Download Ringwood Phase transformations and their bearing on the

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INAUGURAL INGERSON LECTURE
Phase transformations and their bearing on the constitution and dynamics of the mantle*
A. E. RINGWOOD
Research School of Earth Sciences, Australian National University, PO Box 4, Canbena, ACT 2605, Australia
Abstract-The
bulk chemical composition of the Upper Mantle (“pyrolite”) is derived from experimental
and petrological studies of the complementary relationships between basaltic magmas and refractory
peridotites. The phase transformations which are experienced by pyrolite between depths of 100-800 km
are reviewed in some detail, particularly with regard to their capacity to explain the seismic P and S
velocity profiles throughout this region. The transition of olivine and pyroxene to &(Mg,Fe)$iO., plus
garnet provides a satisfactory explanation of the velocity changes associated with the 400 km discontinuity
within the limits of error of the seismic velocity dete~inations.
Seismic velocities between 400 and 6.50
km are likewise consistent with this region crystallizing as an assemblage of &y(Mg,Fe)$304 plus garnet.
The depth of the 650 km seismic discontinuity corresponds closely to the pressure at which (Mg,Fe)$iOe
spine1 disproportionates to MgSi03 perovskite + (Mg,Fe)Q magnesiowtistite. This transformation is completed over a narrow depth interval (~4 km) and is capable of explaining the seismic characteristics of
the 650 km discontinuity. The elastic properties and density of the Lower Mantle are readily explained
within their observational uncertainties by a pyrolite composition crystallizing as an assemblage of perovskites plus magnesiow~stite. A substantial change in chemical composition (e.g., an increase in Si&
and/or a decrease in FeO) at the 650 km discontinuity is not required by available geophysical and
petrological data. The near-chondritic ratios of involatile lithophile elements in pyrolite provide important
boundary conditions for geochemical Earth models and place severe limitations upon hypotheses which
invoke large-scale melting early in the Earth’s history. They also imply that the Mg/Si ratio of the Lower
Mantle is similar to that of the Upper Mantle.
The geochemical evolution and dynamical behaviour of the mantle are strongly influenced by the
petrological differentia~on of pyrolite at mid-ocean spreading centres to form new oceanic lithosphere.
The MORB basaltic crust is underlain by a layer of harzburgite. During subduction, these lithologies
each respond to sequential phase transformations in a different manner, so that at any given depth, they
may differ in density from surrounding pyrolite. Most importantly, between 650 and 750 km, both former
basaltic crust and harzburgite are less dense (-0.15 and 0.05 g/cm3, respectively) than pyrolite. Relatively
young and thin subducted plates may have attained thermal equilibrium with surrounding mantle by
the time they reach the 650 km discontinuity. Because of their buoyancy below 650 km, these plates
would not be able to penetrate the 650 km discontinuity and instead are deflected laterally along the
discontinuity. This process would eventually produce a layer of former basaltic crust (gametite), buoyantly
trapped on top of the 650 km discontinuity, which would partially isolate the convective systems of the
Upper and Lower Mantle. In contrast, older and thicker oceanic plates may be sufficiently cool and
strong to permit their differentiated upper layers to penetrate the 650 km discontinuity. However, the
tips of these plates experience substantial buoyancy stresses at this depth which cause buckling. in consequence, the descending slab piles up and forms a large melange, or “megalith,” of mixed former hanburg&e and former oceanic crust with cross-sectional dimensions amounting to several hundred kilometres.
Its integrity is maintained for a Iimited period (- lo8 a) by high viscosity arising from its lower temperature
as compared to surrounding mantle. The presence of megaliths may explain a number of geophysical
observations including the complex structures present near the intersections of slabs with the 650 km
discontinuity, which have recently been imaged by seismic tomography, as well as associated depressions
in the depth of this discontinuity. The megahth functions as a “cold finger” in the Lower Mantle and
may initiate a sinking convection current. After the cessation of subduction, the megalith ~aduaily warms
up, accompanied by reduction in its viscosity, and ultimately becomes entrained in the convective system
of the Lower Mantle. Mantle convection is thus envisaged as a hybrid system, with a large degree of
independent convection within both the Upper and Lower Mantle, combined with a more limited exchange
of material between these regions. This behaviour is enhanced by the high viscosity of the Transition
Zone as compared to the Upper and Lower Mantle. The origins of intraplate (hot-spot) magmas are
considered in terms of the above model. Partial meiting of the former basaltic crust of the slab at depths
of 200-600 km refertihzes overlying depleted peridotite. This refertilized material, entrained by the subducting plate, accumulates in a thin zone immediately overlying the garnetite layer on top of the 650 km
* Delivered at the Goldschmidt Conference, Baltimore, MD, May 12, 1988.
2083
A. E. Ringwood
2084
discontinuity.
Subsequent reheating of this fertile peridotite causes diapirs to ascend from the 650 km
discontinuity
accompanied
by partial melting within the Upper Mantle to form geochemically enriched
magmas which are erupted at hot spots.
1. INTRODUCIION
THE DEVELOPMENT OF THE THEORY of plate tectonics has
focussed attention on the nature of the “engine” within the
mantle which drives plate motions. It is recognized that this
engine represents a complex form of convection
which is
poorly understood. Important boundary conditions on convective processes have been provided by geochemical studies
of basalt source regions, some of which have maintained their
identities for more than a billion years, despite the mixing
induced by mantle convection. It is essential to establish the
physical locations of these reservoirs and the processes by
which they have formed. The study of these topics has given
rise to an interdisciplinary
field appropriately named “chemical geodynamics”
(ALL~GRE, 1982).
Before meaningful progress can be made towards addressing the basic problems of chemical geodynamics, it is essential
to possess an adequate understanding
of the present physical
and chemical constitution
of the mantle. Major progress towards achieving this objective has been made during the last
few decades. To a large extent, this has arisen from controlled
laboratory experimentation
on the chemical and physical
properties of mantle minerals and rocks over a pressure-temperature regime encompassing
a large proportion
of the
mantle. These topics are reviewed in the first half of the present paper, with particular emphasis on the key role of phase
transformations.
An attempt is then made to apply knowledge
in this field to some topical areas of chemical geodynamics.
The mantle is divided into three regions on the basis of
seismic velocity distributions (Fig. 1). The Upper Mantle em-
braces the region between the Mohorovicic
Discontinuity
(marking the base of the crust) and a major seismic discontinuity occurring near a depth of 400 km. A second major
seismic discontinuity
occurs near a depth of 650-670 km
and the region between these discontinuities
is known as the
Transition Zone. The Lower Mantle comprises the large region extending between the “650 km discontinuity”
and the
Core, which is encountered near a depth of 2900 km. Many
earth scientists use the term “upper mantle” to describe the
entire region between the Moho and the 650 km discontinuity. This terminology
should be discouraged. The Transition Zone possesses an entirely distinct mineralogy to the
regions above and below, and this is manifested in important
differences in physical properties which are of considerable
geodynamic significance. It therefore warrants specific recognition as a major province of the mantle.
Although the existence and positions of major seismic discontinuities near 400 and 650 km are well established, there
are substantial uncertainties
in estimates by various authors
in the velocity changes at these discontinuities
and also in
the seismic velocity gradients between depths of -200 and
800 km, as illustrated in Fig. 1. BENNETT (199 1) has provided
an illuminating discussion of the underlying causes of these
very real uncertainties,
estimates of which are given in Table
1. Accordingly, it is not desirable to select any unique velocity
distribution as the basis for detailed interpretations
of mantle
mineralogy. These interpretations
should always include appropriate allowances for uncertainties in the seismic data. In
a recent comprehensive
seismic investigation,
SHEARER
(199 I) concluded
that the two major intramantle
discontinuities were located at depths of 410 f 3 km and 660
t- 8 km.
2. THE UPPER MANTLE
(a) Chemical Zoning in the Upper Mantle
41
200
300
400
500
600
700
800
DEPTH km
FIG. 1. Compressional and shear velocity profiles derived from
seismic observations: PEM (DZIEWONSKI et al., 1975); SHR14
(HELMBERGERand ENGEN, 1974); K8 (GIVEN and HELMBERGER,
1980); PREM (DZIEWONSKI and ANDERSON, 1981); GCA (WALEK,
1984); GH (GRAND and HELMBERGER, 1984).
The P-wave velocities of most regions of the Upper Mantle
immediately
underlying
the Mohorovicic
Discontinuity
(Moho) are in the range 8.1 -t 0.4 km/set. This property,
coupled with certain broad petrological and chemical limitations, effectively restricts the mineralogical composition of
this region to some combination
of olivine, pyroxene(s), and
garnet. The principal rock types containing these minerals
are peridotite (olivine-pyroxene)
and eclogite (garnet-pyroxene). Mineralogical
intermediaries
between these two rock
types are rare.
A wide range of evidence reviewed by RINGWOOD (1975)
shows that the uppermost mantle or lithosphere is dominantly
composed of peridotite, with eclogite widely distributed as
local segregations but relatively small in total amount. This
interpretation
would probably represent a consensus view
today; however, the consensus is a relatively recent phenomenon. During the 1950s and 1960s there was an active debate
as to whether the bulk composition
of the Upper Mantle was
dominantly
eclogitic or peridotitic. The ill-starred Mohole
project was conceived to resolve this debate and to determine
2085
Inaugural Ingerson Lecture
Table 1
ESTIMATED PERMISSIBLERANGES OF P-ANOS-WAVE
SEISMIC
400 AND 650 KM
AND
S-WAVE
Seismic
Velocity
Velocity
Velocity
Ah’D
BETWEEN
CORRESPONDING
RANGES
FOR
p-
400-650KM (KENNET 1991)
P-waves
S-waves
at 400 km
2.5-5.8
2.8-5.7
at 650 km
3.6-7.3
3.0-7.5
(%)
changes
discontinuity
GRADIENTS
parameter
changes
discontinuity
DISCONTINLRTIES
VELOCITY CHANGES ATTHE
(W)
gradients
between
400-
1.8-2.9 x 1O-3
2.0-5.0 x 10-3
650 km (km/sec)/km
whether or not the Mohorovicic Discontinuity represents a
phase change from gabbroic lower crust to eclogitic upper
mantle. The debate was effectively resolved via an extensive
experimental investigation of the gabbro-eclogite transformation (RINGWOODand GREEN, 1966), which showed that
the geophysical characteristics of the Moho in most regions
were inconsistent with the proposed phase change.
Numerous samples of peridotites from the Upper Mantle
have been transported to the surface as xenoliths in kimberlites and alkali basalts, or have been intruded into the crust
during erogenic activity. Most of these Upper Mantle peridotites are strongly depleted inlow melting-point components
and incompatible elements, so that they would be unable to
produce the common types of basaltic magmas if partially
melted. Nevertheless, we know that basaltic magmas have
been erupted in copious volumes throughout geological time
at localities scattered all over the Earth’s surface, in both
continental and oceanic settings. It therefore appears that
beneath the refractory peridotite layer, there must exist a
more primitive source region which has retained a significant
basaltic component. RINGWOOD( 1962a,b) and GREEN and
RINGWOOD(1963) denoted this primitive source material by
the term “pyrolite,” implying a non-specific olivine-pyroxene
rock capable of yielding basaltic magmas on partial melting
(some workers have preferred to use the term “primitive
mantle” in this context). Peridotite is believed to represent
the refractory residue remaining after basaltic magma has
been extracted from pyrolite. We thus arrive at a chemically
zoned model for the Upper Mantle as indicated in Fig. 2.
Beneath ocean basins the layer of depleted peridotite is
believed to be quite thin (e.g., lo-50 km). The underlying
pyrolite provides the source region for the most abundant
class of magmas erupted at the Earth’s surface: mid-oceanic
ridge basalts (MORBs). The xenolith population in kimberlites which penetrate stable continental cratons implies that
the depleted peridotite layer beneath cratons is much thicker,
probably in the vicinity of 150-200 km. However, where
continental plates are rifted and begin to move apart, MORB
basalts are erupted along the newly formed spreading centres
(e.g., the Red Sea). This implies that the pyrolite layer is of
global extent, continuous beneath ocean basins and continents, although deeper beneath the latter (Fig. 2).
(b) Basalt Petrogenesis and the Pyrolite Model
The pyrolite model is based upon the complementary relationships between basaltic and komatiitic magmas and their
respective peridotitic and dunitic refractory residues. GREEN
and RINGWOOD (1964, 1967) showed how the petrogenesis
of various classes of basaltic magmas could be interpreted in
terms of varying degrees of partial melting of pyrolite at defined depth intervals in the mantle. The methodology employed by these workers during the 1960s was based upon
high-pressure-high-temperature
experimentation combined
with analyses of crystalline and quenched liquid phases via
electron-probe techniques. It was this innovation which made
it possible to interpret the physico-chemical behaviour of
complex, multi-component systems as a function of P and
T; and the methodology has been widely adopted by other
DEPTH
Km
OCEAN
CONTINENT
GREGATIONS
FIG.2. Chemically zoned model for the Upper Mantle.
2086
A. E. Ringwood
laboratories during the 1970s and subsequently. A second
powerful tool for investigating the petrogenesis of basaltic
magmas was introduced by GAST (1968), who showed how
the partitioning of trace elements could be used to constrain
the nature ofthe partial melting and fractional crystallization
processes involved in magma genesis.
Methods for estimating the chemical composition of pyrolite were discussed extensively by RINGWOOD(1975). Petrogenetic relationships between a primitive MORB and residual harzburgite were employed by GREEN et al. ( 1979) to
formulate this parental mantle composition (Table 2). A corresponding estimate was made by SUN (1982) on the basis
of complementary relationships between komatiitic magmas
and residual dunites (Table 2). A second method is based
upon the recognition of types of Iherzolites which have experienced only very small degrees of partial melting and hence
are likely to approach the pyrolite composition (e.g., JAGOUTZ
et al., 1979; Table 1). These compositions are in close agreement with more recent estimates for the composition of the
primitive Upper Mantle obtained by FREY et al. (1985)
BONATTI et al. (1986) ZINDLER and HART (1986) and
MCDONOUGH ( 1987).
