
solidified partial melt originating from a more
primitive chondritic parental composition, as
argued by Herzberg and O'Hara [19], Takahashi
[21] Ohtani et al. [20] and Herzberg [36] there
must be a complementary silica-enriched reservoir
at greater depths consisting of residual phases
a n d / o r early cumulates. Two scenarios can be
envisaged: (1) majorite garnet fractionation (10-25
GPa), and (2) Mg-perovskite fractionation ( > 25
GPa). The mineralogy of the phase involved in
crystallization-differentiation depends on the
pressure regime in which this process occurs.
Voluminous ultrabasic liquids of global extent
could have been produced in a number of different ways. The two limiting cases are: (1) continuous extraction of a partial melt from depths near
or below the base of the present upper mantle,
and (2) fractional crystallization of a magma ocean extending well below the present upper mantle,
and possibly to the core-mantle boundary. In either
scenario, the chemical composition of the upper
mantle should reflect equilibrium between ultrabasic liquid and majorite garnet or perovskite.
The chondrite model mantle composition has
majorite garnet as its liquidus phase at pressures
equivalent approximately to mantle depths of between 450 and 700 km (e.g. [20,21]). Below about
750 kin, Mg-perovskite would become the liquidus
phase [24,34]. We will now consider the chemical
evolution of this model mantle composition which
might be caused either by majorite fractionation
or by perovskite fractionation, with the aim of
evaluating the hypothesis outlined in the Introduction whereby global differentiation of a chondritic
mantle gives rise to a pyrolite upper mantle. In
other words, can subtraction of majorite or Mgperovskite form a chondritic mantle yield pyrolite?
6.1. Majorite garnet fractionation
Our 'experimental data' show that the partition
coefficients of elements studied are not markedly
dependent upon bulk composition for the range of
ultrabasic compositions studied. We have accordingly used the preferred partition coefficient data
in Table 6 to calculate the composition of majorite
on the liquidus of a chondritic bulk composition
(Table 9).
The progressive compositional changes in a
chondritic mantle composition caused by separation of liquidus majorite have been calculated
on the basis of a fractional crystallization model
using our measured partition coefficients.
It should be noted that the two principal potential sources of experimental errors: (a) presence of
a temperature gradient across the charge, and (b)
departures from local equilibrium, would produce
systematic errors in D values in the same direc-
141
TABLE 9
Chemical compositions (wt.%) of chondritic model mantle and
fractionating garnets and perovskites
Chondritic
mantle (1)
SiO 2
AI203
Cr203
FeO
MgO
CaO
Na20
50.5
0.18
3.7
0.81
6.8
35.3
3.0
0.39
Sum
99.9
ZiO 2
Garnet
(2)
MgPv
(3)
52.5
0.09
8.7
1.1
3.3
32.4
1.7
0.18
57.7
0.54
1.5
0.19
2.0
37.4
0.60
0.04
100.0
100.0
CaPv
(4)
50.3
0.70
2.2
0.28
1.3
3.4
41.7
0.04
100.0
(1) "Chondritic" model mantle derived from C1 chondrite
composition [3]; (2) fractionating majorite garnet for chondritic
mantle composition; (3) fractionating Mg-perovskite for
chondritic mantle composition; (4) fractionating Ca-perovskite
for chondritic mantle composition. (2), (3) and (4) are calculated from experimentally observed partition relationships as
described in text.
tion, i.e. D values smaller than unity would be
overestimated and D values greater than unity
would be underestimated. Thus the calculated
fractionation trends and conclusions which follow
and which are shown in Figs. 4, 5 and 6 are
intrinsically conservative. According to these
calculations, fractionation of up to 30% of majorite
from a chondritic composition causes a drastic
decrease of A1203 and modest increases in CaO
and TiO 2. However, it has a very small effect
upon the S i / M g ratio of the residual liquid. This
point is brought out dramatically in Fig. 4a in
which Ca/A1, and A1/Ti and S i / M g ratios are
plotted as a function of majorite fractionation. It
would be necessary to remove much more than
50% of majorite from a chondritic composition in
order to approach the pyrolyte S i / M g ratio. However, the removal of more than 10% of majorite
would cause the Ca/A1 and A1/Ti ratios of the
residual liquid to deviate significantly from the
value of those ratios which are inferred to be
present in the upper mantle. Trace element concentrations provide similar constraints on the likelihood and extent of majorite fractionation. The
evolution of Sm, Yb, and Sc concentrations from
chondritic abundances are calculated using partition coefficients between majorite and ultrabasic
liquid (Table 6) and the results are shown in Fig.
