Download Water transportation from the subducting slab into the mantle

Gondwana Research 11 (2007) 148 – 165
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Water transportation from the subducting slab into the mantle transition zone
Shigenori Maruyama a,⁎, Kazuaki Okamoto b
a
Department of Earth and Planetary Sciences, Tokyo Institute of Technology, O-okayama 2-12-1, Meguroku, Tokyo 152, Japan
b
Institute of Earth Sciences, Academia Sinica, P.O. Box 120, Academic Road Sec. 2, Nankang, Taipei, Taiwan, ROC
Received 1 June 2006; accepted 5 June 2006
Available online 14 August 2006
Abstract
Using a recently developed petrogenetic grid for MORB + H2O, we propose a new model for the transportation of water from the subducting
slab into the mantle transition zone. Depending on the geothermal gradient, two contrasting water-transportation mechanisms operate at depth in a
subduction zone. If the geothermal gradient is low, lawsonite carries H2O into mantle depths of 300 km; with further subduction down to the
mantle transition depth (approximately 400 km) lawsonite is no longer stable and thereafter H2O is once migrated upward to the mantle wedge
then again carried down to the transition zone due to the induced convection. At this depth, hydrous β-phase olivine is stable and plays a role as a
huge water reservoir. In contrast, if the geothermal gradient is high, the subducted slab may melt at 700–900 °C at depths shallower than 80 km to
form felsic melt, into which water is dissolved. In this case, H2O cannot be transported into the mantle below 80 km. Between these two endmember mechanisms, two intermediate types are present. In the high-pressure intermediate type, the hydrous phase A plays an important role to
carry water into the mantle transition zone. Water liberated by the lawsonite-consuming continuous reaction moves upward to form hydrous phase
A in the hanging wall, which transports water into deeper mantle. This is due to a unique character of the reaction, because Phase A can become
stable through the hydration reaction of olivine. In the case of low-pressure intermediate type, the presence of a dry mantle wedge below 100 km
acts as a barrier to prevent H2O from entering into deeper mantle.
© 2006 Published by Elsevier B.V. on behalf of International Association for Gondwana Research.
Keywords: MORB + H2O phase diagram; Subduction zone geotherms; Peridotite + H2O phase diagram; Water transportation to deeper mantle
1. Introduction
The role of water in the genesis of magma in subduction zone
environments has been widely recognized (e.g. Wyllie, 1988).
The dehydration has been attributed either to direct melting of the
slab (e.g. Nicolls and Ringwood, 1973; Wyllie and Sekine, 1982;
Brophy and Marsh, 1986) or to the release of fluids to decrease in
the melting temperature of the mantle wedge (e.g. Ringwood,
1975; Delany and Helgeson, 1978; Tatsumi et al., 1986).
Dehydration due to isobaric amphibole breakdown, traditionally
taken as the amphibolite–eclogite transformation reaction, is
believed to provide the H2O necessary to trigger partial melting of
the overlying mantle wedge. However, several high-pressure
hydrous minerals have been documented to be stable at coesite
⁎ Corresponding author. Fax: +81 03 5734 3538.
E-mail address: [email protected] (S. Maruyama).
and diamond isograds in many ultrahigh-pressure metamorphic
terranes (see reviews by e.g., Liou et al., 1994; Chopin and
Sobolev, 1995; Schreyer, 1995). Recent experimental studies
(Schreyer, 1988; Poli, 1993; Poli and Schmidt, 1995; Okamoto
and Maruyama, 1999) show that hydrous phases such as
lawsonite, chloritoid, staurolite, phengite and zoisite–clinozoisite
are passive transporters of considerable amounts of water in
subducting oceanic crust to depths greater than that of amphibole
stability.
It is important to determine the precise dehydration depth in the
subducting oceanic plate and the mechanism of water transportation into the deep mantle, because several recent experiments
suggest that H2O can be stored in several hydrous phases (A, E,
hydrous-β, and -γ) in the mantle transition zone (e.g. Inoue et al.,
1995; Kawamoto et al., 1996; Smyth and Kawamoto, 1997). The
aim of this study is to propose H2O transportation mechanisms
into the mantle transition zone, combining the experimental
results for phase diagrams of MORB + H2O and peridotite + H2O,
1342-937X/$ - see front matter © 2006 Published by Elsevier B.V. on behalf of International Association for Gondwana Research.
doi:10.1016/j.gr.2006.06.001
S. Maruyama, K. Okamoto / Gondwana Research 11 (2007) 148–165
together with an additional data set on the T-distribution on the
slab surface.
2. Phase diagram of MORB + H2O system
Okamoto and Maruyama (1999) defined the high-pressure
stability limit of lawsonite in the MORB + H2O system (Fig. 1).
The lawsonite–eclogite facies is bounded by the high-pressure
stability limit of a lawsonite–involving a continuous reaction
defined by the compositional enlargement of garnet towards a
grossular-rich end member with increasing pressure and/or
temperature. Lawsonite in metabasite is stable up to 10 GPa at
temperatures lower than 900 °C.
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The so-called eclogite facies is subdivided into five facies.
The dry eclogite facies has the largest P–T space and is bounded
by the jadeite eclogite and garnetite facies at P higher than
16 GPa (Irifume et al., 1986; Okamoto and Maruyama, 2004)
and the dry solidus on the high-T side. The low-pressure limit of
lawsonite-bearing eclogite facies is bounded by the zoisite
eclogite subfacies involving a reaction with a gentle positive
slope at 2–3 GPa (Poli and Schmidt, 1995). The low-pressure
and low-temperature boundary is separated by a gentle reaction
curve with a negative Clapeyron slope from the blueschist facies
(Maruyama et al., 1996; Maruyama, 1997). The amphibole–
eclogite facies is stable below the zoisite–eclogite facies, down
to 1.0 GPa. Both the zoisite– and amphibole–eclogite facies are
Fig. 1. A phase diagram of the MORB + H2O system (Modified after Okamoto and Maruyama, 1999, 2004). Thin solid lines delineate polymorphic transformation of
(1) quartz/coesite (Bohlen and Boettcher, 1982), (2) coesite/stishovite (Zhang et al., 1996), and (3) diamond/graphite (Bundy, 1980), and (4) an univariant reaction,
albite = jadeite + quartz (modified from Holland, 1983, see Maruyama et al., 1996). The pale-colored shadow area delineates the stability field of lawsonite eclogite in
the MORB + H2O system (Poli and Schmidt, 1995; Okamoto and Maruyama, 1999). Dark-colored shadow area delineates the slab melting region in the MORB + H2O
system. Facies boundaries below 4 GPa are after Maruyama et al. (1996). Abbreviations; PA = pumpellyite–actinolite facies, GS = greenschist facies, EA = epidote–
amphibolite facies, AM = amphibolite facies, GR = granulite facies, HGR = high-pressure granulite facies, BS = blueschist facies, Am Ec = amphibole eclogite facies, Zo
Ec = zoisite eclogite facies.
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terminated by wet solidus at T higher than 650 °C. The limits of
metamorphic facies below 2.0 GPa are after Maruyama et al.
(1996) and Maruyama (1997).
3. Fate of water from mid-oceanic ridge to subduction zone
3.1. The average H2O content of hydrated oceanic crust
New oceanic crust is born at mid-oceanic ridges, and suffers
hydrothermal fluid-alteration (“ocean-floor metamorphism” of
Miyashiro, 1973). The crust generated at a mid-oceanic ridge is
about 7 km thick, and the top 800 m is hydrated by reacting with
the circulating hydrothermal water (Fig. 2). Beneath the 800 m
depth, the crust is essentially dry except along fracture zones.