The methodology used to derive the major element composition of pyrolite has also been extended to obtain the
abundances of minor elements in the primitive Upper Mantle
(RINGWOOD, 1966a,b; RINGWOODand KESSON, 1977; SUN,
1982; ZINDLERand HART, 19.86; MCDONOUGH, 1987). The
reliability of these estimates has improved successively as the
data base has expanded. The estimate of the composition of
the primitive Upper Mantle from MCDONOUGH (1987) is
given in Table 3. Compositional models of this type have
proven to be of considerable utility in several areas of geo-
chemistry. For example, the abundance patterns of siderophile and volatile elements in the Earth’s mantle provide
important constraints on models of accretion of the Earth
and formation of the core (e.g., RINGWOOD, 1966a,b, 1984;
W.&NKE,1981; SUN, 1982; W~CNKE
and DREIBUS,1988). They
have also been used widely in modelling magma petrogenesis
by partial melting processes and crust/mantle differentiation
(see, e.g., HOFMANN, 1988).
The data in Table 3 show that many involatile lithophile
elements in pyrolite (e.g., Mg, Ca, Al, Ti, Y, SC, heavy and
intermediate REEs, Zr, and Hf) are present approximately
in chondritic relative abundances (see also NESBI~Tand SUN,
1976; SUN and NESBITT, 1977; ZINDLER and HART, 1986).
Moreover, if the entire mantle is assumed to be of pyrolite
composition, a simple relationship exists between the composition of the bulk Earth (mantle and core) and the composition of primitive Cl chondritic meteorites, which display
similar relative abundances of involatile elements to the Sun
(RINGWOOD, 1966a, 1975; ZINDLER and HART, 1986). Essentially, a geochemically self-consistent Earth model can be
derived from the CI chondrite composition by processes involving partial reduction of iron and nickel oxides to form
a metallic core together with loss of volatiles. It is important
to note, however, that about 20% of the total silicon must be
removed from the mantle composition in order to preserve
this relationship. Processes which could be responsible for
this silicon deficit are discussed in Section 5.
(c) Structure of the Upper Mantle
The seismic velocities of the uppermost 200 km of the
mantle vary in a complex manner both laterally and vertically.
Table 2
PYROLITEI MODEL COMPOSITIONS
1
(Jagoutz,
et al., 1979)
2
(Sun, 1982)
3
(Green,
et al., 1979)
SiO2
45.13
44.49
Ti02
0.22
0.22
0.17
Al203
3.96
4.30
4.4
Cr203
0.46
0.44
0.45
CaO
3.50
3.50
3.4
MgO
38.30
37.97
38.8
Fe0
7.82
8.36
7.6
NiO
0.27
0.25
0.26
MnO
0.13
0.14
0.11
Na20
0.33
0.39
0.4
lOOMg/(Mg+Fe)
1. Least
depleted
2. Komatiite
3. MORB
89.7
ultramafic
- dunite
- harzburgite
89.0
xenoliths
model
model
45.0
90.1
Inaugural lngerson
Lecture
2087
Table 3
ESTIMATEDCHEMICALCOMPOSITIONOFTHEPRIMITIVEUPPERMANTIE-TYROLITE"
(AFTERMCDONOUGH,~~~~)
Element
CI
I’yrolite
Li (ppm)
1.57
1.6
Be
0.027
0.08
B
0.98
0.5
250
C
50000
26
F
61
2545
Na
4950
22.45
Mg%
9.30
2.36
Al %
0.827
Si %
10.25
20.93
95
P (ppm) 1080
350
S
62000
30
Cl
700
240
K
545
Ca %
0.902
2.573
SC
5.98
17.34
Ti
440
1280
82
V
56
2935
Cr
2670
1080
Mn
1900
Fe %
18.10
6.53
105
Co
500
Ni
108ccl
1890
Gl
120
30
56
Zn
310
Ga
10.0
3.9
Ge
32.4
1.1
A5
1.93
0.13
Se
18.6
0.05
Br
3.57
0.075
Rb
2.32
0.635
Sr
7.26
21.05
Y
1.57
4.55
Zr
3.87
11.22
Nb
246
713
MO(ppb) 920
65
4.2
Ru
710
1
Rh
134
5
I’d
555
Pyrolite
(Normalized
to Mg and CI)
Element
CI
Pyrolite
8
40
0.017
0.023
80
13
0.067
1720
175
0.042
160
5
0.013
2260
13
0.002
430
11
0.011
0.073
1.20
1.20
1.20
1.20
1.20
1.20
1.20
1.20
1.20
1.20
1.20
1.20
1.20
1.20
1.20
1.20
1.21
0.09
0.003
0.003
0.003
0.003
0.002
0.010
0.021
0.031
0.009
1.20
1.07
0.42
1.23
Cd
0.21
In
0.02
Sn
0.18
Sb
0.21
Te
1.00
I
1.18
cs
188
33
0.85
Ba
2410
6989
As
200
710
0.036
La
244
0.0023
Ce
1833
0.018
Pr
632
95.7
0.18
Nd
471
1366
1.18
Sm
ELI
153
58
444
1.20
1.21
Gd
205
595
0.61
Tb
0.46
254
0.24
“Y
Ho
0.15
Er
165
0.087
Tm
0.072
Yb
0.10
0.075
0.16
0.014
37.2
56.3
708
278
168
108
737
163
479
25.5
74
481
LN
166
25.4
Hf
107
309
Ta
14
41
w
95
21
0.028
Re
36
0.0011
OS
520
3.4
0.009
Ir
480
3.3
0.11
rt
AU
1010
6.8
140
0.75
Hg
Tl
400
1.20
140
7
1.20
I%
2470
185
1.20
1.20
Pyrolite
(Normalized
to Mg and CI)
73.7
0.28
10
0.029
Bi
110
2.5
0.002
Th
29
84.1
0.003
U
8.1
21
0.004
From Li to Zr element concentrations
and Mg, Al, Si, Ca and Fe are in wt%.
are given
These variations are caused by several factors. Pyrolite is capable of crystallizing in four distinct mineral assemblages
which have well-defined fields in P, T, and fHzo space (GREEN
and RINGWOOD, 1970), and possess distinctly different physical properties. Additional complications
are introduced by
small degrees of partial melting and by anisotropy of olivine.
These topics have been reviewed by RINGWOOD (1975),
GREEN and LIEBERMANN(1976), and LEVEN et al. (198 1).
Below a depth of about 70 km, the stable mineral assemblage
displayed by pyrolite is olivine, garnet, clinopyroxene
rt orthopyroxene.
This assemblage
remains stable throughout
most of the Upper Mantle to a depth of 400 km (AKAOGI
and AKIMOTO, 1979). (However, orthopyroxene is eliminated
in ppm,
Nb to U are given
via solid solution in clinopyroxene
CALO and GASPARIK, 1990.)
in ppb
below about 300 km: PA-
Because of its relative depletion in Fe, Ca, and Al, the
refractory peridotite layer (Fig. 1) is about 0.06 g/cm3 less
dense than underlying garnet.pyrolite (RINGWOOD, 1966~).
In consequence, this layer is gravitationally stabilized and
has therefore been resistant to disruption by convection
(JORDAN, 1979, 198 1; O’HARA, 1975). The development of
the sub-continental
lithosphere as a chemical boundary
layer is probably an evolutionary feature (CLARK and RINGWOOD, 1964). This layer attains its greatest thickness, possibly in the vicinity of 200 km, beneath Precambrian shields
and has experienced
an extremely complex geological and
2088
A. E. Ringwood
geochemical history. Although depleted in its major element
chemistry, it appears to have been subjected to repeated episodes of “metasomatism”
in which local domains enriched
in incompatible
elements were formed.
Several seismic studies of the lithosphere beneath stable
continental regions have provided evidence of a discontinuity
at a depth of about 200 km. A high-resolution study by HALES
et al. ( 1980) indicated an increase in P-wave velocity of about
0.3-0.5 km/set at this depth, followed by a further small
decrease about 30 km deeper. Whereas the P- and S-wave
velocity distributions
in the sub-continental
lithosphere are
well explained at most depths by peridotite and garnet pyrolite
lithologies, this complex seismic feature at 200 km is not so
readily interpreted in these terms (LEVEN et al., 198 1). Moreover, it does not appear to be of global extent (SHEARER,
199 1). For example, it was not recognized in the data of the
Early Rise seismic profile which traversed North America
with a comprehensive
coverage of azimuth (WARREN, 1968).
LEVEN et al. (1981) suggested that the discontinuity
might
be caused by preferred orientation
of olivine and pyroxene
crystals near a depth of 200 km. These minerals are highly
anisotropic in their elastic properties. Preferred orientation
of olivines and pyroxene may have been caused by shearing
near the boundary of the sub-continental
lithosphere with
the underlying asthenosphere,
as continental plates migrated
across the Earth’s surface.
(d) Large-scale Melting of the Upper Mantle
RINGWOOD (1975, pp. 577-579) pointed out that the Upper Mantle would probably have been molten to a depth of
200-400 km immediately
after accretion of the Earth. The
fate of this primitive terrestrial magma ocean has since become the subject of considerable
speculative
discussion.
Ringwood concluded
that the magma ocean would have
crystallized rapidly, and that the differentiated
cumulates
would have been subducted into the mantle and homogenized
by convection. It now seems that the process may have been
more complicated. Recent evidence implies that the densities
of ultrabasic partial melts generated in the mantle at depths
of 250-400 km are higher than those of their olivine plus
pyroxene residue (STOLPER et al., 1981; OHTANI, 1984; RIGDEN et al., 1984; AGEE and WALKER, 1988a; MILLER et al.,
199 1). Thus, an ultrabasic magma ocean could have been
gravitationally
stable in the 250-400 depth interval and
bounded by an overlying layer of olivine and pyroxene crystals. NKBET and WALKER (1982) proposed that such a subterranean magma ocean existed for as long as two billion
years and was the source of much of the komatiitic volcanism
which occurred during the Archaean.
However, ARNDT
(1986) has drawn attention to several geochemical and isotopic difficulties attached to this hypothesis. In particular, it
is unable to provide an acceptable explanation
of the “depleted” geochemical signatures of many Archaean komatiites.
MCFARLANE and DRAKE (1990) also showed from studies
of the partitioning of Co and Ni that Upper Mantle rocks do
not display the signatures that would be expected on the basis
of the olivine-flotation
model of ACEE and WALKER (1988b).
Recent studies (e.g., HERZBERG and O’HARA, 1985; OHTANI, 1985; TAKAHASHI, 1986; TAKAHASHI and ITO, 1987)
have demonstrated
that the temperature interval between the
solidus and liquidus of peridot&e decreases quite markedly
with pressure, from 600°C at atmospheric
pressure to less
than 1OO’C at 16 GPa (see also, JACKSON, 1977). This has
led to the proposal that the Upper Mantle has itself been
formed by partial melting of materials from the Transition
Zone and Lower Mantle (e.g., HERZBERG and O’HARA,
1985). According to this model, the region of melt extraction
was located mainly in the Transition Zone between depths
of 400 and 670 km where majorite garnet is the liquidus
phase on a pyrolite composition
(TAKAHASHI, 1986). Presumably, the mechanism
would have involved subsolidus
convection
throughout
the Lower Mantle with upwelling
plumes
experiencing
partial melting
via adiabatic
decompression
as the plumes ascended into the Transition
Zone.
Although the existence of a narrow solidus-liquidus
temperature interval may be consistent with the above scenario,
it is by no means sufficient to validate it. The small melting
interval could equally reflect the simple facts that olivine
happens to possess a high melting point at zero pressure
combined with a low melting-point
gradient with pressure,
whilst pyroxenes and garnet have much lower melting points
at zero pressure but possess relatively high melting-point gradients. Given this conjunction
of properties, it is inevitable
that peridotite would possess a wide melting interval at low
pressures and a small melting interval at high pressure. These
characteristics
alone do not justify the petrogenetic interpretation of the Upper Mantle as a product of partial melting.
It was pointed out earlier that many involatile lithophile
elements (Mg, Ca, Al, Ti, Zr, Hf, SC, Y, and heavy and intermediate REEs) are present in pyrolite in near-chondritic
ratios. Experimental
results by KATO et al. (1988a; summarized in Table 4) show that majorite garnet on the liquidi
of chondritic and komatiitic melts at 16-20 GPa is enriched
by factors of 1.5-2.5 over coexisting liquid in Al and SC and
is correspondingly
depleted in Ca, Ti, Sm, and La by more
than a factor of 2. These observations
imply that a partial
melting process in which majorite was the principal residual
phase would have caused much larger fractionations
of these
elements than are observed in pyrolite. If there had been a
primordial magma ocean, all traces of its existence must have
been removed by subsequent subsolidus convection. Hafnium
isotopic compositions
of Archaean komatiites point to the
same conclusion (GRUAUet al., 1990).
3. THE TRANSITION
ZONE
(a) Historical Background
In a classic paper on the elasticity of the mantle, BIRCH
( 1952) concluded that the increases of seismic velocities between depths of about 300-900 km were too large to be caused
by self-compression
of homogeneous
material, and that
mantle silicate minerals therefore transformed
to denser,
closer-packed
phases in this depth interval. Testing Birch’s
hypothesis provided a challenge for experimentalists
since
the pressures and temperatures in this region were well beyond
the capacities of existing high P,T apparatus. However, the
plausibility of his interpretation
was strengthened
by observations that many germanates;isostructural
with silicates at
2089
Inaugural lngerson Lecture
Table 4
MgSiO3
PARTITION COEFFICIENTS FOR SELECTED ELEMENTS BETWEENLIQUIDUS
PEROVSKITE, PYROPERICH
GARNETS
AND
IJLTRABASIC MELTS, AND
Casio3 PEROVSKrrEAND A BASALTICMELT. @ROMUT0
MgSiOs(pv)’
Element
CaSi03
BETWEEN
et al., 1988a)
(pv)
D mpv/liq
Dcpv/liq
Dgnt/liq
Ca
0.2
0.8
0.6
Al
0.5-0.8
0.6
2.5
Na
0.02
0.4
0.1
Ti
3
3
0.4
Zr
9
5
0.6
Hf
14
6
0.8
U
25
Th
20
Nb
-1
1.4
<O.l
SC
5
0.2
1.7
Y
3
2.5
1.3
Yb
2
1.5
1.4
Sm
0.2
6
0.2
La
<O.l
5
<O.l
10
PO
Sr
3
Ba
<O.l
<O.l
K,Rb,Cs
<O.l
<O.l
*
This is the preferred
set of partition
coefficients
obtained
in the
investigations
of Kato et al. (1988a, 1988b). It is believed that they are likely to be
accurate to +50% and in many cases to +-30%. The set of partition
coefficients
obtained
by Kato et al.(1988b)
represents
limits only, owing to constraints
imposed by the experimental
method employed.