4b. It is evident that these parameters are likewise
sensitive indicators of majorite fractionation.
As noted in the Introduction, the present upper
mantle source region of MORBs displays approximately chondritic relative abundances of
most refractory lithophile elements. We have
shown that the extensive majorite fractionation
necessary to derive the present upper mantle
S i / M g ratio from a chondritic parental composition would necessarily be accompanied by gross
fractionations of A1/Ca, A1/Ti, S c / S m and
S m / Y b ratios. Since these fractionations are not
observed, we reject the hypothesis of gross mantle
differentiation proposed in references [19-21,36].
More generally, the similarity between these ratios
1.4
(a)
o
'
/
1,2 ~
'
Ca/AI
1.0
I
I
Si/Mg
o
,~.=
0.8
E
I
0.6
=~,=
10
20
Garnet fractlonation, %
1.4
(b)
o 1
'
.
2
P
N 1.0
I
I
5 0.a
0.6
I
0
10
20
G a r n e t fractlonatlon, %
Fig. 4. Effect of majorite garnet fractionation on a chondritic
model mantle composition, calculated for partition coefficients
given in Table 7. Elemental ratios for the mantle composition
have been normalized to corresponding ratios in C1 chondrites.
(a) Variations in Si/Mg, Ca/A1, and A1/Ti ratios as a function of majorite garnet fractionation. The S i / M g ratio for the
present upper mantle is also indicated. (b) Variations in S m / Y b
and S c / S m ratios as a function of majorite garnet fractionation.
142
in the present upper mantle MORB source region
and chondritic ratios implies that majorite fractionation has had at most only a minor influence
upon the composition of this region.
6.2. Mg-perovskite fractionation
The arguments here are closely analogous to
those used in the previous section. Ito and Takahashi [24] showed that Mg-perovskite would be
the liquidus phase on peridotitic and probably on
chondritic compositions at pressures above about
26 GPa. Kumazawa [22] and Ohtani [23] proposed
that during formation of the Earth, the mantle
experienced complete or extensive partial melting
to depths greater than 700 km accompanied by
gross crystal-liquid fractionation involving the
separation of Mg-perovskite. This, in turn led to
the formation of a lower mantle dominantly composed of Mg-perovskite and an upper mantle of
pyrolite composition.
This model is tested in Fig. 5. The major element composition of Mg-perovskite crystallising
on the liquidus of a chondritic composition is
estimated analogously to the previous case for
garnet and is given in Table 9. It is seen that it
would require the separation of more than 50% of
Mg-perovskite to change the S i / M g ratio of the
mantle from the chondritic value of 0.95 to that of
upper mantle pyrolite (0.78). The variations of
S c / S m and S m / H f ratios as a function of Mgperovskite fractionation are also shown in Fig. 5.
It is apparent that these ratios are markedly influenced by quite small degrees of Mg-perovskite
fractionation. The observed near-chondritic
S c / S m and S m / H f ratios in the MORB source
region in the upper mantle do not permit prior
fractionation of more than a few percent of Mgperovskite.