The H2O added to the oceanic crust by hydrothermal metamorphism is difficult to estimate precisely, simply because it
depends on the degree of development of hydrothermal fracture
systems at the mid-oceanic ridge, which we cannot directly
observe. The H2O is fixed as loosely bounded water in sheet
silicates or pore space and structural water in the crystalline
phases. The measured structural water from the DSDP hole
504B (Alt et al., 1986) is approximately 1 to 2 wt.% from the
top of the pillow lava down to 600 m. However, it increases to
4–6 wt.% from 600 down to 800 m depth (the transition from
the pillow to dike zone). The structural water of the dredged
basaltic samples by multiple regression analysis also ranges
from 1 to 8 wt.% (Hart, 1976). Staudigel et al. (1995) estimated
the structural water at about 2.7 wt.% and the interstitial water
is at least 2 wt.%, based on the void volume of the crust (approximately 5 vol.% by Johnson, 1979). The present study uses
5–6 wt.% H2O+ which is stored in hydrous minerals due to
ocean-floor hydrothermal metamorphism.
Dredged gabbros and ultramafic rocks from fracture zones
are highly altered and suffered high-grade greenschist to amphibolite facies metamorphism. These rocks contain several
kinds of hydrous minerals. However, such hydrated rocks are
restricted to fracture zones (Fig. 2).
During transport of oceanic crust from the mid-ocean ridge
to the trench, no significant metamorphic and dynamic events
occur between mid-oceanic ridge and subduction zone if OIB
volcanism does not affect the MORB crust (Maruyama, 1991).
Fluid trapped in the pore space of the subducting oceanic crust
and overlying sediments is expelled within the 10–15 km depth
range (e.g. Moore, 1992). Loosely bounded water (H2O−) is
also released at shallow depths between 15 to 20 km. Subsequently, the subducted oceanic crust suffers subduction-zone
Fig. 2. Water circulation system and plate tectonics.
S. Maruyama, K. Okamoto / Gondwana Research 11 (2007) 148–165
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Fig. 3. P–T diagram showing the estimated H2O wt.% of the subducted oceanic crust (modified after Okamoto and Maruyama, 2004). Two representative geotherms
(young vs. old slabs) along the slab surface are shown.
metamorphism and continues to release H2O, by dehydration
reactions, to the overlying mantle wedge (Fig. 2). The change in
H2O content by dehydration reactions due to subduction zone
metamorphism will be reviewed below.
3.2. The maximum H2O contents in the MORB + H2O system
The maximum H2O content in the hydrated subducting
oceanic crust can be calculated, based on the modal variation of
hydrous minerals in metabasalt for given metamorphic facies
(e.g. Peacock, 1993). The mineralogical changes in metabasalt
as a function of P and T are best illustrated using the concept of
metamorphic facies (Eskola, 1920), because most hydrous
minerals in metabasalt break down by a complex series of
continuous reactions. Peacock (1993) calculated the maximum
H2O content in each metamorphic facies, except for the eclogite
facies, in the NCMASH system. He calculated a metamorphic
mineral norm (that is similar to the CIPW norms) in order to
estimate the maximum H2O content in each of the metamorphic
facies. He treated the eclogite facies as completely dry, and fate
of hydrous minerals.
However, to estimate the amount of H2O transported from
the surface to the mantle by a subducting oceanic plate, both the
blueschist–eclogite and the epidote amphibolite–eclogite
transformations are of critical importance. The change in the
amount of H2O due to the blueschist – lawsonite–eclogite
transformation was estimated to be 3 wt.% by Poli and Schmidt
(1995) and that of the amphibolite–eclogite facies transformation to be 1.5 wt.% by Poli (1993). They calculated the mineral
proportions employing a mass balance program which includes
a Monte Carlo error propagation. In this study modal wt.% of
each phase was calculated with the least square method for
major elements. The modal wt.% of lawsonite at 8 GPa, 700 °C
was estimated to 4.0 which corresponds to 0.5 wt.% of water in
the hydrated MORB crust (Okamoto and Maruyama, 1999).
The maximum H2O content of hydrated MORB crust has been
estimated by Peacock (1993), Poli (1993) and Poli and Schmidt
(1995) to be below 6 GPa, and those are compiled in Fig. 3.
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Based on the previous experimental studies, most workers
considered that most H2O in the subducting oceanic crust is
released by the amphibolite – eclogite transformation at 60 km
depth, and less than 1 wt.% H2O is retained beneath the volcanic
front (at ca. 100–120 km depth), if the temperature is below
650 °C (within the stability field of lawsonite–eclogite facies).
However, if the temperature is above 650 °C at 30–90 km
depths, subducting oceanic crust reaches the solidus temperature condition, causing the partial melting of amphibole-bearing
metamorphosed oceanic crust. The released H2O dissolves in
the melt; then the remaining subducting oceanic crust transforms to dry eclogite. If the eclogite becomes anhydrous, it will
not melt due to the high-T solidus over than 1500 °C at about
200 km depth. The H2O content in the subducting oceanic crust
depends on the variation of P–T path along the slab surface.
4. T-distribution on the slab surface
The fate of water in subducting oceanic lithosphere, thus,
strongly depends on the T distribution in the subducting slab. For
an accurate estimate, it is critical to apply a multi-disciplinary
approach, because numerical simulations cannot give a unique
solution, because there are too many independent parameters to
control the thermal structure along the slab surface. In the
following, we first discuss the approach by numerical simulation,
then summarize the field constraints, and finally define the
predominant parameters which control the thermal structure.
4.1. Numerical simulation
Models presented over the past 25 years demonstrate that the
thermal structure of a subduction zone depends on numerous
parameters, including: (1) the thermal structure (age) of the
incoming lithosphere, (2) the convergence rate, (3) the geometry
of subduction, specifically subduction dip angle, (4) the
distribution of radiogenic heat-producing elements, (5) the rate
of shear heating along the subduction shear zone, and (6) the
geometry and vigor of induced convection in the overlying
mantle wedge (Fig. 4) (e.g. Oxburgh and Turcotte, 1970; Hasebe
et al., 1970; Toksöz et al., 1971; Turcotte and Schubert, 1973,
Anderson et al., 1978; Hsui and Toksöz, 1980; Honda and
Uyeda, 1983; Wang and Shi, 1984; Cloos, 1985; Honda, 1985;
van den Beukel and Wortel, 1988; Toksöz and Hsui, 1989;
Peacock, 1990, 1991; Davies and Stevenson, 1992; Staudigel
and King, 1992; Furukawa, 1993a,b; Peacock et al., 1994).
Moreover, although poorly known, (7) the role of fluid
dehydrated from the downgoing slab (Anderson et al., 1976,
1978; Delany and Helgeson, 1978), and (8) the exothermic or
endothermic reaction within it would affect the thermal structure
of the top of the downgoing slab.