Nevertheless,
they demonstrate
that Dmpv/liq (Sc,Zr,Hf) > 3 and Dmpv/liq (Ca,Sm) < 0.2.
low pressures, transformed
to much denser polymorphs
at
modest pressures and that germanates tended to serve as highpressure, crystal-chemical
models for silicates (e.g., RINGWOOD, 1962c, 1966~). Moreover,
studies of phase relationships displayed by germanate-silicate
systems enabled quantitative predictions to be made of the pressures required to
transform
silicates into denser polymorphs.
Thus, from a
study of the thermodynamics
of Mg$i04-NizGe04
solid solutions (at ambient pressure), RINGWOOD (1956, 1958a) calculated that MgzSiOd should transform from the olivine to
the spine1 structure at 17.5 t 5.5 GPa, 15OO”C, and that the
spine1 polymorph
would be 11 + 3% denser than olivine.
(These predictions have subsequently been verified.) A similar
result was obtained by direct extrapolation
of high-pressure
phase boundaries in this system (RINGWOOD and SEABROOK,
1962) and is shown in Fig. 3. It was also found that the olivines
Fe2Si04, N&SiO+ and Co2Si04 transformed to denser spine1
+ The designation
+ FeO).
Mg’ refers to the molar
ratio
100 MgO/(MgO
structures at 2-7 GPa (RINGWOOD, 1958b, 1962d, 1963) and
that Si02 transformed
to the rutile structure near 10 GPa
(STISHOV and POPOVA, 1961). These results, reviewed by
RINGWOOD (1962c, 1966~) and CLARK and RINGWOOD
(1964), were widely considered to have established the validity
of Birch’s hypothesis.
Nevertheless, the challenge to establish the detailed mineralogical nature of the Transition Zone remained. In 1966,
Ringwood and Major developed an apparatus capable of
achieving - I8 GPa and 12OO”C, and employed it to discover
many new phase transformations
in mantle silicates, including those of magnesian olivines (Mg’80)t to spine1 and of
(Mg’& to a spinel-like (beta) phase (RINGWOOD and MAJOR,
1966). Transformations
of several silicate pyroxenes into a
new kind of garnet structure containing up to 25% of octahedrally coordinated silicon were also observed (RINGWOOD,
1967; RINCWOOD and MAJOR, 1971). The latter workers also
synthesized a perovskite containing -80% CaSi03, and obtained strong evidence that pure Casio3 perovskite had been
synthesized
in their experiments.
This synthesis was later
confirmed by LIU and RINGWOOD (1975). Reviews of this
2090
A. E. Ringwood
tories using multi-anvil systems. Preeminent in this field has
been the laboratory of Akimoto at the University of Tokyo.
Some of the more important experimental investigations
carried out by this group are described by AKIMOTO (1972,
1987), AKIMOTO and FUJISAWA(1966, 1968) AKIMOTO et
al. (1976, 1977), AKAOGI and AKIMOTO (1977, 1979), and
AKAOGI et al. ( 1987). More recently, the laboratory of Ito at
the University of Okayama has made a wide range of major
contributions (discussed later). During the last few years, there
has been considerable proliferation of multi-anvil and diamond anvil laboratories and the field is now extraordinarily
active.
,
20
NizGeOd
40
60
mole % forsterite
OLIVINE
solid
solutions
a0
Mg,SiO,
FIG. 3. The System Ni,GeO,-Mg2SiOdat 600°C and O-9 GPa.
Extrapolation of the phase boundaries indicates that pure Mg2Si04
olivine should transform to the spine1 structure near 16 GPa (after
RINGWOOD and SEABROOK, 1962).
work, mainly carried out at the Australian National University, were given by RINGWOOD( 1970, 1975).
The early 1970s witnessed the introduction of two new
techniques of generating high pressures and temperatures
which have since played key roles on this field. BASSETTet
al. ( 1967) developed an improved diamond anvil cell capable
of generating 30 GPa and, in 1972, combined it with laser
heating. Using this system, BASSETTand MING (1973) observed the disproportionation
of FezSi04 spine1 into Fe0
+ stishovite around 25 GPa and 3000°C. LIU (1974, 1975)
used the laser-heated diamond anvil cell to discover the key
transitions of MgSi03 to the perovskite structure, the disproportionation of Mg,Si04 spine1 to perovskite plus periclase,
and the transformation of pyrope garnet to perovskite. (The
existence of MgSi03 perovskite had been predicted by RINGWOOD 1962c, 1966c.) A review of the early ANU diamond
anvil work is given by LKJ (1979a), whilst a more comprehensive treatment of high-pressure phase transformations is
provided by LIU and BASSETT( 1986).
The other important advance in high-pressure technology
during this period was the development of the two-stage multianvil system by KAWAI and ENDO (1970) which permitted
pressures of up to 25 GPa and temperatures of -2000°C to
be generated in significant volumes (a few mm3). Using these
systems, RAWA et al. (1974) synthesized MgSiO, ilmenite
whilst IT0 et al. (1974) synthesized pure MgzSi04 spinel.
So far, we have been concerned primarily with the early
discoveries of high-pressure phases which are important in
mantle mineralogy. However, the systematic exploration and
definition of their phase relationships and the characterization
of their properties has been an enterprise of equal importance.
This work has been carried out mainly in Japanese labora-
(b) Transformations
in a Pyrolite Mantle
In Section 2, a mode1 bulk chemical composition (pyrolite)
for the Upper Mantle was formulated (Table 1). We now
wish to inquire whether phase transformations occurring in
material of pyrolite composition are capable of explaining
the variations of known physical properties (mainly seismic
velocities and densities) throughout the Transition Zone. The
most significant seismic characteristics in this region are the
two major velocity discontinuities near depths of 400 and
650 km (Fig. 1). As discussed earlier, substantial uncertainties
exist in the detailed velocity profiles, especially the magnitudes
of the velocity increases and in the intervening velocity gradients.
The phase transformations which are expected to occur
with increasing depth in a mantle of pyrolite composition
are summarized in Fig. 4. Phase relationships in the system
Mg2Si04-Fe2Si04 (Fig. 5) play a key role in determining
mantle structure around depths of 400 km and have been
successively refined during the last 25 years (RINGWOODand
MAJOR, 1966, 1971; AKIMOTO, 1972; KAWADA, 1977; AKIMOTO~~~ FUJISAWA,1966,1968; AKAOG~et al., 1989, 1984;
PYROLITE
DEPTH
km -1
DENSITY
g/cm3
100
200
300 t
3.38
OPX
OLIVINE (Mg,Fe),SiO,
I
\
G+PX
3.42
3.59
3.62
3.68
3.71
a00
0
0.2
0.4
VOLUME
0.6
0.8
1.0
FRACTION
FIG. 4. Mineral assemblages and (zero-pressure) densities displayed
by pyrolite to a depth of 850 km. It is assumed that the temperature
at 400 km is near 1400°C and at 650 km is near 1600 “C in accordance
with the mantle geotherm of BROWN and SHANKLAND (1981).
209 1
Inaugural Ingerson Lecture
1600” C
18
SPINEL (y) SOLID SOLUTION
16
8
6
OLIVINE (cc) SOLID SOLUTION
80
Mg&iOd
60
40
mole % forsterite
20
FenSi
FIG. 5. Phase relationships in the system Mg$iO.,-FezSi04 at 422 GPa and at 1600°C (after AKAOGIet al., 1989).
KATSURA and ITO, 1989). The most recent work (Fig. 5)
shows that the olivine component of pyrolite (Mgls9) would
transform to the beta-phase at a depth close to 400 km, with
a transition interval of 9-17 km and at temperatures in the
vicinity of 1400- 1600°C. This is accompanied by a density
increase of about 8% (referred to ambient P and 7) and is
believed to be primarily responsible for the major seismic
discontinuity near 400 km. AKA~GI et al. ( 1989) obtained a
gradient of 1.6 MPa/“C for this transition, which is substantially smaller than earlier estimates.
Pyroxenes and garnet with MSi03-A1203 stoichiometry (M
= Mg,Fe,Ca) constitute the second most abundant class of
minerals in the Upper Mantle. At a depth of about 300 km,
appreciable amounts of pyroxenes began to dissolve in the
preexisting garnet phase to form complex solid solutions
M3(mi,A12)Si3012 (RINGWOOD, 1967, 1975; AKA~GI and
AKIMOTO, 1977; AKAOGI et al., 1987). Up to one-quarter of
the silicon atoms in these high-pressure garnets, named “majorite,” are octahedrally coordinated. For the pyrolite composition, complete conversion of the pyroxene component
to garnet is achieved at a depth of 460 km, assuming a temperature of about 1500°C at this depth (IRIFUNE, 1987;
IRIFLJNEand RINGWOOD, 1987a). The slope of this transition
is relatively insensitive to temperature, with dP/dT - I.5
MPa/“C. The pyroxene-garnet transformation causes marked
positive velocity gradients to occur on either side of the 400
km discontinuity in a zone between depths of 300-460 km.
The transformation is accompanied by a density increase of
- 10% in the pyroxene component of the Upper Mantle.
At a somewhat greater depth, /3(Mg,Fe)zSi0, transforms
to the spine1 (y) structure (Fig. 5). Assuming a mantle temperature of about 15OO”C, the phase transformation would
occur over a depth interval between -500 and 530 km and
would be accompanied by a density increase of about 2%.
The elastic properties of @and y Mg$i04 are similar, and
the transition may be reflected mainly by a small increase in
seismic velocity gradients between these depths (WEIDNER,
1985; WEIDNER and ITO, 1987; RICDEN et al., 1991). Evidence for a small seismic discontinuity near 520 km has been
presented by SHEARER( 1990). This discontinuity may, however, be spread out over a depth interval of 30-50 km (CUMMINSet al., 1991).
The next major phase transformation occurring at greater
depths involves the formation of MgSi03 perovskite (LIu,
1974, 1975). The perovskite polymorph of MgSi03 possesses
a density of 4.10 g/cm3 which is about 4% greater than that
of an isochemical mixture of periclase plus stishovite. LIU
(1979a, 1975) and IT0 et al. (1984) showed that at temperatures in the vicinity of 16OO”C, Mg,Si04 spine1 disproportionates to MgSi03 perovskite plus MgO periclase at a pressure near 23 GPa. This matches the depth of the “650 km
discontinuity,“* variously placed between 650 and 690 km
according to different seismic studies (Fig. 1). The corresponding transformation in the system Mg$i04-FezSi04 has
been extensively studied by YAGI et al. (1979), ITO and YAMADA (1982), ITO et al. (1984), and ITO and TAKAHASHI
(1987, 1989). The latter results (Fig. 6) demonstrate that the
transition pressure is essentially constant at 23 GPa ( 1600°C)
through the composition range Mg*,,-Mg’,, and that the
transition is remarkably sharp, being completed within a
depth interval smaller than 4 km. Accordingly, the transition,
which is accompanied by a (zero-pressure) density increase
of I l%, would cause a steep seismic discontinuity at a depth
close to 650 km. The transition possesses a negative temperature gradient of -4(*2) MPa/“C (IT0 et al., 1990; IT0 and
TAKAHASHI, 1989), a characteristic which has important
geophysical implications.
Majorite garnet is the second most abundant mineral occurring in the Transition Zone (assuming a pyrolite bulk
composition). Its stability has been studied experimentally
by IRIFUNE and RINGWOOD(1987a) in a composition identical with the majorite phase of pyrolite and also by IT0 and
TAKAHASHI(1987) in a simplified composition. The former
investigation showed that majorite is stable from 16 to 20
GPa. Above 20 GPa, a CaSi03 perovskite phase is exsolved
from majorite garnet and this is joined near 22 GPa by
MgSiO, ilmenite. Exsolution of these phases causes the
1600” C
’
24.
PERO”SK,TE
MAGNESIkV~STITE
3;w
b I lbH”“l
It
ST,SH&TE
2
a
MpSiOd
SPINEL
80
Mg $30,
*Also referred to as the “670 km discontinuity.”
2 MAGNESlOWkTlTE
PEROVSKITE
MSiO~
60
mole %
40
20
FezSiO,
FIG. 6. Pseudo-binary phase diagram for the system Mg,Si04Fe2Si04at 21-26 GPa and 1600°C (after ITOand TAKAHASHI,
1989).
2092
A. E.
Ringwood
residual garnet to become richer in A1203 and MgO, approaching the pyrope composition. Compositions of coexisting 2 perovskites + garnet at 24.5 GPa, 1400°C, obtained
by IRIFUNE and RINCWOOD (1987a) are given in Table 5.
Work by ITO and TAKAHASHI(1987) on a simplified composition showed that further transformation of garnet at still
higher pressures proceeded over a significant pressure interval
and that complete transformation into an assemblage of two
perovskites plus a new highly aluminous phase was achieved
at a pressure near 27 GPa. MADON et al. (1989a) found that
the aluminous phase produced during transformation of a
naturally occurring pyrope garnet to perovskites possessed a
hollandite-like structure, with a composition close to CaMgAl&O,,.