Following the publication of our preliminary
results on perovskite fractionation [25,26], Ito and
Takahashi [24] also attempted to place constraints
on the role of perovskite fractionation, based upon
their measurements of partition coefficients
Dmpv/liq for CaO (0.05-0.1) and of A1203 ( - 1.0)
in a run at 25 GPa and 2500 o C. However, their
arguments are of doubtful validity. We have already described evidence suggesting that these
authors overestimated Ornpv/liq for A1203. Moreover, Ringwood [9] pointed out that the partial
molar volume of A1203 in solid solution in Mgperovskite is similar to that of corundum and that
at pressures substantially higher than 25 GPa, it is
likely to be exsolved from Mg perovskite to form
an ultra dense accessory phase, ~MgAI204 [35], in
which the partial molar volume of A1203 is substantially smaller than that of corundum:
MgSiO.xA1203 (perovskite) + MgO
MgSiO3-perovskite + xcMgA1204
1.4
I
o
1.2
+ (1 - x ) M g O
i
I
.
I
j
/
m/HI
Si/Mg
5
N
~
0.8
0.6
0
I
5
E
i
10
Mg-perovsklte fractionatlon,
Accordingly, it is expected that throughout most
of the lower mantle, at pressures above 30 GPa,
Mg-perovskite on the liquidus of mantle compositions would contain only small amounts both of
CaO and A1203. This casts additional doubt on
the significance of Ito and Takahashi's calculations.
%
Fig. 5. Variation of S m / H f , Sc/Sm, and S i / M g ratios as a
function of Mg-perovskite fractionation from a chondritic
model mantle composition, calculated from partition coefficients given in Table 7. The S i / M g ratio of the present upper
mantle is also indicated. Shaded areas indicate uncertainties
which would be caused by errors of + 2 in the S m / H f and
S c / S m ratios used in the calculation.
6.3. The Si / Mg ratio of the mantle
Models which assume that the bulk mantle has
a chondritic S i / M g ratio (0.95) and that the
silica-depleted, olivine-rich upper mantle ( S i / M g
= 0.78) is therefore underlain by a silica-enriched
lower mantle of pyroxenitic stoichiometry ( S i / M g
- 1 . 0 ) have hitherto proposed that this layered
structure was formed by crystal-liquid differentiation during early melting of the Earth (e.g.
[17,22,23]). Existing geophysical arguments which
143
have been used to support this interpretation are
demonstrably fragile, as pointed out by Ringwood,
[9]. Moreover, results described in the previous
section contradict this model and show that refractory lithophile element abundances in the upper mantle are inconsistent with the extensive
degree of crystal/liquid fractionation which is
required to produce the observed Si./Mg ratio in
the upper mantle from parental material possessing the chondritic S i / M g ratio. We conclude,
therefore, that no valid case now exists for maintaining that the major element bulk composition
of the lower mantle is substantially different from
that of the upper mantle. (This implies that the
depletion of silicon in the upper mantle (compared to chondrites) is shared by the entire mantle.)
7. Constraints on convective homogenization of an
early differentiated mantle
Advocates of an extensively molten primordial
Earth might nevertheless attempt to escape the
arguments presented above by proposing that subsequent subsolidus mantle convection effectively
rehomogenised the mantle after an early episode
of gross melting and differentiation, so that all
traces of this episode have been removed. We will
now evaluate this hypothesis.
Sun [10,11] has investigated the compositions
of high-Mg .basalts and komatiites (of the nonBarberon type) from Archaean terrains and has
concluded that the composition of the source regions of these voluminous magmas was generally
similar to that of m o d e m MORBs, and was characterized by approximately chondritic relative
abundances of lithophile involatile elements. This
situation has apparently prevailed throughout the
last 3.8 G a [10,11]. Hence, if the mantle indeed
melted and differentiated early in its history it
must have become effectively remixed by 3.8 Ga.
Recently, ancient zircon crystals with ages of
4.10-4.28 G a have been discovered in two localities in Western Australia at Mt. Narryer [12] and
at Jack Hills [13]. Kinny [14] has analyzed the
hafnium isotopic composition of 4.20-Ga zircons
from Mt. Narryer and has demonstrated that they
evolved in geochemical environments in which the
L u / H f ratio was approximately chondritic (within
30%). Our experiments (Table 7) show that crys-
1.4
1.2
Garnet fractlonatlon
o
4.2 Ga
zircon
1.0
1
o.s
E
~_ 0.6
t
Mg-perovsklte fracUonatlon
0.4
•
0.2
k
0
t
~,
10
~"
20
•
30
FracUonatlon, %
Fig. 6. Effect of majorite garnet and Mg-perovskitefractionations on Hf/Lu ratios in a chondritic model mantle composition, calculated from partition coefficients given in Table 7.