Proposed P–T paths for the oceanic crust at the top of the
subducted slab differ sharply in the various models shown in the
above papers. The results of selected numerical models are
summarized by Peacock (1996). Although temperature distributions within the mantle wedge and in the subducted slab are
relatively similar in all of these models, dramatic differences
occur at the interface between the mantle wedge and the
subducted slab. In the absence of shear heating, estimates of the
temperature in the subduction shear zone at 100 km depth
ranges from 300 to 750 °C. The lower calculated temperatures
result from experiments in which the mantle wedge is assumed
to be rigid. Qualitatively, induced convection in the mantle
wedge brings warm mantle material into close proximity to the
subducting slab; induced convection warms the subducting slab
at the expense of cooling of the adjacent mantle wedge. Most
models suggest that induced mantle–wedge convection heats
the top of the slab by several hundred degrees. In the Peacock et
al. (1994) estimation, induced convection increases the
temperatures of the slab surface at 100 km depth from 150 °C
to 450 °C (Fig. 5a).
In the absence of shear heating, faster convergence rates
result in cooler subduction shear zone temperatures (Fig. 5b).
Below 70 km depth and above 100 km depth, temperatures of
the faster subducting slabs are lower than that of the slower
subducting slabs. At 100 km depth, differing convergence rates
do not cause a large temperature difference on the slab surface;
T = 450 °C if the V = 100 mm/yr and T = 550 °C if the
V = 10 mm/yr.
The thermal structure (age) of the oceanic lithosphere prior to
subduction strongly influences the thermal structure of the
subduction zone (Fig. 5c). If the age is 50 Myr, the temperature
of the slab surface is approximately 450 °C at 100 km depth. If
the age is less than 2 Myr, although the convergence rate is
as fast as 100 mm/yr, subducting slab causes partial melting.
Fig. 4. Schematic cross section of subduction zone and the parameters to control the thermal structure of subduction zone (modified from Peacock, 1996).
S. Maruyama, K. Okamoto / Gondwana Research 11 (2007) 148–165
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Fig. 5. Steady-state P–T paths along the subduction shear zone calculated by Peacock et al. (1994). (a) Calculated steady-state P–T conditions based on analytical
expressions (Molnar and England, 1990) and numerical calculations (Peacock et al., 1994). V = 100 mm/yr. Age of subducting lithosphere = 50 Ma. (b) Subduction
shear zone T conditions for the two different convergence rates (10 and 100 mm/yr). Age of incoming lithosphere = 50 Ma. (c) Subduction shear zone P–T conditions
as a function of age of the subducting lithosphere. τ = 0. V = 100 mm/yr. (d) Subduction zone P–T conditions for different shear stresses. Age of incoming
lithosphere = 50 Ma.
Temperature at 100 km depth is 750–1050 °C (Fig. 5c). If the age
of the oceanic plate is older than 50 Myr and shear stress is
negligible, the top of the subducting oceanic plate suffers
blueschist facies metamorphism at depth range from 15 km to
60–70 km, and transforms to lawsonite–eclogite thereafter. The
average age of oceanic plate is approximately 100 Ma on the
modern Earth, and the oldest plate dates back to 200 Ma.
Shear heating, caused by shear stresses of the order of
100 MPa, dramatically increases calculated temperatures in the
subduction shear zone to values in excess of 1000 °C at 100 km
depth (e.g.Toksöz et al., 1971; Turcotte and Schubert, 1973)
(Fig. 5d). Subduction-zone shear stress is poorly known. Recent
estimates of shear stresses in subduction zones range from about
100 MPa (Honda, 1985; Scholz, 1990; Molnar and England,
1990) to several tens of MPa (Bird, 1978; van den Beukel and
Wortel, 1988; Peacock, 1992; Titchelaar and Ruff, 1993) to
approximately zero (Hyndman and Wang, 1993), and is
negligible at depths greater than 50–60 km, where the
subducting crust and the overlying mantle are mechanically
coupled (Furukawa, 1993b). Furthermore, shear heating may be
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absorbed by dehydration reactions taking place in the subducting slab (Anderson et al., 1976, 1977; Delany and Helgeson,
1978). Presence of H2O also lowers rock strengths and frictional
dissipation even lowers shear stress at depth.
Additional variables that lead to significant uncertainties in
calculating the thermal structure of subduction zones include
hydrothermal circulation in the oceanic lithosphere, the variation
of thermal conductivity with T and P, the amount of heat consumed and released by metamorphic reactions, and the effects of
fluid flow at shallow depths. Metamorphic dehydration reactions
in the subducting slab may consume significant amounts of heat
(Anderson et al., 1976, 1978; Delany and Helgeson, 1978).
Similarly, hydration reactions in the overlying mantle wedge
may release significant amounts of heat (Peacock, 1987).
Thermal calculations by Peacock (1990) that assumed 2 wt.%
bound H2O in the subducting oceanic crust indicate that metamorphic reactions have only a minor effect on the calculated
thermal structure (< 5 °C).
Summarizing the numerical simulation, the critical factors to
control the thermal structure of the subducting lithosphere are the
age of the lithosphere, the convergence rate and the induced
mantle convection. Among these factors, the age the lithosphere
has systematic spatial variation in the present- and paleosubduction zones in circum-Pacific region. Shear stress is difficult
Fig. 6. (a) Y or Yb content in the adakites and island arc volcanics as a function of age of subducting lithosphere (Drummond and Defant, 1990). (b) P–T conditions on
the top surface of young slab calculated by Peacock et al., 1994).
S. Maruyama, K. Okamoto / Gondwana Research 11 (2007) 148–165
to estimate, but several studies show that it is less than several tens
of MPa, giving a negligible effect at depth larger than 50 km. Thus
we further argue the P–T conditions in various subduction zones
reconstructed from the age of the lithosphere.
4.2. Field constraints
4.2.1. (1) Adakites from slab melting of MORB
The trace element and REE patterns of adakites and high-Al
Archean TTG suites suggest that they are derived from partial
melting of a garnet-bearing MORB source (Kay, 1978; Martin,
1986). Recently, several researchers have described adakites in
modern-arc environments where < 25 Ma oceanic crust is currently being subducted (Fig. 6a; Drummond and Defant, 1990).
These recent volcanics crop out between the main calc–alkaline
arc and the trench, do not have typical arc geochemical
signatures, and may represent direct partial melts of subducting
oceanic crust. If the source of these volcanics is direct partial
melting of the downgoing slab, then these magmas provide a
powerful constraint on the thermal structure of subduction
zones. The slab surface of the subducting lithosphere younger
than 25 Myr may reach temperatures over 650 °C at 50 km
depth or further; that is, when it reaches the wet solidus in the
high-P amphibolite and granulite facies (Fig. 6b). Applying the
convergence rate and the age of subducted lithosphere beneath
southernmost Chile and Mt. St Helens suites, P–T paths for the
top of the subducted lithosphere were estimated (Peacock et al.,
1994). The estimated P–T paths (condition satisfying both
induced mantle wedge convection- and shear heating-free) also
reach the wet solidus.
4.2.2. (2) T-estimates based on the depth range of the
seismogenic zone of subduction thrust faults
Subduction zone faults generate earthquakes over a limited
depth range. Hyndman et al. (1997) demonstrated that the up
dip (upper limit) and the down dip (lower limit) of the
seismogenic zone are thermally controlled by the slab surface.