Experiments by ITO and TAKAHASHI(1987)
and IRIFUNE and RINGWOOD (1987b) show that the substantial amounts of alumina present in the compositions
studied cause the garnet-perovskite(s) transformation to be
smeared out over a pressure interval of about 3-4 GPa and
may not be completed until depths of 700-720 km in the
pyrolite composition (Fig. 4).
It should be recognised that uncertainties in seismic velocity
distributions permit substantial tradeoffs between the magnitudes of the velocity jumps at the 400 and 650 km discontinuities and the seismic velocity gradients between 400 and
650 km (Fig. 1, Table 1). Likewise, uncertainties in the elastic
properties of mantle minerals (particularly the temperature
derivatives of the elastic moduli) hinder efforts to provide a
precise interpretation of the mineralogy appropriate to a given
seismic velocity distribution.
WEIDNER (1985), IRIFIJNE (1989), AKAOGI et al. (1987),
WEIDNER and ITO (1987), BINA and WOOD (1987), BINA
(199 1), YEGANEH-HAERIet al. (1989), and GWANMESIAet
al. (1990) have calculated the seismic velocity and/or density
distributions within the 200-700 km region of the mantle
for a pyrolite bulk composition, using the preceding evidence
(Section 3a) on mineral stability fields combined with recent
elasticity and density data for the individual mineral phases.
They conclude that the pyrolite composition is capable of
explaining the positions of the major seismic discontinuities
near 400 and 650 km, the velocity changes across the discontinuities, and the velocity gradients and density distributions between them, within the (substantial) uncertainties
mentioned above.
BASS and ANDERSON(1984) and ANDERSONand BASS
(1986) argued that the seismic velocity jump across the 400
km discontinuity and the velocity gradients below 400 km
are inconsistent with a pyrolite composition. Their conclusion
depends rather critically upon the choice of a preferred set
of seismic velocity distributions (PREM) in the mantle and
a selected set of elastic properties of minerals.
ANDERSONand BASS(1986) claim instead that the 400650 km region is composed of “piclogite” which is chemically
equivalent to an eclogite containing an additional 15-20s
of olivine. The high positive seismic velocity gradients between 400 and 600 km characteristic of the PREM model
are attributed to the persistence of clinopyroxene in this region, accompanied by its increasing solubility in garnet. According to their model, clinopyroxene is not consumed until
a depth of 600 km. However, this is contradicted by detailed
experimental studies of the stability field of clinopyroxene in
relevant compositions, which show that clinopyroxene would
be eliminated by 460 km via solid solution in garnet (IRIFUNE,
1987; IRIFUNEand RINGWOOD,1987a; IRIFLJNEet al., 1989).
Accordingly, high gradients between 460 and 600 km cannot
be caused by this mechanism. A further weakness in the Anderson-Bass model was their choice of a bulk modulus for
Table 5
COMPOSITIONS
OF CO-EXISTING PHASES PRESENT
IN A PYROWTE
COMPOSITION
(COLUMN
2)
AT 24.5 GPA AND14OO’C(FROM IRIFUNEAND RINGWOOD, 1987a)
Bulk
Mg
Garnet
Perovskite
Composition
Ca
Perovskite
SiOz
51.8
46.6
52.3
51.2
TiOz
0.5
0.1
1.8
1.5
Al203
11.4
18.7
5.3*
1.9
Cr.203
0.9
1.8
0.8
0.2
Fe0
3.1
3.0
3.7
0.6
MgO
23.1
25.3
34.8
3.1
CaO
9.5
4.1
1.0
41.3
NqO
0.9
0.9
0.1
0.3
* This value may be too high because
of electronprobe
beam overlap
on garnet
Inaugural Ingerson Lecture
majorite which is now known to have been far too high (YAGI
et al., 1987).
DUFFY and ANDERSON (1989) reexamined the issue in
the light of subsequent seismic and mineral-physics data and
concluded that the S-wave distribution implied an olivine
content of -35% whilst the P-wave distribution was consistent with 53% olivine. Their preferred olivine content was
40%. This is not so very far removed from the pyrolite composition of 55-60% olivine (RINGWOOD, 1975). The bulksound (G/p where K = bulk modulus, p = density) velocity
distribution across the 400 km discontinuity was determined
by BINA (1991) using Duffy and Anderson’s data set. Bina
found that the velocity change was best matched by an olivine
content of 68 +_ 11%.
Recent progress in the determination of accurate pressure
derivatives of elastic moduli for Transition Zone minerals
promises to clarify the question of chemical homogeneity
between 400 and 650 km (GWANMESIAet al., 1990; RIGDEN
et al., 199 1). These results suggest that the velocity contrasts
across the 400 km discontinuities for P and S waves can be
reconciled with a range of compositional models varying in
orthosilicate content from about 45 to 65%, each associated
with a different but not unreasonable choice for the values
of the (as-yet-unmeasured) temperature derivatives of elastic
wave velocities. Compressional wave velocity gradients within
the Transition Zone are also compatible with such compositional models, which include the pyrolite composition.
However, the S-wave gradients calculated from acoustic
measurements may be significantly smaller than most of the
seismically determined S-wave gradients (RIGDEN et al.,
199 1). If confirmed, this would suggest the presence of some
degree of chemical inhomogeneity within the Transition
Zone.
4. THE LOWER MANTLE
(a) Phase Relationships
BIRCH (1952) concluded that the variation of seismic velocities and density throughout the Lower Mantle between
depths of 900 and 2700 km could be explained by selfcompression of chemically homogeneous material within the
Earth’s gravitational field, and that there was no evidence for
the occurrence of major phase transformations or for substantial changes of chemical composition in this region. This
interpretation has been supported by static high-pressure investigations on the stabilities of MgSi03 perovskite and MgO
which show that these phases remain stable throughout the
entire Lower Mantle (KNITTLE and JEANLOZ, 1987; MAO
and BELL, 1977).
The distribution of iron between perovskite and magnesiowiistite has been studied by YAGI et al. (1979), IT0 et al.
(1984), GUYOT et al. (1988), and ITO and TAKAHASHI(1989)
who demonstrated a strong preferential partition of Fe0 into
magnesiowtistite. This effect tends to decrease with depth.
The compositions of Mg-perovskite and magnesiowtistite
below 1000 km in a mantle of pyrolite composition would
be Mg’,,., and Mg*80,respectively (GUYOT et al., 1988).
LIU (1977) demonstrated that up to 25 mol% of A&O3 can
be accommodated in solid solution in MgSi03 perovskite. In
a study of the transformation of a natural pyrope garnet to
2093
perovskite, MADON et al. (1989b) showed that perovskite
(approximately Mg’,,) was capable of accommodating up to
20 atom% of Al in the octahedral Si site at a pressure of 50
GPa. (At higher Al contents an additional aluminous phase,
CaMgA1&0i6, would be present.) These results suggest that
with the lower (-4%) A&O3 contents characteristic of the
pyrolite composition, the magnesian perovskite phase would
be capable of accepting all of the AllO3 into solid solution
throughout most of the Lower Mantle, and that an additional
aluminous phase is unlikely to be present (contrary to a suggestion by RINGWOOD, 1982).
Calcium silicate (CaSi03) adopts the perovskite structure
at high pressures (RINGWOODand MAJOR, 197 1; LIU, 1975).
Experimental investigations by IRIFUNE and RINGWOOD
(1987a) show that in the Lower Mantle, calcium would be
contained in CaSi03 perovskite and that solid solution between this phase and coexisting MgSi03 perovskite is very
limited (Table 5). These results also show that the modest
sodium content of the Lower Mantle would be accommodated in solid solution within a CaSi03 perovskite mineral.
The experimental evidence cited above indicates that the
mineral assemblage adopted by pyrolite in the Lower Mantle
would comprise MgSiOJ perovskite + CaSi03 perovskite
+ (Mg,Fe)O magnesiowtistite. This mineral assemblage was
recently identified in a natural (fertile) peridotite composition
crystallized at 54 GPa and 1900°C (O’NEILL and JEANLOZ,
1990).
(b) Composition
The nature of the major seismic discontinuity near 650
km has been the subject of considerable debate. Its depth is
coincident with a major phase transformation to perovskite
which occurs both in orthosilicate (M2Si04) and metasilicate
(MSi03) stoichiometries. The question under debate is
whether the discontinuity is primarily caused by the isochemical disproportionation of silicate spine1 to perovskite
plus magnesiowilstite or whether, in addition, there is a
change in chemical composition at this depth.
Chemical models proposed for the Lower Mantle range
between the pyrolite composition and an essentially pure
perovskititic mineralogy which would be consistent with a
chondritic bulk MgO/SiO* ratio for the entire mantle. Extensive shock-wave data on olivines and pyroxenes at Lower
Mantle pressures and temperatures show that for the same
MgO/(MgO + FeO) ratio, the density difference between these
two model mantle compositions is smaller than 0.06 g/cm3
(WATT and AHRENS, 1986). Further ambiguity is introduced
by uncertainties over the temperature distribution in the
Lower Mantle, by the effects of other minor components
(CaO, A1203),by the possibility of additional minor (probably
second-order) phase changes in the Lower Mantle, and, finally, by experimental uncertainties in the existing mineralphysics data base. It is evident that it would be extremely
difficult, if not impossible, to determine whether the Lower
Mantle is of perovskititic or pyrolitic composition from density data alone. An identical dilemma is encountered in comparisons of the elastic properties of the Lower Mantle with
those estimated for pyrolite and perovskite compositions.
Most workers who have discussed the likely composition
2094
A. E. Ringwood
of the Lower Mantle in the light of elasticity and density data
have concluded that they are consistent with the interpretation
that the Lower Mantle possesses a similar chemical composition to the Upper Mantle, within the uncertainties of the
seismic and mineral-physics data base (e.g., RINGWOOD, 1975;
JACKSON, 1983; ITO et al., 1984; WATT and AHRENS, 1986;
WEIDNER and ITO, 1987; WOLF and BUKOWINSKI, 1987;
CHOPELASand BOEHLER, 1989; BUKOWINSKIand WOLF,
1990). However, the experimental and observational uncertainties do not preclude the possibility of a change in chemical
composition at the 650 km discontinuity. Thus, a perovskititic
Lower Mantle may be permitted, particularly if the Fe0 content is increased to offset the effect of increasing SiOZ content
(e.g., JACKSON, 1983).
In contrast to the above conclusion, some workers have
concluded that the geophysical properties of the 650 km discontinuity and the Lower Mantle are inconsistent with a pyrolite composition and that a change in chemical composition, involving an increase in Fe0 content and possibly also
of SiOz, is definitely required (e.g., ANDERSON, 1977, 1983;
ANDERSONand BASS,1986; LILJ,1979b; KNITTLEet al., 1986;
JEANLOZand KNITTLE, 1989). Most of these workers, and
also others (e.g., HERZBERG and O’HARA, 1985; OHTANI,
1985; TAKAHASHI, 1986; AGEE and WALKER, 1988b), have
preferred a perovskititic mineralogy (metasilicate, MSi03
stoichiometry) for the Lower Mantle.
Without exception, the arguments used by these authors
to justify this conclusion have proved to be seriously flawed.
LIU (1979b) claimed that the depth of the 650 km discontinuity was not consistent with the pressure at which
(Mg,Fe)$i04 spine1 transformed to perovskite + magnesiowtistite. However, significant errors are now believed to
have occurred in Liu’s pressure calibrations (JEANLOZand
THOMPSON,1983). Subsequently, more accurate phase equilibria studies have shown that the pressure of the transformation (ITO and TAKAHASHI, 1989) is consistent with the
depth of this discontinuity. LEES et al. (1983) argued that the
650 km discontinuity is too sharp to be caused by the above
phase transformation; however, this conclusion is contradicted by the recent experiments of ITO and TAKAHASHI
(1989).
JACKSON( 1983) demonstrated that the viability of pyrolite
versus perovskititic models for the Lower Mantle would be
critically dependent on the mean thermal expansion coefficient (5) of MgSi03 perovskite between about 25 and 1600°C.
If this value were relatively “high,” exceeding 25 X 10-6/oC,
a perovskititic Lower Mantle would indeed be favoured. On
the other hand, if (Ywere smaller than 25 X 10-6/oC, a pyrolite
composition would be consistent with the observed density
of the Lower Mantle.
KNITTLE et al. (1986) and ROSS and HAZEN (1989) measured the thermal expansion coefficient of MgSiOJ perovskite
at room temperature, both groups obtaining a value of 22
X 10-6/oC. Knittle et al. attempted to measure (Yat higher
temperatures and observed an anomalous increase to -40
X 10-6/oC at -600°C. Taking this at face value, KNITTLE
et al. (1986) and JEANLOZand KNITTLE (1989) concluded
that the “high” thermal expansivity of MgSi03 perovskite
implied that the Lower Mantle must be substantially richer
in iron than the Upper Mantle (in accordance with the study
of JACKSON, 1983). However, the results of HILL and JACKSON(1990) and CHOPELASand BOEHLER(1989) indicate that
KNITTLE et al. (1986) overestimated the thermal expansivity
of MgSi03 perovskite. CHOPELAS and BOEHLER (1989)
showed that for a wide range of materials, the variation of (Y
at high pressure follows a relationship (a In al/a P)= = 5.5
? 0.5, which implies a considerable decrease in Cuat Lower
Mantle pressures. Thus, at the core-mantle boundary they
found that 01is reduced to only 5 X 10d6.
CHOPELASand BOEHLER(1989) calculated the densities
of perovskite and magnesiowiistite (Mg’**)appropriate to the
pyrolite composition using standard procedures but including
the effect of pressure on (Y.The calculated density for the
pyrolite model throughout the Lower Mantle was found to
agree closely with the seismically determined density
throughout this region, showing clearly that the pyrolite
composition remains viable. Chopelas and Boehler’s estimate
of the thermal expansion coefficient of MgSiO, perovskite at
high pressures has recently been confirmed by direct measurements (WANG et al., 1991). I conclude that the density
of the Lower Mantle is consistent with a pyrolite composition,
but that uncertainties in the data base do not preclude the
possibility that the Lower Mantle may possess a significantly
different composition.