The shaded region indicates uncertainties which would be
caused by errors of +2 in the Hf/Lu ratio used in the
calculation. Also shown are the permitted bounds of Hf/Lu
ratios present in the geochemicalenvironment in which 4.2 Ga
zircon from Mt. Narryer, Western Australia, evolved between
4.55 and 4.2 Ga.
tallization of Mg-perovskite on the liquidus of
ultramafic melts would cause strong fractionation
of L u / H f ratios, making the reasonable assumption that Dmpv/li q for Lu is similar to that of Yb.
This effect is shown in Fig. 6, which has been
calculated in a similar manner to Figs. 4 and 5. It
follows from Fig. 6 and from Kinny's data that if
the primordial mantle experienced extensive crystallization-differentiation involving substantial
separation of Mg-perovskite, it must nevertheless
have become rehomogenised by convection by
about 4.20 G a ago. Fig. 6 shows that the H f / L u
ratio is less sensitive to majorite garnet fractionation and does not place such a strong constraint
on very early mantle differentiation involving this
latter phase.
The Earth is believed to have formed between
4.55 and 4.45 b.y. ago [42]. If it were completely or
extensively melted during formation to the extent
that substantial fractionation of Mg-perovskite occurred, the above results imply that it must have
become re-homogenised by 4.20 Ga. It is difficult
to conceive of a plausible differentiation-convec-
144
not experience an extensive degree of melting during the formation of the Earth and has retained its
original state of large-scale chemical homogeneity
since that time. It should nevertheless be understood that we are not excluding models according
to which a relatively small magma ocean (e.g. 200
km deep) may have been produced during formation of the Earth, as discussed by Ringwood [1].
Nor are we excluding second-order chemical inhomogeneities connected with the differentiation of
the mantle over geological time arising from the
operation of subduction processes as discussed by
Ringwood [39].
tion regime that could achieve this effect. Extensive early separation of Mg-perovskite from a
chondritic melt is expected to cause enrichment of
A1 and Ca in residual liquids (Table 7). The results
and discussion by Ringwood and Irifune [37] show
that this would have considerably expanded the
stability field of garnet in the upper mantle, leading to a stable density stratification of the upper
mantle in relation to the lower mantle which may
well have been immune to convective re-homogenization.
Even if the crystallization-differentiation process led initially to an unstable gravitational
stratification, as, for example might be caused by
strong enrichment of iron relative to magnesium
in overlying late differentiates [38], formidable
difficulties still confront convective re-homogenization on short timescales. For a wide range of
initial conditions, the denser high-level cumulates
would develop instabilities and simply sink
through the less-dense earlier cumulates to collect
near the base of the mantle. Continuation of this
process would lead, ultimately, to a gravitationally
stable density distribution throughout the mantle
which would inhibit convective rehomogenization.
The authors are indebted to Dr. S. Kesson for
critical reading of the manuscript and especially to
Mr. N. Ware for invaluable assistance in carrying
out the electronprobe analyses and to Mr. W.O.
Hibberson for converting some of our glass starting materials to amphibolites. We are also grateful
to Mr. P. Kinny and Mr. N. Ware for permission
to quote unpublished experimental data.
8. Conclusion
References
The results discussed in this paper show that
extensive melting and crystallization of the mantle
would lead to gross chemical differentiation and
fractionation of key elemental ratios away from
their chondritic values. The'observation that the
present upper mantle contains near-chondritic relative abundances of many involatile lithophile elements strongly suggests that the bulk compositions of the upper and lower mantles are similar.
Moreover, it implies either that the mantle has
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ago implies that re-homogenization must have been
completed prior to this time. It is difficult to
understand how convective re-homogenization of
a highly differentiated mantle could be achieved
in the relatively brief period between 4.45 and
4.20 Ga. It seems more likely that the mantle did
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