The subduction zone faults are aseismic in their seaward updip
portions and landward downdip of a critical point (Fig. 7; see
above). The seaward shallow aseismic zone, commonly
beneath accreted sediments, may be a consequence of
unconsolidated sediments, especially stable-sliding smectite
clays. If the clays are dehydrated and transformed to illite, then
the fault may become seismogenic where the temperature
reaches 100–150 °C (e.g. Hower et al., 1976), at 5–15 km
depth (Fig. 7; see above). The downdip seismogenic limit for
subduction of young, hot oceanic lithosphere is also temperature-controlled. The maximum temperature for seismic
behavior in crustal rocks is 350 °C (brittle–ductile transition).
For subduction beneath thin island arc crust and beneath
continental crust in some areas, the forearc mantle is attached to
the subducting oceanic crust at temperatures higher than the
350 °C. In such a case, the effect of the serpentinization is
avoided. Using the determined seismogenic zone from
Cascadia and Southwest Japan, where very young and hot
plates are subducting, and from Alaska and most of Chile,
where the forearc mantle is reached in the crust, updip and
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downdip temperatures were determined (see Hyndman et al.,
1997; Oleskevich et al., 1999) (Fig. 7, below).
For Cascadia and SW Japan, melting P–T conditions of the
subducting slab causing production of the adakite suites are also
plotted. The P–T path for SW Japan is higher P and lower T than
that for Cascadia. This may be due to the greater age of the
subducting lithosphere of the westward descending Pacific plate.
The P–T paths for Chile (North Chile, Taltal, Coquimbo, and
Valparaiso) and for Southern Alaska are steeper than those of
Cascadia and SW Japan, showing that they are due to the older
the subducting slab, and the steeper the geotherm along the
Benioff thrust. However, the slab descending beneath South
Chile, with the fastest subducting speed (9 cm/yr) of the present
examples, has the steeper gradient in spite of being younger age
(5 Ma) than the Cascadia (6.5–8 Ma) and the SW Japan (0–
15 Ma) analogue. The Valparaiso slab (9 cm/yr) also has the
steeper gradient in spite of being younger (37 Ma) than the North
Chile, Taltal, and Coquimbo (45–50 Ma) and the South Alaska
(50 Ma) plates. This tendency may be partly due to a higher
convergence rate (9 cm/yr vs. 4–6 cm/yr). Generally speaking,
all of these data support the idea that the age of the subducting
slab dominantly controls the thermal structure on the Benioff
plane although high convergent rate is recognized in some
young slabs. Normal subduction-zone geotherm passes through
the lawsonite–eclogite facies.
4.2.3. (3) P–T conditions of HP and UHP metamorphic belts
Fossil geotherms in paleo-subduction zones are wellpreserved in HP and UHP metamorphic belts all over the
world (Maruyama et al., 1996). Using these fossil geotherms,
radiometric and fossil ages, reconstructed paleo-plate geometries and convergence rates, we can estimate the dominant factors controlling the thermal structure on the slab surface in a few
well-constrained metamorphic belts.
The fossil geotherm and age of the subducted oceanic crust
are well reconstructed for the Cretaceous high-P–T Sanbagawa
belt, SW Japan (Maruyama and Seno, 1986; Isozaki and
Maruyama, 1991; Maruyama, 1997). The reconstructed P–T
path is shown in Fig. 8. The metamorphic field gradients have
been determined precisely by Banno and his co-workers for the
last three decades (e.g. Banno, 1964; Higashino, 1975; Banno et
al., 1978; Banno and Sakai, 1989; Higashino, 1990; Otsuki and
Banno, 1990, etc). The facies series ranges from the pumpellyite–actinolite to transitional blueschist–greenschist, greenschist, epidote–amphibolite, and amphibolite facies, and is
classified of the HP intermediate type (Miyashiro, 1961).
Eclogite–facies assemblages occur locally in Fe2+-rich highest-grade metabasites (Banno et al., 1978). Thus, the metamorphic grade ranges from 0.4–0.5 GPa and 250–300 °C for the
lowest grade (Maruyama and Liou, 1983) to 1 GPa and 550–
600 °C for the highest grade (Banno and Sakai, 1989).
The youngest fossils from the blueschist–facies rocks in the
Sanbagawa belt are latest Jurassic to earliest Cretaceous
radiolarians (Iwasaki et al., 1984). Radiometric ages range
from 120 Ma to 60 Ma, depending on the distance from the root
zone and metamorphic grade (Itaya and Takasugi, 1988; compilation by Isozaki and Maruyama, 1991). 40Ar/39Ar dates yield
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Fig. 7. Subduction zone P–T conditions deduced from the depth and temperature of the updip and the downdip seismogenic zone (see text for details). The updip and
downdip temperatures are estimated by the temperatures of smectite–illite transformation and of the brittle–ductile transition of the crustal material, respectively (see
Hyndman et al., 1997; Oleskevich et al., 1999).
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Fig. 8. The fossil geotherm of the high P–T Sanbagawa metamorphic belt (e.g. Banno and Sakai, 1989; Maruyama et al., 1996), and Alpine UHP–HP metamorphic
belt (e.g. Chopin, 1984; Kienast et al., 1991, etc.).
the same results (Takasu and Dallmeyer, 1990; Takasu et al.,
1994; Dallmeyer et al., 1995). Recent U–Pb SHRIMP age of the
eclogite ranges 120–110 Ma as peak metamorphic age
(Okamoto et al., 2004). Thus the ages of formation of the
accretionary complex of the Sanbagawa protoliths and of
metamorphism are very close to each other, at ca. 145–130 Ma
and 120–60 Ma respectively, with a clustered metamorphic age
peak at 80 Ma. The plate geometry from the Jurassic to the
present in the NW Pacific Ocean margin was reconstructed by
Engebretson et al. (1985). According to them, the Kula/Pacific
mid-oceanic ridge subducted under NE Asia at about 80 Ma.
Information recorded in the accretionary fold belt of the
Shimanto belt indicates that the Kula/Pacific mid-oceanic ridge
subducted underneath Japan at ca. 75 Ma (Taira, 1985),
correlating the exhumation of the Sanbagawa high P–T belt
from the 30 km depth, with approach of the mid-oceanic ridge
to the subduction zone (Maruyama, 1997). Convergence rate of
the plate was estimated to be approximately 202 mm/yr
(Engebretson et al., 1985).
Another example is from the Alps (Fig. 8). The fossil
geotherm of the subduction of the Piemont Oceanic crust
beneath the Apulian plate is well-preserved in the HP–UHP
metamorphic belt of the Western Alps (e.g. Chopin, 1984;
Kienast et al., 1991, etc.). The metamorphic facies range
from the pumpellyite–actinolite to blueschist, amphibole–
eclogite facies, and zoisite–eclogite facies (see Maruyama et
al., 1996). The Piemont oceanic plate was born at ca.
180 Ma (Jurassic to Triassic age; Channel et al., 1979),
resulting from the break-up of Pangea and the separation of
Laurasia + Gondwana to initiate the opening of the central
Atlantic Ocean in the Early Jurassic. The metamorphic age
of the UHP rocks was determined as ca. 35 Ma by
Gebauer et al. (1993, 1997) by ion-probe dating (SHRIMP)
of UHP zircons. The age of the subducting Piemont oceanic
plate has been assumed to be 145 Ma. The convergence
rate was estimated to be very low, approximately 0.2 mm/
yr, according to the plate reconstruction of Dewey et al.
(1989).