A related constraint on Lower Mantle compositional
models can be obtained from the observed density distribution throughout this region. Assuming that the Lower Mantle
is chemically homogeneous, the observed p versus P relationship can be extrapolated to zero pressure using an appropriate equation of state (CLARK and RINGWOOD, 1964).
The decompressed density thereby obtained can be compared
with different mineralogical models with the aid of known
zero pressure densities of individual minerals and their corresponding thermal expansion coefficients. The most rigorous
investigation of this type yet carried out was reported by BUKOWINSKI and WOLF (1990). They demonstrated that a
Lower Mantle of pyrolite composition was consistent with
the Lower Mantle density distribution and that perovskititic
models were “far less satisfactory than pyrolite-like models”
although they could not be conclusively eliminated. On the
other hand, they found that a class of models possessing higher
levels both of SiOz and Fe0 than would be present in a perovskititic stoichiometry were permitted. However, it is difficult to justify these latter models on geochemical grounds.
(c) The Lowermost (D”) Region of the Mantle
It has long been known that the lowermost region of the
mantle immediately overlying the core displays anomalous
seismic velocities (e.g., JEFFREY& 1939; BULLEN, 1949; GuTENBERG,1958, 1960). The thickness of this anomalous region (defined by Bullen as the D” layer) is believed to be
about 200 km. Seismic P- and S-wave velocity gradients are
smaller than elsewhere in the Lower Mantle and may even
be negative. Moreover, the seismic properties of the D” layer
imply the presence of a high degree of lateral heterogeneity
(e.g., YOUNG and LAY, 1987; DZIEWONSKIand WOODHOUSE,
1987).
It has been suggested that the anomalous seismic properties
could be caused by the presence of a chemical boundary layer
2095
Inaugural Ingerson Lecture
(e.g., RINGWOOD,1959, 1961) or of a thermal boundary layer
(e.g., STACEYand LOPER, 1983) or perhaps by a combination
ofthese effects (LAY, 1989). Interpretation of D” as a thermal
boundary layer would imply that a superadiabatic temperature gradient of a few degrees per kilometer exists in this
region (STACEY and LOPER, 1983). This interpretation seemed
to be supported by measurements of the effect of high pressure
on the melting point of iron by WILLIAMSet al. (1987) which
would imply that the Outer Core (near the core-mantle
boundary) was about 1000°C hotter than the Lower Mantle
(at depths of 2500-2700 km). However, a subsequent careful
redetermination of the melting point of iron at high pressure
by BOEHLERet al. (1990) indicates that the core is probably
much cooler than was inferred by WILLIAMSet al. (1987),
thereby casting doubt on the need for a thermal boundary
layer in the D” region.
It is well known that the Outer Core contains substantial
amounts of one or more light elements in solution in molten
iron. RINGWOOD(1977, 1984) advanced several arguments
which suggested that oxygen is the principal light element in
the core. Experimental investigations of the system Fe-Fe0
at high pressures by KATO and RINGWOOD(I 989) and RINGWOOD and HIBBERSON(1990) provided strong support for
this hypothesis and showed that the solution of oxygen would
cause a large depression of the melting point of iron, thereby
causing the metal phase to melt at substantially lower temperatures than mantle pyrolite. Their results, in conjunction
with those of BOEHLERet al. (1990) confirm that the case
for a strong thermal boundary layer involving a superadiabatic
temperature change of 500-1000°C across the D” region is
decidedly weak. A milder superadiabatic temperature gradient
cannot be excluded, but this would not have a large effect
upon seismic velocity gradients in the D” region.
RINGWOOD (1959, 1961) pointed out that the process of
core-formation within the Earth would be likely to lead to a
condition of chemical disequilibrium between core and
mantle at their mutual interface. He suggested that chemical
reactions consequently
occurred near the core-mantle
boundary (CMB), leading to transport of components (Si,
Fe, and Ni) across the boundary, and that the anomalous
seismic properties of the lowermost 200 km (D”) zone were
caused by these processes. Ringwood estimated that the
chemical disequilibrium would generate an EMF in the vicinity of I volt across the CMB and that under certain conditions, charge separation might occur with electrons and
ions following different paths, leading to the generation of
electrical currents across the CMB. (These effects might be
enhanced if perovskite in the Lower Mantle should indeed
be in a superionic conductive state as argued by PRICE et al.,
1989.) RINGWOOD(1959) suggested that the electrical currents
thereby produced might play a significant role, in conjunction
with the core dynamo, in generating the Earth’s magnetic
field.
Some of these ideas have recently been exhumed by KNITTLE and JEANLOZ(1989) and JEANLOZ(1990) who also propose that the core is not in equilibrium with the mantle and
that chemical reactions accordingly occur at and near the
CMB, ultimately producing the seismic heterogeneities observed in the D” layer. JEANLOZ(I 990) maintains that this
is the most dynamically and chemically active region in the
mantle. This interpretation was based on observations by
KNITTLE and JEANLOZ (1989) that metallic iron and
(Mg,Fe)SiOs perovskite chemically react with one another at
very high pressure (270 GPa) and temperature (23000°C).
Similar observations have been reported by GUYOT et al.
(1988).
The detailed nature of these reactions was studied by
RINGWOODand HIBBERSON(199 1) who showed that several
common metallic oxides are soluble in molten iron at high
P and T, and at oxygen fugacities near the iron-wiistite buffer.
Individual oxides dissolved quasi-congruently, but there are
large differences in the solubilities of individual oxides. Fe0
is much more soluble than SiOz, which in turn is much more
soluble than MgO. Thus, at the CMB, molten iron would
selectively dissolve firstly Fe0 and then SiOZ, leaving a layer
enriched in periclase. The rate of this process would depend
on the ionic conductivity of the lower-most mantle and would
be accelerated if this region were in the superionic conductive
state (PRICE et al., 1989). The periclase-enriched chemical
boundary layer would possess lower density and seismic velocities than the overlying region. Because of the resultant
gravitational instability of this layer, diapirs of periclase-enriched material would ascend from the CMB. The seismic
heterogeneity of the D” region might be caused by the presence
of large numbers of these diapirs. Ultimately, it is expected
that convective mixing in the Lower Mantle would cause the
diapirs to be dispersed, thereby producing a relatively high
level of chemical homogeneity at depths above 2700 km.
5. GROSS MANTLE DIFFERENTIATION
AND THE Mg/Si RATIO
We have seen that geophysical evidence permits the Lower
Mantle to be either of pyrolitic or perovskititic composition.
The former alternative might be favoured on the grounds of
economy of hypotheses. Moreover, we should note that it
would be coincidental for the depth of a major change in
mantle chemical composition to be identical with the depth
at which (Mg,FehSiOd spine1 happens to transform to perovskite plus magnesiowtistite. Nevertheless, as noted previously, many scientists have continued to prefer a perovskititic Lower Mantle with metasilicate stoichiometry. In
many cases, the primary reason for this preference is that the
perovskititic interpretation would permit the bulk mantle to
possess a chondritic Mg/Si ratio. A mantle of pyrolite composition possesses an Mg/Si ratio of 1.27 (atomic) whereas
the chondritic ratio is 1.05. The pyrolite model thus requires
an explanation for the apparent deficiency of SiOZ in the
mantle.
Advocates of a chemically zoned mantle comprising a pyrolite outer region and a perovskititic Lower Mantle have
called upon the processes of crystallization differentiation
from the molten state or of extensive partial melting of the
mantle to achieve this state (e.g., KUMAZAWA,198 1; OHTAN],
1985; HERZBERGand O’HARA, 1985; ANDERSONand BASS,
1986; TAKAHASHI, 1986; AGEE and WALKER, 1988b). This
hypothesis can be tested directly by appropriate experiments.
As discussed in Section 2, a wide range of involatile lithophile
elements is present in the Upper Mantle in near-chondritic
relative abundances (Table 3). This observation provides a
strong constraint on proposed global differentiation processes.
A, E.
2096
Consider the crystallization of a completely molten mantle.
MgSi03 perovskite is found to be the liquidus phase on chondritic and pyrolite compositions below about 700 km (IT0
and TAKAHASHI, 1987), and because of its high relative
abundance (-80% of a chondritic Lower Mantle) it is expected to display a broad crystallization field before being
joined by magnesiowiisite (- 10%). Casio3 perovskite is not
expected to crystallize as a liquidus phase because of its low
abundance (- 5%). The partition coefficients of many minor
elements between silicate perovskites, garnets, and a range
of ultrabasic liquids were determined by KATO et al. (1988a).
Some of their results are shown in Table 4. Note that the
partition coefficients of Zr, Hf, and SC (Group A) between
MgSiOs perovskite and liquid are much higher than unity,
whilst the corresponding partition coefficients of the light
and intermediate rare earths (Group B elements) are much
smaller than unity. These results show that quite modest degrees of fractionation of MgSiOj perovskite, whether caused
by crystallization from the melt or by partial melting processes, would result in rather drastic fractionations between
Group A and B elements in the residual liquid. An example
is given in Fig. 7. It is seen that a large degree of MgSi03
perovskite fractionation is necessary in order to cause substantial changes in the Si/Mg ratio during crystallization of
a chondritic mantle from the molten state. However, the separation of only 5-10% of MgSi03 perovskite would have
driven the Sm/Hf and Sc/Sm ratios of the residual melts well
outside the near-chondritic ratios which are presently observed in the extensive Upper Mantle source regions of midocean ridge basalts. The results of KATO et al. ( 1988a) imply
rather strongly that the present mantle does not reflect the
operation of comprehensive perovskite-controlled fractionation postulated by the authors cited above, whether or not
these arise from global fractional crystallization or partial
melting processes. They strongly suggest that the bulk composition of the Lower Mantle is similar to that of the Upper
Mantle.
0.8
0.6
0
5
Mg-perovsklte
10
fractionation,
%
FIG. 7. Variation of Sm/Hf
perovskite fractionation from
and Si/Mg ratios as a function of Mga chondritic mantle composition, calculated from partition coefficientsgiven in Table 4. The Si/Mg ratio
of the present Upper Mantle is also indicated. Shaded areas indicate
uncertainties which would be caused by errors of f2 in the Sm/Hf
and Sc/Sm ratios used in the calculation
(from UT0
et al., 1988a).
Ringwood
It appears that either (i) the mantle was initially formed
in an approximately homogeneous state, or (ii) ifglobal melting and differentiation indeed occurred during, or soon after,
formation of the Earth, the record of these processes has been
destroyed by subsequent subsolidus convection, which rehomogenised the mantle. RINGWOOD(1990a) has advanced
several arguments which support the first, and are unfavourable to the second, of the two alternatives. However,
these are beyond the scope of the present review. The primary
conclusion is that the Mg/Si ratio of the Upper Mantle is
similar to that of the Lower Mantle and is substantially higher
than the chondritic ratio. Possible causes of this deviation
from the chondritic Mg/Si ratio are considered below.
RINGWOOD (1989a) showed that the similar A&O3contents
of terrestrial and Venusian basalts imply that the ratio of
normative olivine to pyroxene in the Venusian mantle resembled that of the terrestrial mantle and that the Venusian
Mg/Si ratio was therefore terrestrial rather than chondritic.
He also showed that geochemical relationships among pallasites, howardites, diogenites, and eucrites indicate that the
Mg/Si ratio of the eucrite parent body was closer to the terrestrial than to the chondritic value. Moreover, he noted that
the spectra of S-type asteroids which predominate in the inner
asteroid belt imply high olivine/pyroxene ratios combined
with high metal contents. These relationships suggest that Sasteroids possess terrestrial rather than chondritic Mg/Si ratios. This array of evidence suggests that the terrestrial mantle
Mg/Si ratio is not unique to the Earth but may be representative of the inner region of the solar system between Venus
and the inner asteroid belt.
RINGWOOD(1989b) suggested that, rather than being depleted in Si, the Earth’s mantle and the inner solar system
may actually possess the primordial solar Mg/Si ratio and
that CI chondrites may have been enriched in Si via a cosmochemical process. (Nevertheless, the Mg/Si ratios of the
Earth’s mantle and CI chondrites are both consistent with
the Mg/Si ratios of the solar photosphere within the errors
of measurement of the latter; ANDERSand GREVESSE,1989.)
RIETMEIJER(1987, 1988) showed that both chondritic interplanetary dust particles and Halley’s Comet dust particles
from the outer parts of the solar system are substantially enriched in relatively volatile elements (Mn, Cu, K, Na, S, Zn,
Bi) and also in silicon relative to CI chondrites (outer asteroid
belt). He suggested that there had been a radial chemical
differentiation of volatile elements (including Si) between the
two source regions in the inner and outer solar system and
that the existence of this chemical zonation challenges the
assumption that the CI chondrites possess primordial compositions. An alternative possibility suggested by RINGWOOD
(1979) is that CI chondrites indeed possess the primordial
(solar) Mg/Si ratio, but that there has been a significant depletion of silicon (accompanied by more volatile elements)
from the inner solar system via high-temperature processing
connected with the early evolution of the protosun and solar
nebula.
6. PHASE TRANSFORMATIONS AND DENSITY
RELATIONSHIPS IN SUBDUCTED
OCEANIC LITHOSPHERE
So far, this review has focussed on the characteristics of a
mantle composed of pyrolite. However, it should be recog-
2097
Inaugural lngerson Lecture
MATURE PHANEROZOIC
ASEISMIC RIDGES
ARCHEAN OCEANIC
OCEANIC LllHOSPHERE
OCEANIC PLATEAUS
LlTHtXPHERE
...................................................
...................................................
...................................................
...................................................
...................................................
...................................................
..................................................
...................................................
...................................................
...................................................
...................................................
...................................................
::f::$::;::;::RESIDUAL
:::j:::::::::::::::::i::::::::~
Fo
DUNlTE
:;::;:::;::;:;:::
.................
~:.~~.~.~~.~.-.~_-.~:,~:.~
92-94
...................
...................................................
..........................
...................................................
SLIGHTLY DEPLETED PYRCMTE 1
(minus0.1-l% alkallc
liquid
62% 01
36% Opx, Cpx, Gnt
60
c
La
sn
ml
(b)
(a)
(c)
FIG. 8. Idealized structure of oceanic lithosphere showing chemical and petrological zoning developed during partial
melting and differentiation: (a) current mid-oceanic spreading centrcs; (b) oceanic plateaus and aseismic ridges developed
over hot spots; and (c) postulated structure of the Archaean oceanic lithosphere.
nised that the formation of new lithosphere at mid-oceanic
ridges is accompanied
by melting and differentiation,
leading
to a petrologically zoned structure. Oceanic lithosphere displays marked variations in its vertical stratification.
The
structure
of normal lithosphere
formed at mid-oceanic
spreading centres during most of the Phanerozoic
is shown
in Fig. 8a. The basaltic crust, -7 km thick, overlies a layer
of harzburgite 5-20 km thick which is underlain by lherzolite
and pyrolite which have experienced successively smaller degrees of partial melting (e.g., RINGWOOD, 1982). Typical
compositions
of these lithologies are given in Table 6. There
are also extensive oceanic regions where the basaltic crust is
Table 6
CH~ICAL COMPOSITIONS
Harzburgitecb)
43.6
MORB(C)
49.7
SiO2
44.6
Al203
4.3
0.6
16.4
cao
3.5
0.5
13.1
MgO
38.1
46.5
10.1
FeO
9.1
8.8
8.0
Na20
0.4
2.0
Ti02
0.7
88.2
(a)
iW
(c)
90.5
69.2
(Sun, 1982)
(Michael and Bonatti, 1985) (corrected for NiO, MnO and CrO)
(Green, et al., 1979) (primitive MORB)
FeO* = Fe0 (total) + NiO + MnO = CrO
2098
A. E. Ringwood
they are accompanied by substantial density contrasts which
may be of considerable geodynamic significance.
The sequence of phase transformations displayed by pyrolite with increasing depth is shown in Fig. 4. The zeropressure densities of the various mineral assemblages can be
accurately calculated from those of their individual components and are also shown in Fig. 4. Mineral assemblages and
zero-pressure densities for subducted basaltic oceanic crust
are shown in Fig. 9 which is based on experiments by IRIFUNE
and RINCWOOD (1987a). These authors showed that subducted oceanic crust of MORB composition would remain
about 0.1-0.2 g/cm3 denser than pyrolite down to 650 km
and below about 750 km. However, between about 650 and
780 km, a large amount ofgamet remains stable in the MORB
composition, whereas pyrolite has essentially transformed to
the denser perovskite-magnesiowiistite assemblage. Because
of the persistence of garnet, the MORB composition is actually about 0.1 g/cm3 less dense than pyrolite over this depth
interval.
RINGWOOD(1990b) noted that oceanic crust is probably
subjected to partial melting during subduction, resulting in
the elimination of excess SiOz (stishovite) and that it would
be more realistic to calculate the density of subducted oceanic
crust at depths below about 500 km on the basis of stishovitefree mineral assemblages. The densities of this modified
oceanic crust and pyrolite are shown as a function of depth
much thicker than normal and is believed to be underlain
by thicker sections of harzburgite. These include oceanic hot
spots, oceanic plateaus, and passive continental
margins
(MCKENZIE and BICKLE, 1988). In these regions, the basaltic
oceanic crust may be as much as 20 km thick (Fig. 8b).
Mantle temperatures are believed to have been 200-300°C
higher in the Archaean than at present (MCKENZIE and
WEISS, 1975; BICKLE, 1978, 1986). Higher mantle temperatures would be expected to have produced more advanced
degrees of partial melting, leading to copious generation of
komatiitic magmas and a thickened oceanic crust (BICKLE,
1978, 1986). The refractory residuum remaining after extraction of these magmas would have consisted primarily of
dunites or orthopyroxene-poor harzburgite containing ohvine
of about Mg*94composition (GREEN, 1975). The structure
expected for Archaean oceanic lithosphere is illustrated in
Fig. 8c,
The formation of chemically differentiated oceanic plates
is a major geodynamic process which may have involved a
large proportion of the volume of the mantle. As differentiated
oceanic lithosphere plates descend into the mantle, they experience a complex series of successive phase transformations
with increasing depth. The sequences and characteristics of
these phase transformations are considerably influenced by
the differing chemical compositions of the basalt, harzburgite,
dunite, and depleted pyrolite layers (Fig. 8). In some cases,
DENSITY
DEPTH
g/cm3
BASALTIC OCEAN CRUST
km
100
3.50
CLINOPYROXENE
200
\
300
/
400
\
GARNET
ST
500
600
700
LiIF
SH-
I
600
Mg PEROVSKITE
0.1
0.2
0.3
+
0.4
VOLUME
0.5
0.6
I
I,
Ca PEROVSKITE
0.7
ITE
I
0.8
0.9
1.0
FRACTION
FIG. 9. Mineral assemblages and (zero-pressure) densities displayed by subducted basaltic (MORB) oceanic crust to
a depth of 850 km.
2099
Inaugural Ingerson Lecture
in Fig. 10 and the corresponding
density difference between
pyrolite and oceanic crust is shown in Fig. 11. It is seen that
elimination
of stishovite would substantially
enhance the
buoyancy of subducted oceanic crust (relative to surrounding
mantle) between depths of 650 and 780 km.
Mineral assemblages and associated zero-pressure densities
displayed by harzburgite to a depth of 850 km are shown in
Fig. 12. It is assumed that the harzburgite follows the same
geotherm as pyrolite in Fig. 4, with temperatures of - 1400°C
at 400 km and - 1600°C at 650 km. Stability fields shown
for mineral assemblages in Fig. 12 are based mainly on references discussed in Section 3, together with the experimental
results of IRIFUNE and RINCWOOD (1987b). The density of
harzburgite is compared with that of pyrolite (along the same
geotherm) in Fig. 10, and the density difference between
harzburgite and pyrolite as a function of depth is shown in
Fig. 11. Harzburgite is seen to be 0.05-0.08 g/cm3 less dense
than pyrolite to a depth of 500 km, owing to its lower Fe0
content and also to its smaller A&O3 content which is responsible, in turn, for a smaller amount of dense garnet.
Harzburgite is also -0.05 g/cm3 less dense than pyrolite below 750 km because of its lower Fe0 and higher M2Si04/
MSi03 ratios (RINGWOOD, 1982). However, between depths
of 650 and 700 km, harzburgite is up to 0.08 g/cm3 denser
than pyrolite. This is because harzburgite fully transforms to
a perovskite-magnesiotistite
assemblage at 650 km, whereas
the broader stability field of garnet in pyrolite extends the
completion of this transformation
to about 720 km.
The implications of these density relationships in subducted
oceanic lithosphere for processes of mantle dynamics are discussed in the following section.
7. PHASE TRANSFORMATIONS
AND MANTLE DYNAMICS
(a) Viscosity
Stratification
in the Mantle
THERMALLY
+0.3 - (a) Basalt
c
fi
E
+O.l
r (b) HarzburgiteminusPyrolite
0.0
I
SLAB
I
I
I
I
I
I
:
I
I
I/!\
A
: _______
I
I
I
I
-0.1 -
I
-0.2
I I
300
I
500
400
1
I
700
600
I
800
DEPTH km
FIG. 11. Density differences (obtained from Fig. IO), between a
modified (stishovite-free) MORB composition and pyrolite, and between harzburgite and pyrolite as a function of depth along the
BROWN and SHANKLAND (198 1) geotherm.
with depth in the mantle will be influenced by the successive
phase transformations
which occur with increasing pressure.
KARATO (1989) demonstrated
the existence of systematic relationships
between (normalized)
temperatures
and flow
stresses for families of dense oxides possessing crystal structures relevant to the Earth’s mantle. These relationships suggested that intrinsic viscosities of crystal structures decreased
in the sequence: garnet > spine1 2 olivine 2 perovskite
> rocksalt. Karat0 noted that this sequence is consistent with
observations
of the relative strengths of these minerals in
deformed rocks and high-pressure experimental
charges. He
drew attention to some geophysical implications of these observations, including a high viscosity for the Transition Zone
as compared to the Upper Mantle.
MEADE
Crystal structure has an important influence on the rheology of solids. Accordingly, it is likely that variation of viscosity
I
EQUlLlSFiATED
minus Pyrolite
and
JEANLOZ
strengths of the principal
(1990)
determined
silicate minerals
HARZBURGITE
DEPTH
the
relative
of the mantle
by
IENSITY
g/cm3
THERMALLY
EQUILIBRATED
SLAB
4.6
44
:
OLIVINE
;
Gi
5
(Mg,Fe),SiO,
42
4.0
3.32
3.8
3.6
3.4
200
- -
3.69
300
400
500
DEPTH
600
700
800
600
3.71
km
4.10
FIG. 10. Density profiles displayed by pyrolite, harzburgite, and
modified MORB compositions along the geotherm of BROWN and
SHANKLAND(1981) at depths between 200 and 900 km in the mantle.
Densities have been calculated from experimentally determined phase
assemblages in these compositions, corrected for compressibility and
thermal expansion along the above geotherm (based on IRIFUNE and
RINGWOOD, 1987b; RINGWOOD
and IRIFUNE, 1989;and RINGWOOD,
1990b). Densities of the MORB composition below 500 km have
been calculated on a stishovite-free basis as discussed in the text.
Mg-PEROVSKITE
0
0.2
0.4
0.8
1.0
VOLUME FRACTION
FIG. 12. Mineral assemblages and densities displayed by subducted
harzburgite to a depth of 850 km.
2100
A. E. Ringwood
direct measurements in a diamond anvil cell at high pressures
and ambient temperature. They demonstrated that strength
decreased in the order: spine1 > perovskite > olivine. in
agreement with Karato’s results. They also cited literature
evidence relating to the high strength of garnet, relative to
other Upper Mantle minerals. MEADE and JEANLOZ(1990)
concluded from these results that the Transition Zone may
represent a region of high viscosity in the mantle, as compared
to the Upper and Lower Mantle. This conclusion, if subsequently verified, would be of considerable geophysical significance. For example, it could have an important influence
on the style of convection in the Earth’s mantle.
The methodologies employed by the above workers assume
that mantle viscosity is governed mainly by dislocation creep.
This assumption seems well founded for the Upper Mantle
(KARATO et al., 1986). The possibility that a different mechanism (diffusion-controlled creep) might predominate in the
Lower Mantle should nevertheless be entertained. If so, it
would imply even smaller viscosities in the Lower Mantle.
WALL and PRICE (1989) calculated the self-diffusion characteristics of MgSiOj perovskite and obtained an electrical
conductivity in good agreement with observed values for the
Lower Mantle. The large magnitude of this conductivity suggested that the Lower Mantle might be in a superionic conductive state, which in turn implies high atomic diffusivities
and would favour a relatively low viscosity for this region.
A number of authors-including
HAGER ( 1984), HAGER
and RICHARDS (1989) GURNIS and DAVIES (1986), and
DAVIES and RICHARDS (199 I)-have developed models of
whole-mantle convection which depend rather critically upon
the assumption that the viscosity of the mantle increases by
factors of 30-100 on passing from the Transition Zone to
the Lower Mantle. The evidence for this assumption is weak.
Although not yet definitive, the available evidence from mineral physics suggests that the Lower Mantle may possess a
lower viscosity than the Transition Zone.
ALL~GRE (pers. comm., 1988) suggested that a high-viscosity Transition Zone would enhance the tendency for layered convection in the mantle. This is a potentially important
effect which warrants further study. MCCULLOCH (199 1)
pointed out that a high viscosity in the Transition Zone may
have caused the Upper Mantle to convect independently of
the remainder of the mantle during the Archaean. The Archaean crust would then have been extracted only from a
reservoir shallower than 400 km. This could provide an explanation of the relatively strong geochemical “depletion signatures” that are indicated by the neodymium isotopic compositions of the oldest Archaean rocks (23.8 Ga).
(b) Collision of Slabs with the 650 km Discontinuity:
The Formation of Megaliths
GIARDINIand WOODHOUSE(1984) concluded from a study
of deep seismicity in the Tongan subduction zone that the
descending slab encounters strong resistance at the 650 km
discontinuity and is deflected laterally. More generally,
earthquake source mechanisms show that the lower sections
of slabs (below 300 km) are in a state of compressive stress
(ISACKS and MOLNAR, 197 1) which would be consistent with
the presence of some kind of barrier to slab penetration at
650 km. HAGER (1984) and HAGER and RICHARDS(1989)
suggested that the resistance to subduction below 650 km is
caused by the high viscosity of the Lower Mantle. However,
as noted in the previous section, this interpretation is not
supported by experimental evidence on the viscosities of
mantle minerals.
Nevertheless, there are two other factors which would cause
the tips of slabs to experience buoyant resistance when they
encounter the 650 km discontinuity. Firstly, the interiors of
old slabs are likely to be a few hundred degrees cooler than
surrounding mantle at 650 km (OXBURGH and TURCOTTE,
1970). Because of the negative gradient of the transition of
spine1 to perovskite + magnesiowtistite (Section 3b), the spine1
+ garnet field (Fig. 4) would extend 30-40 km below the 650
km discontinuity in the interior of a slab which was -400°C
cooler than surrounding mantle. The lower density of the
spine1 would cause a significant buoyant stress at the tip of
the slab. Secondly, it was shown in Section 6 that the former
basalt and harzburgite components of the slab are 0.20-0.05
g/cm3 less dense than surrounding mantle in the region between 650 and 750 km (Fig. 11). Because of the combination
of these two buoyancy effects, the tip of the slab may be
subjected to a resistive stress of -500 bars (0.05 GPa) as it
sinks below 650 km.