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In the above contrasting examples (Fig. 8), an old slab with a
very slow convergence rate vs. a young slab with an extremely
fast convergence rate, clearly supports the idea that the age of
the subducting oceanic plate is a critical parameter in determining the P–T condition of the slab surface of the subducting
oceanic plate. Applying this approach to estimate the P–T path
of the slab surface, we can roughly estimate the path in each
respective subduction zone. The P–T path beneath NE Japan is
much colder than that of the Alps, because the age of the
subducting oceanic plate is about 100 Myr with much faster
subduction speed than that of Alps. The convergence rate is
approximately 100 mm/yr, which is 50 times higher than that of
Alps, indicating a path within lawsonite–eclogite facies.
5. Phase diagram of peridotite + water
Hydrated oceanic crust releases H2O to the overlying mantle
wedge under various P–T conditions. The released H2O should
be mainly retained in hydrous phases; some is stored in
anhydrous minerals (several tens ppm in olivine, garnet and
several hundreds ppm in pyroxenes; see Bell and Rossman,
1992), in aqueous fluid phase, and in the melt (Thompson,
1992). The fate of released water is controlled by the stability
relations of hydrous phases in the overlying mantle wedge. Thus
to discuss the transportation mechanism of water into the deeper
mantle, we need a phase diagram for peridotite + H2O (Fig. 9).
Stability relations of hydrous minerals for ultramafic compositions in the mantle wedge has been determined recently by
ultrahigh-pressure experiments (e.g. Yamamoto and Akimoto,
1977; Bose and Ganguly, 1995; Ulmer and Trommsdorff, 1995;
Kawamoto et al., 1996; Kawamoto and Holloway, 1997). Phase
relations of serpentine, chlorite, amphibole and others are shown
schematically in Fig. 9. Amphibole, chlorite and serpentine are
stable at depths shallower than 200 km; amphibole is stable below
3 GPa, 1000 °C, chlorite below 5 GPa, 800 °C and serpentine
(antigorite) at temperatures lower than 500–600 °C below
6.5 GPa. Thus, with increasing pressure up to 6.5 GPa, maximum
temperatures of hydrous mineral stability decreases from 1000 °C
to 500–600 °C. With a further increase in pressure, hydrous
silicates expand their stability fields towards higher temperatures.
The hydrous minerals (e.g. phase A, E and hydrous-β, -γ) are
candidates for H2O containers in the mantle wedge (peridotite +
H2O system) above 6.5 GPa. The high-T maximum stability limit
of those phases shifts from 600 °C at 6.5 GPa to 1500 °C at
13 GPa. The hydrous β-olivine coexists with melt at temperatures
higher than T = 1100 °C and P = 12 GPa (Fig. 9).
Fig. 9. (a) A phase diagram of the MORB + H2O system and subduction zone geotherms. Open area: hydrous minerals are unstable, shaded areas: hydrous minerals are
stable. Broken lines: phengite stability limit in pelitic rocks (D&H by Domanik and Holloway, 1996; S by Schmidt, 1996). (b) A phase diagram of peridotite + H2O
system (Kawamoto et al., 1996, 1997; Bose and Ganguly, 1995; Ulmer and Trommsdorf, 1995). Open area: hydrous minerals are unstable, shaded areas: hydrous
minerals are stable. A, B, C and D; representative P–T paths of the slab surfaces.
S. Maruyama, K. Okamoto / Gondwana Research 11 (2007) 148–165
The water released from the subducting lithosphere is
trapped by the overlying mantle wedge in which hydrous
silicates are stable, if the geotherm is cold enough to stabilize
hydrous phases. However if the geotherm is hot, the released
H2O is not stable in the hydrous phases in the overlying mantle.
Instead, the H2O would move upwards as porous flow, and
cannot be transported into the deeper mantle through dragging
down by the lithospheric counter-flow.
As already discussed above, the P–T path of the slab surface
varies, and is strongly dependent on the age of the subducting
slab, and the convergence rate. Temperature conditions of the
overlying mantle wedge depend also on the T of the slab
surface.
6. Mechanism of water transportation into the mantle
transition zone
Combining the two phase diagrams in Fig. 9 in addition to
Fig. 3, the water transportation mechanism can be estimated as
follows. First, the following four different T path cases along
159
the slab plane (Fig. 9) are discussed: (A) steepest geotherm
generated by old slab; (D) most gentle geotherm produced by
very young slab; and intermediate types (B) and (C).
In the case of subduction of an old oceanic plate along a low
geothermal gradient, the subducting oceanic crust suffers
progressive high P–T metamorphism from blueschist to lawsonite–eclogite facies (Fig. 9). Lawsonite–eclogite facies assemblages are stable over a wide P–T field extending from
2.5 GPa, T < 500 °C, through 8 GPa, T < 800 °C to 9.5 GPa,
<700 °C. That is, lawsonite is stable at depths ranging from
75 km to 300 km. As described above, lawsonite is continuously
dehydrated with increasing temperature in the coesite field, and
with increasing pressure and temperature in the stishovite field.
Subduction of old oceanic plate (Path A) releases 3 wt.% H2O
continuously from 75 km to 100 km depth by the blueschist–
lawsonite–eclogite transformation and retains 1 wt.% H2O in the
subducting slabs. Therefore, the small wedge corner triangle
(trenchward from aseismic front) must be hydrated extensively
with successive subduction of the oceanic plate with time (Fig.
10a). Intermediate-depth intraslab earthquake has been
Fig. 10. Schematic profile showing the transportation mechanism of H2O down to the mantle. Case A, B, C and D correspond to the P–T paths in Fig. 9. See detail in
the text.
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considered to be caused by dehydration embrittlement in
subducting oceanic crust (e.g. Kirby et al., 1996). The aseismic
front (AF) in NE Japan may correspond to the pressure-sensitive
facies boundary between the blueschist and lawsonite–eclogite
facies because approximately 5 wt.% of H2O would be released
due to continuous dehydration reaction from the blueschist
facies to the boundary. The small triangle mantle wedge may not
join the dragged counter-flow of the mantle because of the
buoyancy of strongly serpentinized peridotite (Fig. 10a).
At depth in the lawsonite–eclogite facies, continuous dehydration of lawsonite releases small quantities of H2O; the
subducting oceanic crust retains about 0.5 wt.% H2O even at
depths greater than 200 km. In the stishovite field, the
descending slab continuously releases all the remaining H2O
until the lawsonite–eclogite stability limit, at 300 km depth. The
H2O released from the subducting oceanic crust, would migrate
upward to hydrate the mantle wedge above the subduction zone.
This H2O is stored mainly in chlorite and serpentine in the
hanging wall peridotite just above the MORB crust down to
200 km depth. The peridotite containing those hydrous phases is
dragged further down and is heated up gradually to release
water-rich fluids. The released H2O moves upward to decrease
the melting temperature of the mantle wedge (Fig. 10a). Islandarc tholeiite magma may be formed by this process (e.g. Tatsumi
et al., 1986). In the mantle wedge adjacent to the slab, however,
serpentine is stable above 200 km depth; below this depth, phase
A is stable from a depth of 200 km to at least a depth of 300 km
or further down to below 300 km depth. These two phases
retain H2O by slab dehydration through the lawsoniteconsuming reaction and the induced mantle wedge convection
may move all the way down to the mantle transition depth.
Below 250 km, phase E is stable at T > 700 °C; the hydrous layer
in the mantle wedge thickens with continuous subduction beneath the mantle wedge. Hydrous phase β-olivine is stable at
depths below 350 km, T > 1100 °C (Fig. 9b). Thus, the surface
water trapped in the oceanic lithosphere can be partly transported
into the mantle transition zone at 410–660 km, if the geotherm is
cold enough to stabilize lawsonite down to a depth of 300 km.