RINGWOOD(1982, 1986, 1989b, 1990b) and RINGWOOD
and IRIFUNE (1988) have argued that the combined buoyancy
stresses would cause the tip of the slab to buckle when it
encounters the 650 km discontinuity. Buckling may also be
accompanied by plastic thickening. If the slab were, for example, about 400°C cooler than surrounding mantle, its mean
viscosity would intensify the buckling process. After a limited
amount of buckling (thickening) had occurred at the tip of
a subsiding plate, additional penetration through 650 km
would be further impeded by the accumulation of high viscosity material at its base. Continuation of this process over
an extended period, e.g., IO* years, would lead to the development of a large “megalith” of relatively cool but deformed
former oceanic lithosphere, with a mean cross-sectional diameter of several hundred km (Fig. 13). The shapes of megaliths are expected to be highly variable and would be dependent upon the relative rates of subduction versus trench
migration. High subduction velocities and low trench migration velocities would yield vertically elongated megaliths,
whereas slow subduction and high trench migration velocities
would yield megaliths with relatively large horizontal dimensions.
Megaliths are believed to be comprised of a melange of
former oceanic crust and harzburgite (Fig. 13). The mean
density of this differentiated assemblage is slightly smaller
than that of chemically equivalent pyrolite (Fig. 11).However,
the megalith is likely to be a few hundred degrees cooler than
the surrounding mantle, which would tend to cancel the density deficit, causing it to be neutrally buoyant. Megaliths provide “cold fingers” or heat sinks in the Lower Mantle and
may generate sinking convection currents in this region. The
integrity of megaliths is maintained, initially. by high viscosity
arising from their relatively lower temperature. As the outer
regions of megaliths are warmed, their viscosity falls and they
are likely to be entrained into the convective system of the
Lower Mantle. Ultimately, this process of thermal erosion
would cause the entire megalith to be mixed into the Lower
Mantle. Individual megaliths are thus expected to have a
Inaugural lngerson Lecture
island arc
Calc-alkaline volcanism
LITHOSPHERE
Resorbtion of
depleted pyrolite
into Upper Mantle
PYROLITE
by melts derived from
former basaltic crust
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . ...!
400 km discontinuity
Accretion of
fertilized peridotite \
650 km discontinuity
Former basaltic crust
Former harzburgite
*:.:,:.:.:.:.:.:.~
,,,.,.,.,.,.,.,.
,.,.,
.,.,.,.
:::i:::::::::::::::::::::~::,:,:,~:.~,:,
1
/-
Descending 1
convection I
current
t
/
I
:
FIG. 13. Model showing structure of mantle and subduction ofa cool, thick plate of differentiated oceanic lithosphere.
Previous subduction during the Archaean and also more recent subduction of thickened oceanic crust (Figs. 8b, 14)
have produced a gravitationally stable layer of garnetite overlying the 650 km discontinuity. The tip of a cool, thick
plate experiences buoyant resistance when it penetrates this layer and encounters the seismic discontinuity at 650 km.
Here, the former oceanic crust and harzburgite layers may buckle and plastically thicken to form a large melange
(megalith) situated mainly below the seismic discontinuity. The megalith is a transient feature and ultimately becomes
entrained in the convective regime of the Lower Mantle. The lower layer of the descending plate of sub-oceanic
lithosphere becomes delaminated and resorbed into the Upper Mantle because of its inability to penetrate the gametite
layer (-600-650 km) owing to the buoyancy relationships depicted in Figs. 10 and I I. Partial melting of subducted
oceanic crust between 200 and 600 km causes geochemical refertilization of adjacent depleted peridotite which is
entrained in the flow and becomes concentrated in a thin layer immediately overlying the gametite. This layer of
refertilized peridotite ultimately provides a source region for geochemically enriched intraplate (hot-spot) basaltic
magmas.
transient existence, comparable
with the lifetimes of their
parental subduction systems.
A number of geophysical observations are consistent with
the existence of these structures. DZIEWONSKI and WOODHOUSE (1987) found that the degree of seismic heterogeneity
at depths of -650 km is greater than at shallower (~500
km) and deeper (>800 km) depths. RINCWOOD (1982,1989b)
suggested that megaliths might explain the existence of large
positive gravity anomalies (-50 mgL) which extend for up
to 1000 km behind subduction zones in the direction of the
arc or continent (e.g., Fig. 12 in PHILLIPS and LAMBECK,
1980).
CREAGER and JORDAN (1984, 1986) concluded
that
anomalously
seismic high velocities were present at depths
of 650- 1200 km below subducting plates in the Western Pacific. This has frequently been cited as evidence for direct
penetration
of slabs to great depths in the Lower Mantle.
However, more recent seismic studies reveal a more complex
2102
A. E. Ringwood
picture. FISHER et al. (1988) found that residual sphere
anomalies from deep earthquakes in the Mariana slab were
consistent with thickening of the slab below 650 km by a
factor of 5 or more. This could be explained by the presence
of a megalith. ZHOU and CLAYTON (1990) found that highly
complex structures underlie the intersections
of slabs with
the 650 km discontinuity
and that they vary considerably for
different slabs. The structures appear to include interfmgering,
and sub-horizontal
extensions, as well as thickening of the
slabs. Zhou and Clayton concluded that penetration of slabs
into the Lower Mantle was impeded by a barrier at 650 km.
SHEARER (199 1) investigated
the depths to the 650 km
discontinuity (actually at 660 km) on either side of subducted
slabs in the northwest Pacific. He found that the discontinuity
was depressed by -30 km, compared to its normal depth,
for a distance of - 1000 km in the back-arc direction, and
for a distance of -500 km in the opposite direction. The
depressions could be explained in terms of the spine1 - perovskite + magnesiowiistite
transition, if the mantle in these
regions were a few hundred degrees cooler than average.
Shearer pointed out that the implied temperature distribution
is not predicted by models in which slabs slice cleanly through
the discontinuity
(e.g., CREAGER and JORDAN, 1986) but,
models in which the subducting
slab
instead, “supports
spreads out near 660 km. Such models include the gravitationally trapped megalith hypothesis.”
CHRISTENSEN and YUEN (1984) concluded that slabs could
penetrate directly into the Lower Mantle without major deformation,
despite encountering
substantial buoyant resistance at 650 km caused, for example, by a negative phase
transition gradient and/or a change in chemical composition.
However, their analysis applied only to vertically oriented
slabs. Most slabs encounter the 650 km discontinuities
at
angles which are substantially smaller than 90 degrees. Buoyant stresses exerted on the tips of these inclined slabs seem
much more likely to lead to buckling.
The above scenario for the formation of megaliths has been
based upon the subduction of mature oceanic plates, typically
80-120 km thick and 50-100 Ma old (at the point of subduction). Mature plates of this type possess sufficient thermal
inertia to be much cooler than surrounding
mantle when
they reach 650 km. In contrast, relatively young (<30 Ma)
and thin (~50 km) slabs possess comparatively
low thermal
inertia and may also be subducted more slowly than older,
thicker slabs. Accordingly, their internal temperatures
may
approach thermal equilibrium with surrounding
mantle by
a depth of 650 km. Such slabs become aseismic and ductile
at relatively shallow depths and may behave quite differently
from old, thick slabs when they reach 650 km (RINGWOOD
and IRIFUNE, 1988). Their fate is considered in the following
section.
discussed in Section 3). RINGWOOD (1982) initially opposed
Anderson’s perched eclogite model on the basis of some reconnaissance
experiments
by LIU (1980). However, subsequent, more elaborate experiments
by IRIFUNE and RINGWOOD (1987a) showed that subducted
basalt would indeed
be substantially less dense than pyrolite (RINGWOOD, 1990b;
RINGWOOD and IRIFUNE, 1988) between 650 and 750 km
(Fig. 11) and that gravitational trapping of subducted oceanic
crust on the 650 km discontinuity
deserved serious consideration.
RINGWOOD and IRIFUNE (1988) suggested that the subduction of relatively young and thin oceanic plates which
approached
thermal equilibrium
with surrounding
mantle
may facilitate this trapping process. Density relationships depicted in Fig. 11 show that the tip of the plate would be
buoyant when it encountered
the 650 km discontinuity.
Moreover, it would be ductile because of elevated temperatures and transformational
softening. The plate is therefore
likely to be deflected forward along the 650 km discontinuity,
as shown by KERR and LISTER (1987) and leading to the
trapping of former basaltic crust immediately
overlying the
discontinuity.
Subducted oceanic plateaus with crustal thicknesses of up to 20 km (Fig. 14) would be particularly susceptible to gravitational trapping because of their enhanced
net buoyancy below 650 km.
The most favourable conditions for gravitational trapping
of subducted lithosphere
probably existed during the Archaean. A section through Archaean oceanic lithosphere is
displayed in Fig. 8c. Former dunitic lithologies would have
been 0.12 g/cm3 less dense than pyrolite below 700 km. The
thickened (-20 km) mafic or komatiitic crust would have
been even more buoyant. It seems improbable
that such
buoyant differentiated
lithosphere
would enter the Lower
Mantle. Rather, it would be expected to spread out as density
currents along the 650 km discontinuity
to form a globally
encircling layer (KERR and LISTER, 1987). If the mantle were
200-300°C hotter than today, as suggested earlier, its viscosity
TRANSITION ZONE
Spine1 + Garnet
(c) Gravitational Trapping of Oceanic Crust on the 650
km Discontinuity
ANDERSON (1979, 1980) proposed that subducted basaltic
crust would be less dense than pyrolite immediately
below
650 km and would therefore be trapped during subduction
to form a “perched eclogite layer.” His proposal formed part
of a more complex petrogenetic and structural model which
envisaged a layer of “piclogite” between 250 and 650 km (as
Transformational softening
LOWER MANTLE
Perovskite + Mw
AP Pyrolite minus Garnetite = 0.18 g/cm 3
dp
Pyrolite minus Harzburgite = 0.06 g/cm 3
FIG. 14. Encounter of a thin, thermally equilibrated plate with the
650 km discontinuity, Ieading to buoyant trapping of former oceanic
crust (garnetite) immediately above the discontinuity.
Inaugural Ingerson Lecture
would be 2 or 3 orders of magnitude lower. This would facilitate the disposition of the dunitic and basaltic components
to form a stably stratified structure as shown in Fig. 15, ultimately isolating the convective systems of the Upper and
Lower Mantle (RINGWOOD, 1990b; RINGWOOD and IRIFUNE,
1988). Convection currents in the Upper and Lower Mantles
would erode this differentiated barrier via entrainment at its
upper and lower surfaces, whilst it would be replenished by
capture of continually subducting lithosphere. This dynamic
equilibrium may have led to a steady-state thickness for the
barrier and to convective rehomogenisation of the mantle
regions on either side. The layer of garnetite-facies former
oceanic crust trapped above 650 km may have had a thickness
in the vicinity of 50 km.
(d) The Style of Mantle Convection
The body forces which drive the plate into the mantle arise
from the higher densities of the former oceanic crust and
underlying depleted harzburgite layers, as compared to surrounding mantle. The lower lithosphere of depleted pyrolite
(Fig. 13) is warmer and more ductile (BODINE et al., 1981;
KIRBY, 1980) whilst the density difference between the lower
lithosphere and underlying asthenosphere is relatively small.
RINGWOOD( 1982) suggested that because of the small density
difference and its ductile behaviour, the lowermost zone of
the descending lithosphere may not move coherently with
the upper lithosphere and may gradually become resorbed
into the asthenosphere. Thus, the effective width of the slab
decreases as it descends, and, by 650 km, it consists mainly
of the former basaltic and harzburgite components.
Mid-ocean ridge basalts (MORBs) are believed to be produced by partial melting of pyrolite which has experienced
episodic extraction of small amounts of alkalic magmas,
highly enriched in incompatible elements, over a period exceeding lo9 years (GAST, 1968). Resorbtion and recycling of
the lower oceanic lithosphere into the Upper Mantle during
subduction appears to provide a promising means of generating MORB source regions (RINGWOOD, 1982).
The chemical and isotopic compositions of basalts erupted
ARCHAEAN
MANTLE
km
PYROLITE:
SpineCGarnet
facies
2103
from the mantle provide a wide range of constraints on the
existence and lifetimes of geochemical reservoirs in this region. These topics have been extensively discussed by RINGWOOD ( 1982, 1990b) and KESSON and RINGWOOD (1989) in
the context of the model of mantle dynamics depicted in Fig.
13. Actually, this model was formulated with these constraints
very much in mind. It proposes a hybrid form of convection
in which the Upper and Lower Mantles convect essentially
independently, and are separated by a garnetite boundary
layer within the Transition Zone. Superimposed on this bimodal regime, limited transfer of material from the Upper
Mantle to the Lower Mantle occurs in the form of old, thick
oceanic plates that penetrate the 650 km discontinuity,
forming megaliths which are ultimately entrained into the
Lower Mantle convective system. An equivalent amount of
material must be transferred from Lower Mantle to Upper
Mantle, perhaps as penetrative diapirs overlying rising convection currents in the Lower Mantle.
An increasing array of geochemical evidence appears to
be more compatible with a hybrid convective system of this
type than with more simplistic models of exclusively layered
or exclusively whole-mantle convection. For example, trace
element and isotopic mass balances indicate that the continental crust was derived by differentiation from a mantle
reservoir possessing a substantially larger volume than the
mantle regions above a depth of 650 km, but which was much
smaller than the volume of the entire mantle (e.g., ZINDLER
and HART, 1986; GALER et al., 1989). These studies suggest
that the Upper Mantle possesses a highly depleted geochemical signature, whereas the Lower Mantle is only modestly
depleted. O’NIONS (1987) and GALER et al. (1989) concluded
that highly incompatible elements such as U, Th, and Pb are
efficiently extracted from the Upper Mantle and transferred
to the crust on a comparatively short timescale (-600 Ma).
Replenishment of these elements occurs by injection of less
depleted material from the Lower Mantle into the Upper
Mantle accompanied by rapid convective re-homogenization
of the Upper Mantle. O’NIONSand OXBURGH(1983) inferred
from a comparison of the helium and heat fluxes at midocean ridges that whereas most of the radiogenic heat produced in the Lower Mantle was transferred to the Upper
Mantle, there was some kind of barrier to the transport of
the helium produced by radioactive decay. They suggested
that the imbalance might be explained by the presence of a
barrier near 650 km, which impeded the transport of helium,
but not of heat.