The water containers in the mantle transition zone are phase E,
hydrous β- and γ-olivine. The mantle transition zone at 410–
660 km depth is an enormous water reservoir, as demonstrated
by the 3 wt.% of H2O content of the most dominant phase of βolivine (> 50%).
In the case of Path B (shown in Fig. 9), lawsonite is stable
down to 270 km depth in the subducting oceanic crust. In the
overlying mantle wedge, chlorite and serpentine are not stable
below 150 km and no hydrous phase is stable at depths between
150 km (T > 600 °C) and 240 km (T > 700 °C). Phase A should be
stable below 240 km depth if the H2O is present in the mantle
wedge. However, no hydrous silicates are stable between
150 km and 240 km in the convected mantle wedge. Nevertheless lawsonite is stable down to 270 km, so less than 0.5 wt.%
H2O can be supplied from the slab down to 270 km depth by
continuous dehydration of the remaining lawsonite. Concurrently, the released H2O from the downgoing slab moves
upwards to hydrate the mantle wedge to form hydrous phase A
below 240 km to 270 km (Fig. 10b). Once retained in phase A,
H2O can be carried to greater depths due to the induced
convection of the mantle wedge. Below 270 km, phase E
becomes the main H2O container, and is stable in the mantle
transition zone. At greater depth, phase E is not stable, and H2O
is stored in hydrous β-phase of olivine containing up to 3 wt.%
H2O (Fig. 10b).
From the 150 to 240 km depth, the H2O released from the
subducting oceanic crust by lawsonite-consuming reaction,
migrates upward through the overlying mantle wedge if the
fluid is dominated by H2O. The ascending aqueous fluid causes
partial melting of the mantle wedge, and produces island-arc
magma. However, the fluid dissolves considerable amounts of
Si, Mg and other elements under ultrahigh-pressure conditions
(Fujii et al., 1997). The addition of aqueous fluid may produce
hydrous phases such as K-richterite and Ti-clinohumite in the
mantle wedge.
A moderately old subducting oceanic plate may follow path
C of Fig. 9. The P–T trajectory passes from the blueschist to
zoisite–eclogite facies at 70 km depth. The transformation from
zoisite–eclogite to dry eclogite occurs at 110 km, within the
coesite stability field (Fig. 9a). Most of the H2O is leached out at
110 km depth when the slab is heated up to 650–700 °C. H2O
content of the subducting slab decreases from 0.5 wt.% at 75 km
depth to nearly 0 at 110 km depth. Released H2O could be
stored in serpentine and chlorite of the overlying mantle wedge
along the Benioff thrust. However, no hydrous mineral is stable
at depths below 130 km in the mantle wedge right above the
slab surface. Stored H2O in these hydrous minerals at shallow
levels is migrated upward in the convected mantle wedge at
130 km. The ascending aqueous fluids cause the partial melting
of the highest temperature region of the mantle wedge, to
produce island-arc magma (Fig. 10c).
In the case of subduction of young oceanic crust (path D in
Fig. 9), slab melting occurs at depths 30–70 km and the released
H2O is stored in the melt ascending to the mantle wedge.
Extensive melting of hydrated MORB crust forms calc–alkaline
magma, in which all H2O is incorporated. Hence, H2O cannot
be transported into the mantle wedge below 70 km depth. Once
H2O is leached out of the subducting oceanic crust, it is difficult
to produce melt again, because the melting temperature defined
by the dry solidus of the MORB is higher than 1220 °C at 70 km
depth. The hydrated SiO2-rich felsic melt would react with αolivine during the magma ascent into the overlying mantle, to
form orthopyroxene as an interface curtain wall in the melt
pathway (Fig. 10d; Maruyama, 1997).
7. Role of serpentinized peridotites along transform faults
on the oceanic floor
Oxygen isotope analyses of ophiolites (e.g. Gregory and
Taylor, 1981) and direct drilling of oceanic crust at DSDP site
504B (Alt et al., 1986) indicate that hydrothermal circulation at
mid-oceanic ridges is shallower than 800 m based on boiling
study of the seawater detected at 504B, or at least less than
5 km, estimated by water/rock ratio by oxygen isotope on
ophiolite. Maede and Jeanloz (1991) estimated the maximum
thickness of the hydrated zone to be 12 km, which is far thinner
S. Maruyama, K. Okamoto / Gondwana Research 11 (2007) 148–165
compared to the thickness of slab 600–100 km if it is older than
50 Myr. If this is true, major parts of oceanic peridotite under the
MORB crust are predominantly dry, and are unrelated to the
transportation of surface water into the deeper mantle.
However, it is well-known that serpentinized peridotites are
exposed along transform faults on the ocean-floor (Fig. 2).
Therefore, some parts of oceanic peridotite are undoubtedly
hydrated and may transport surface water into the mantle transition zone, although it is difficult to estimate its role
quantitatively. If the lower plane of the double seismic zone is
representing the dehydration of the hydrated peridotite, the
hydrous peridotite in the slab is not only the part hydrated at the
transform faults. Further speculative hypothesis has been
proposed by Seno and Yamanaka (1996), therefore we do not
argue further here. Thus, the presence of those hydrated peridotites promotes the water transportation into deeper mantle.
8. Role of phengite in subducting pelitic rocks
Capping the top of the oceanic lithosphere, the hemipelagic
sediments and trench turbidites may subduct together with
oceanic crust in deeper parts of a subduction zone, although
most would be offloaded to form an accretionary complex at
shallow depths on the hanging wall of the overlying plate. It is
not certain whether they could subduct below Moho depth into
the deep upper mantle, lie above the transition zone. The
stability relations of hydrous phases in a pelitic system have
been investigated on a wide P–T range (Schmidt, 1996;
Domanik and Holloway, 1996). Phengite–muscovite as a
most common hydrous phase in pelitic rocks has a maximum
P-stability up to 8–10 GPa and up to 1060 °C. Other accessory
hydrous phases are topaz–OH, Mg–pumpellyite, lawsonite, and
K–cymrite (Domanik and Holloway, 1996). Phengitic mica has
a stability field 100–200 °C higher than the maximum stability
limit of lawsonite in metabasite at P = 3–10 GPa. Therefore,
some amounts of water may be transported by pelitic rocks
down to 300 km depth only if the geothermal gradients are low
enough including A, B and C in Fig. 9. Thus, pelitic rocks
could, possibly help the transportation of H2O to great mantle
depth.
9. Discussion
Before the advent of plate tectonics in Earth Science, most
earth scientists never discussed the fate of surface water in terms
of global material circulation on a whole mantle scale, because
of the lack of critical petrologic observations, experimental data
sets, and geophysical data on mantle circulation. However, the
presence of phlogopite and K–amphibole in mantle xenoliths
clearly indicate that water is present as structural OH− in mantle
minerals. When plate tectonics appeared, to interpret subduction
zone magmatism, most volcanologists inferred that calc–
alkaline magmas are derived from slab-melting of hydrated
MORB or partial melting of the mantle wedge, to which the
released seawater is added to lower the melting T. Some
petrologists considered that almost all dehydration occurs to
generates arc magma at depths shallower than 150–200 km
161
(Ernst, 1999), hence no H2O penetrates into the mantle transition
zone with time. Therefore, the Earth's mantle transition zone
(410–660 km), if hydrated when the magma ocean had
consolidated during Early Earth at 4.55 Ga, must have gradually
dehydrated by plume activity through geologic time (Kawamoto
et al., 1996). In contrast, there is an opposing idea that
subduction zones deliver six times more water into the mantle
(8.7 × 1011 kg yr− 1) than is delivered to the surface by arc
volcanism (1.4 × 1011 kg yr− 1) (Peacock, 1990), implying the
on-going water transportation into the mantle transition zone. If
this is true, most subducted water is being swallowed into the
mantle transition zone or even into the lower mantle at present,
implying that the total amount of seawater in the oceans is now
being reduced. From the standpoint of mantle mineralogy and
petrology, dense hydrous magnesium silicates (DHMS) which
are stable under mantle conditions, count more than 28 species
(e.g. review by Thompson, 1992). For example, for upper mantle
conditions, after the pioneering work by Ringwood and Major
(1967), Yamamoto and Akimoto (1977) investigated the system
MgO–SiO2–H2O at P = 2.9–7.7 GPa and T = 470–1225 °C, and
found DHMS such as brucite, phase A, phase D, clinohumite,
serpentine, talc, 10 Å phase, and chondrodite. Since then,
superhydrous B (Pacalo and Parise, 1992; Burnley and
Navrotsky, 1996), phase A (Bose and Ganguly, 1995; Pawley
and Wood, 1996), and phase E (Luth, 1995) have been
synthesized and analyzed. Akaogi and Akimoto (1980), Liu
(1986, 1987), Kanzaki (1991), and Gasparik (1993) have
investigated the phase relations near and in the mantle transition
zone. Under lower mantle conditions, phase D (Li and Jeanloz,
1991; Yang et al., 1997; Shieh et al., 1998), phase F (Ohtani et
al., 1995), and phase G (Ohtani et al., 1997; Kudoh et al., 1997)
have been described as stable hydrous silicates if we presume the
presence of water in the lower mantle.
Thus, a variety of hydrous silicates has been demonstrated as
possible stable phases within the whole depth range of upper
mantle and topmost parts of the lower mantle.
In spite of wide and high-T stability fields of those DHMS in
the deeper mantle, surface water cannot be simply transported into the deeper mantle, because (1) the subduction zone
geotherm that they assumed is too warm, resulting in
dehydration of all DHMS at relatively shallow depth ranges
(see Fig. 9b), (2) subducting oceanic lithosphere is capped by
7 km thick MORB, and MORB cannot carry water into mantle
deeper than 300 km. Although the core is the coldest part of the
subducting slab, hydrous parts of the slab are less than the
topmost 12 km (Maede and Jeanloz, 1991), or even less
(probably less than 800 m as discussed before), and (3) water
transportation into the mantle transition zone depends strongly
on the subduction-zone geotherm as explained below. If it is
warm, no water can enter into the deeper mantle. If the presence
of limited lateral arrangement of arc volcanoes usually 120–
200 km right above the slab surface, means that all DHMS
dehydrate below 200 km depth, then a suggestion arises that no
water enters into mantle depths greater than 200 km (Kawamoto
et al., 1996).
To solve the debate, phase relations of MORB + H2O and
peridotite + H2O under the whole upper mantle conditions, are
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S. Maruyama, K. Okamoto / Gondwana Research 11 (2007) 148–165
crucial. The summarized phase relations in Fig. 9b provide a
concept regarding the fate of water. Two critical points at X1
(6.5 GPa, 550 °C) and X2 (5.5 GPa, 600 °C), are marked on Fig.
9b. X1 corresponds to the lowest T at which dry mineral
assemblages are stable in mantle peridotite. From this point
toward higher-P and -T, the stability field of anhydrous mineral
assemblages is spread over a wide range. Therefore, if the
Benioff geotherm passes the lower-T side than X1, dragged
downflow of the mantle wedge can carry the water within
hydrous phases into the deeper mantle. If the Benioff geotherm
passes on the higher-T side of X2, water cannot be transported
into deeper mantle not only by downflow of the wedge mantle
but also by hydrated MORB crust. If the Benioff geotherm
passes between X1 and X2, lawsonite can transport the
crystalline water down to 240 km depth (point X3) to be
relayed to the overlying downflow of mantle in which phase A
is stable. Note a unique character of the reaction to form phase
A connecting X1 with X3, because it is a hydration reaction
forsterite + water = phase A (Pawley and Wood, 1996). If no
water is supplied by the downgoing oceanic crust, the mantle
wedge cannot be hydrated. In the case of the Benioff geotherm
passing on the higher-T side of X2, water cannot be transported
into the deeper mantle at all. All hydrous phases in downgoing
MORB crust and in serpentinized peridotites along transform
faults would be dehydrated by 200 km depth, thereafter no
water moves into the transition zone (Fig. 9b). Thus a small
triangular area defined by X1, X2, X3 is critical to control the
fate of water in a subduction zone. The significance of the role
of triangular zone in the P–T space is portrayed in Fig. 10b and
c where, the presence of a dry hanging wall peridotite prevents
dehydrated water moving down as hydrous phases along the
slab surface. Instead, free-water moves upwards to trigger the
partial melting of peridotite.
Thus, the geothermal gradients along the slab surface
determine whether or not the surface water can be transported
into the deeper mantle. The variable P–T gradients along the slab
surface are a key to understanding the fate of water in subduction
zones. If subduction zones do not follow P–T paths like A or B,
then we must support the interpretation by Kawamoto et al.
(1996). But, as discussed before, if subduction zone geotherms
are very variable and depend highly on the age of the slab, and
follow paths such as A or B, we must support the idea by Peacock
(1990). If the slab is older than 50 Myr, it would pass into the
lawsonite stability fields or even much colder P–T ranges, as
well-demonstrated by (1) numerical simulation (Peacock, 1990,
1996), (2) lawsonite–eclogites (reviewed by Okamoto and
Maruyama, 1999), (3) a few young regional metamorphic belts
with independent information on age of subducted slab and
relative plate motion, and (4) P–T estimates by frequency of
active seismicity in several active subduction zones discussed in
this paper. Considering the average age of subducting slabs (ca.
100 Ma) on the modern Earth, we must conclude that surface
water moves down into the deeper mantle in most present
subduction zones. This leads to an idea that surface water is now
decreasing.
Additional supporting observations are derived from seismological aspects of the mantle. Earthquakes deeper than
100 km, down to 650 km in the Earth may easily be explained if
dehydration is on-going at those depths by subduction (Maede
and Jeanloz, 1991; Stein, 1995). At those depths rocks are
expected to deform by ductile flow rather than brittle fracturing
or frictional sliding on fault surfaces (Maede and Jeanloz,
1991).
Acknowledgements
We express sincere thanks to Dr. Chris Parkinson, Profs. E.
Takahashi, J.G. Liou and M. W. Schmidt who gave us useful
comments and suggestions. We appreciate detailed reviews by
Prof. W.G. Ernst.
References
Akaogi, M., Akimoto, S., 1980. High-pressure stability of a dense hydrous
magnesian silicate Mg23Si8O42H6 and some geophysical implications.
Journal of Geophysical Research 85, 6944–6948.
Alt, J.C., Honnorez, J., Laverne, C., Emmermann, R., 1986. Hydrothermal
alteration of a 1 km section through the upper oceanic crust, deep sea drilling
project Hole 504 B: mineralogy, chemistry, and evolution of seawater–basalt
interactions. Journal of Geophysical Research 91, 10,309–10,335.