The question of slab-mantle coupling
PYROLITE:
Perovskitite
facies
iOOMg/(Mg+Fe)=88
FIG. 15. Possible petrological structure of the mantle during the
Archaean. Gametite layer represents subducted former (komatiitic)
oceanic crust which has been buoyantly trapped above the 650 km
discontinuity. Dunite (Fo& complementary to the komatiite is
trapped in the layer immediately below 650 km. This stably stratified
structure causes the Upper and Lower Mantle to convect independently.
A sinking slab would entrain adjacent mantle material by
viscous coupling, causing it to follow the path of the slab. In
the model depicted in Fig. 13, it is assumed that the width
of the entrained region is quite narrow, no more than 10-20
km. In the mantle convection models of DAVIESand RICHARDS (1991), the entrained region extends several hundred
kilometers on either side of the slab. In these models, the
slab drives the circulation of a large region of adjacent mantle.
RICHARDS and DAVIES ( 1989) pointed out that under these
circumstances, the integrated negative thermal buoyancy of
the sinking region would completely overwhelm the chemical
buoyancy effects (Fig. 11) associated with the narrow layers
2104
A. E. Ringwood
of former oceanic crust and harzburgite, and that the segregation of former basaltic crust on the 650 km discontinuity
illustrated in Fig. 13 would not be possible.
Their argument is correct if the style of mantle convection
is as they propose. This, however, is a matter of debate. These
authors assume that the rheology of the slab-mantle system
can be interpreted primarily in terms of temperature-dependent Newtonian viscosities. However, several observations
suggest that the actual rheology of the mantle may be considerably more complicated and that deformation might
therefore be highly localised.
HELFRICH et al. (1989) showed that the upper boundary
of the slab to depths of -350 km is defined by a discontinuity
at which seismic waves are reflected and converted. A velocity
contrast of 5-10s through a layer - lo-20 km thick is indicated. HELFRICHet al. ( 1989) concluded that preferred orientation of olivine within this layer was one of the factors
contributing to the velocity anomaly. ANDO et al. (1989) and
OHKURA et al. (1990) identified an analogous seismic discontinuity defined by converted waves at a boundary about
35-40 km below the main Wadati-Benioff zone. This lower
discontinuity coincides with a second planar set of earthquake
epicentres, analogous to the Wadati-Benioff zone. These parallel seismic discontinuities may define the rigid core of the
subducted slab, bounded by narrow zones within which displacements are confined. It is difficult to conceive of these
structures existing in a mantle deforming according to the
model of DAVIESand RICHARDS(1991).
A suite of strongly sheared garnet peridotite xenoliths found
in kimberlite pipes in Africa, Siberia, and the USA provides
direct evidence of mantle rheology. The xenoliths have been
deformed at exceptionally high strain-rates amounting to lofold elongation occurring in minutes or hours (GOETZE, 1975;
MERCIER, 1979). Moreover, deformation occurred at remarkably high temperatures, frequently in the vicinity of
1400-16OO”C, and at pressures of 4-7 GPa. During deformation, these garnet peridotites have been invaded by silicate
melts causing them to become “refertilized.” USSON and
RINGWOOD(1989) concluded that these rocks probably represented actual samples of Wadati-Benioff zones and that the
combined high temperatures and high strain rates had been
caused by episodic shearing associated with Wadati-Benioff
zones. These samples testify to the occurrence of deformation
processes deep in the Upper Mantle which obey flow laws
differing rather profoundly from those assumed by DAVIES
and RICHARDS(199 1).
It appears that some essential element is lacking in many
current treatments of mantle convection based on computer
modelling. The missing element may be the heating caused
by viscous dissipation along the boundaries of the sinking
slab (SCHMELING,1989). This heating would cause a considerable decrease in viscosity together with a parallel increase
in strain-rate, thereby concentrating the heating along thin
boundary layers. Schmeling found that these effects caused
the formation of low-viscosity layers on either side of the
sinking slab, with the result that the motion of the slab became
largely decoupled from the neighbouring mantle. These pro-
5Members of the Intraplate BasalticAssociation erupted in oceanic
regions are often called “Ocean Island Basalts” or “OIBs.”
cesses may produce the seismic discontinuities on either side
of the slab, as discussed above, and also the style of deformation displayed by high-temperature peridotites. Moreover,
they provide mechanisms for generating narrow slabs which
are mechanically decoupled from surrounding mantle, as assumed in Fig. 13.
8. HOT SPOTS, PLUMES, AND MANTLE DYNAMICS
An important class of magma is erupted through the lithosphere, both in oceanic and continental regions, and can be
referred to as the “Intraplate Basaltic Association” (IBA).”
Most IBA suites are characterized by enrichments of incompatible elements, whereas MORBs are depleted in these elements. The IBA are produced by widely varying degrees of
partial melting and include picrites, tholeiites, alkali basalts,
basanites, and nephelinites. Experimental investigations (e.g.,
GREEN and RINGWOOD, 1967) show that MORBs and the
IBA are both produced by partial melting of peridotitic source
regions, and that their depleted/enriched characters are inherited from these source regions. If we compare magmas of
both types formed by similar degrees of partial melting (e.g.,
MORBs and Hawaiian tholeiites), the former are depleted in
highly incompatible elements (e.g., Cs, Ba, U) by factors of
lo-20 as compared to the latter.
It was noted earlier (Section 2d; Table 4) that the formation
and fractionation of magmas in the presence of MgSi03 and
Casio3 perovskites would cause certain characteristic geochemical signatures in the resultant melts. There is no evidence of these signatures in either MORBs or in the IBA. On
the contrary, the geochemical characteristics of these magma
series are consistent with source-region lithologies dominated
by garnet, pyroxene(s), olivine, and (silicate) spinel. We conclude, therefore, that the processes which produce these geochemical characteristics occurred not in the Lower Mantle,
but rather in the Upper Mantle and/or the Transition Zone.
There is an emerging consensus that the petrogenesis of
the IBA is related to the subduction of oceanic crust, which
itself consists predominantly of MORBs (e.g., HOFMANNand
WHITE, 1980, 1982; CHASE, 198 1). The processes which
caused enrichment of incompatible elements in the IBA
source region along with characteristic isotopic signatures are
widely debated. According to a model proposed by RINGWOOD (1982, 1990b) and illustrated in Fig. 13, bodies of
subducted oceanic crust experienced small degrees of partial
melting under hydrous conditions at depths between 200 and
600 km. The resultant melts, enriched in incompatible elements, reacted with the adjacent depleted mantle, thereby
transferring their isotopic and trace element signatures into
these peridotitic protoliths. A significant proportion of this
“refertilized peridotite” may have become trapped immediately above the garnetite boundary layer at 650 km (Fig.
13) where it formed a long-lived (_ 1O9a) reservoir. The subcontinental lithosphere may have provided a second environment in which refertilized peridotite was trapped, as illustrated by RINGWOOD(1990b; Fig. I). This material would
possess the capacity to produce alkaline IBA magmas if subjected to further episodes of small-degree partial melting.
Isotopic compositions of Nd, Hf, Sr, and Pb in the IBA
suggest that long time intervals (e.g., lo’-lo9 a) may have
elapsed between initial formation of their refertilized peri-
2105
Inaugural Ingerson Lecture
dotitic source regions and the subsequent reactivation processes which caused partial melting and actually produced
IBA magmas at the Earth’s surface. However, estimates of
the time intervals are uncertain because of the possibility that
isotopic evolution (particularly
in the Pb and Sr systems)
may have been affected by sediment contamination
and hydrothermal alteration of oceanic crust (e.g., CHAUVEL et al.,
199 1). Despite these chronological uncertainties, the evidence
previously cited that IBA magmas are formed by partial
melting of (refertilized) peridotite firmly indicates a two-stage
petrogenetic process separated by a significant time interval.
According to the model shown in Fig. 13, the garnetite
layer overlying the 650 km discontinuity
provides a thermal
boundary layer, which is overlain in turn by a thin layer of
fertilized former peridotite, representing entrained material
formerly adjacent to subducted oceanic crust. It is envisaged
that the garnetite layer was subsequently
heated by rising
convection
currents from the Lower Mantle, causing the
overlying fertilized peridotite to become buoyant and to ascend as plumes. During ascent, partial melting occurred,
leading to the formation of geochemically
enriched magmas
which erupted from hot spots.
Plumes,fiom
the core-mantle boundary?
According to a currently popular hypothesis, IBA volcanism is caused by plumes which ascend directly from a thermal boundary layer surrounding the core. The enriched geochemical and isotopic characteristics
of hot-spot volcanism
have been attributed to relative concentration
of dense bodies
of former oceanic crust by sinking into and within this
boundary layer (e.g., HOFMANN and WHITE, 1982; DAVIES,
1990; DAVIES and RICHARDS, 199 1).
This scenario encounters certain difficulties:
9 Oceanic crust subducted
l-2 Ga ago would have possessed depleted geochemical signatures. Any viable process of remelting of ancient oceanic crust which had collected near the core-mantle
boundary should provide a
great preponderance
of geochemically depleted magmas.
However, tholeiites formed by similar degrees of partial
melting to MORBs display lo-20-fold
enrichment
of
highly incompatible
elements. DAVIES and RICHARDS
( 199 1) suggested that oceanic crust may contain locally
enriched
“inclusions”
which would be preferentially
melted because of lower solidus temperatures.
This explanation is not viable since the temperatures of ascending
plumes exceed those of surrounding
mantle and would
therefore lead to extensive degrees of partial melting, exceeding those which occur beneath mid-oceanic ridges.
Under these conditions,
the former depleted MORB
component of plumes should be almost totally consumed
during melting, and the compositions
of plume tholeiites
and pi&es
should therefore accurately reflect this depleted character.
ii) It was demonstrated
earlier that the geochemical characteristics of the IBA were produced by multistage petrogenic processes which occurred exclusively in the Upper Mantle and Transition Zone and which seem to have
extended over considerable periods oftime (e.g., IO’-lo9
a). It is difficult to reconcile this history with a model
which postulates an ultimate source for the IBA at the
core-mantle boundary. According to the latter model this
source comprised
a heterogeneous
mixture of former
oceanic crust and peridotite (? sediments) which had
been stored as a closed system for - lo9 years at very
high temperatures
without
experiencing
significant
chemical exchange between its separate components.
It
appears doubtful that the complex geochemical and isotopic characteristics
for the IBA could have been generated from this heterogeneous
source material within the
few million years (GRIFFITHS and CAMPBELL, 1990) that
it took the plumes to ascend from the 650 km discontinuity to the lithosphere.
iii) The most widely distributed IBA members belong to the
alkaline association, which includes alkali basalts, basanites, and nephelinites. These occur extensively in typical hot-spot environments
such as Hawaii, where they
are allegedly derived from plumes claimed to be ascending
from the core-mantle boundary. However, numerous and
widely distributed occurrences
of alkaline rocks which
are petrologically, geochemically,
and isotopically indistinguishable from these major hot-spot occurrences are
erupted through lithosphere above currently active subduction zones, or where subduction has recently ceased
(RINGWOOD, 1982, 1990b). These latter IBA members
are clearly derived from sources situated within the Upper
Mantle. Few would argue in favour of ultimate source
regions near the core-mantle boundary. It does not seem
reasonable
to attribute these petrologically
and geochemically identical alkaline rock types to two fundamentally different categories of source regions.
iv) Recent studies of the melting point of iron at high pressures and of the depression of this melting point by dissolved solutes (e.g., oxygen), as reviewed in Section 4c,
do not support the existence of a strong thermal boundary
layer overlying the core-mantle
boundary, which is believed to be essential for the generation of plumes (e.g.,
STACEY and LOPER, 1983).
9. CONCLUDING
REMARKS
An important objective of the Earth Sciences is to explain
the radial variation of physical properties throughout
the
mantle in terms of the chemical compositions
and mineral
assemblages of its various domains. Major progress towards
achieving this objective has been made in recent years and
the principal results are reviewed in the first part of this paper.
The stage has now been set for progress towards achieving a
second major objective-an
understanding
of the dynamical
behaviour of the mantle, including the convective system
which drives plate motions.
Currently,
this objective is
impeded somewhat by the absence of an adequate understanding of mantle rheology. However, geochemical studies
have been of considerable
importance
in defining the existence of long-lived, chemically distinct reservoirs in the mantle which place powerful empirical constraints on the style
of mantle convection.
The model of mantle dynamics proposed in this paper
represents an attempt to reconcile geochemical evidence on
mantle reservoirs with recent advances in understanding
the
physical and mineralogical
constitution
of the mantle. It
makes a number of assumptions
about possible rheologjcal
behaviour of the mantle which have yet to be tested. A basic
A. E. Ringwood
2106
issue in mantle dynamics is the extent to which mantle convection homogenises pre-existing chemical heterogeneities
(e.g., KELLOGG and TURCOTTE, 1986) or, alternatively, the
extent to which chemical heterogeneities themselves modulate
or control the convective circulation of the mantle. It has
been shown in this and earlier papers (e.g. RINGWOOD, 1982)
that large-scale chemical heterogeneities caused by petrological differentiation and phase charges in the mantle can give
rise to density contrasts at least as large as those which are
believed by many geophysicists to drive mantle convection.
Currently, it does not seem possible to decide, CIpriori,
whether thermally driven convection would overwhelm the
effect of chemical differentiation or vice versa. These two
scenarios, however, have quite different implications with
regard to the formation and maintenance of geochemical reservoirs. There are grounds for optimism that progress in understanding these latter topics may soon provide firm constraints which govern the style of mantle convection.
Acknowledgments-The
author is greatly indebted to Dr. I. Jackson,
Dr. S. E. Kesson, Dr. M. McCulloch, Dr. M. Drury, Dr. S. Cox, and
Dr. B. Kennett who read the manuscript carefully and offered numerous critical and constructive suggestions.
Editorial handling: G. Faure
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