Anderson, R.N., Uyeda, S., Miyashiro, A., 1976. Geophysical and geochemical
constraints at converging plate boundaries: Part I. Dehydration in the
downgoing slab. Geophysical Journal of the Royal Astrological Society 44,
333–357.
Anderson, R.N., DeLong, S.E., Schwarz, W.M., 1978. Thermal model for
subduction with dehydration in the downgoing slab. Journal of Geology 86,
731–739.
Banno, S., 1964. Petrologic studies of the Sanbagawa crystalline schists in the
Bessi–Ino district, central Shikoku, Japan. Journal of Faculty of Sciences,
University of Tokyo, Sec II 15, 203–319.
Banno, S., Sakai, C., 1989. Geology and metamorphic evolution of the
Sanbagawa metamorphic belt, Japan. In: Daly, J.S., Cliff, R.A., Yardley, B.
W.D. (Eds.), Evolution of Metamorphic Belts. . Special Publication, vol. 43.
Geological Society of London, pp. 519–532.
Banno, S., Higashino, T., Otsuki, M., Itaya, T., Nakajima, T., 1978. Thermal
structure of the Sanbagawa belt in central Shikoku. Journal of Physics of the
Earth 345–356 (Supplement).
Bell, D.R., Rossman, G.R., 1992. Water in Earth's mantle: the role of nominally
anhydrous minerals. Science 255, 1391–1397.
Bird, P., 1978. Stress and temperatures in subduction shear zones: Tonga and
Mariana. Journal of Royal Astrological Society 55, 411–434.
Bohlen, S.R., Boettcher, A.L., 1982. The quartz–coesite transformation: a
precise determination and the effects of other components. Journal of
Geophysical Research 87, 7073–7078.
Bose, K., Ganguly, J., 1995. Experimental and theoretical studies of the stabilities
of talc, antigorite and phase A at high pressures with applications to
subduction processes. Earth and Planetary Sciences Letters 136, 109–121.
Brophy, J.G., Marsh, B.D., 1986. On the origin of high-alumina arc basalt and
the mechanics of melt extraction. Journal of Petrology 27, 763–789.
Bundy, F.R., 1980. The P,T phase reaction diagrams for elemental carbon.
Journal of Geophysical Research 85, 6930–6936.
Burnley, P.C., Navrotsky, A., 1996. Synthesis of high-pressure hydrous magnesium
silicates: observations and analysis. American Mineralogist 81, 317–326.
Channel, J.F.T., D'Argenio, B., Horvath, F., 1979. Adria, the African
promontary. Mesozoic Mediterranean Paleogeography. Earth Science
Review, vol. 15, pp. 213–292.
Chopin, C., 1984. Coesite and pure pyrope in high-grade blueschists of the
western Alps: a first record and some consequences. Contributions to
Mineralogy and Petrology 86, 107–118.
Chopin, C., Sobolev, N.V., 1995. Principal mineralogic indicators of UHP in
crustal rocks. In: Coleman, R.G., Wang, X. (Eds.), Ultrahigh Pressure
Metamorphism. Cambridge University Press, pp. 96–131.
S. Maruyama, K. Okamoto / Gondwana Research 11 (2007) 148–165
Cloos, M., 1985. Thermal evolution of convergent plate margins: thermal
modeling and re-evaluation of isotopic Ar-ages for blueschists in the
Franciscan complex of California. Tectonics 4, 421–433.
Dallmeyer, R.D., Takasu, A., Yamaguchi, K., 1995. Mesozoic tectonothermal
development of the Sanbagawa, Mikabu and Chichibu belts, southwest
Japan: evidence from 40Ar/39Ar whole-rock phyllite ages. Journal of
Metamorphic Geology 13, 271–286.
Davies, J.H., Stevenson, D.J., 1992. Physical model of source region of
subduction zone volcanics. Journal of Geophysical Research 97, 2037–2070.
Delany, J.M., Helgeson, H.C., 1978. Calculation of the thermodynamic
consequences of dehydration in subducting oceanic crust to 100 kb and
>800 °C. American Journal of Sciences 278, 638–686.
Dewey, J.F., Helman, M.L., Turco, E., Hutton, D.H., Knott, S.D., 1989. Kinematics
of the western Mediterranean. In: Coward, M.P., Dietricu, D.D., Park, R.G.
(Eds.), Alpine Tectonics. . Special Publication, vol. 45. Geological Society of
London, pp. 265–283.
Domanik, K.J., Holloway, J.R., 1996. The stability and composition of phengitic
muscovite and associated phases from 5.5 to 11 GPa: implications for deeply
subducted sediments. Geochimica Cosmochimica Acta 60, 4133–4150.
Drummond, M.J., Defant, M.S., 1990. A model for trondjemite–tonalite–dacite
genesis and crustal growth via slab melting: Archean to modern
comparisons. Journal of Geophysical Research 95, 21503–21521.
Engebretson, D.C., Cox, A., Gordon, R.G., 1985. Relative plate motion between
oceanic and continental plates in the Pacific basin. Special Paper, vol. 206.
Geological Society of America, pp. 1–59.
Ernst, W.G., 1999. Hornblende, the continent marker – evolution of H2O during
circum-Pacific subduction versus continental collision. Geology 27, 675–678.
Eskola, P., 1920. The mineral facies of rocks. Norsk Geologisk Tidsskrift 6,
143–194.
Fujii, T., Mibe, K., Yasuda, A., 1997. Aqueous fluid at high pressure: another
silicate magma. Kazan 42, S321–S326 (in Japanese).
Furukawa, Y., 1993a. Magmatic processes under arcs and formation of the
volcanic front. Journal of Geophysical Research 98, 8309–8319.
Furukawa, Y., 1993b. Depth of the decoupling plate interface and thermal
structure under arcs. Journal of Geophysical Research 98, 20,005–20,013.
Gasparik, T., 1993. The role of volatiles in the transition zone. Journal of
Geophysical Research 98, 4287–4299.
Gebauer, D., Tilton, G.R., Shertl, H.-P., Schreyer, W., 1993. Eocene/Oligocene
ultrahigh-pressure (UHP)-metamorphism in the Dora-Maira massif (western
Alps) and its geodynamic implications. Fourth International Eclogite
Conference (Calabria, Italy). Terra Abstract, vol. 4, pp. 10–11.
Gebauer, D., Schertl, H.-P., Brix, M., Schreyer, W., 1997. 35 Ma old ultrahighpressure metamorphism and evidence for very rapid exhumation in the Dora
Maira massif, Western Alps. Lithos 41, 5–24.
Gregory, R.T., Taylor Jr., H.P., 1981. An oxygen isotope profile in a section of
Cretaceous oceanic crust, Samail ophiolite, Oman: evidence for δ18O
buffering of the oceans by deep (>5 km) seawater-hydrothermal circulation
at mid-oceanic ridges. Journal of Geophysical Research 86, 2737–2755.
Hart, R.A., 1976. Progressive alteration of the oceanic crust. In: Yeats, R.S., et al.
(Eds.) Initial Reports of the DSDP, vol. 34, pp. 433–437.
Hasebe, K., Fujii, N., Uyeda, S., 1970. Thermal processes under island arcs.
Tectonophysics 10, 335–355.
Higashino, T., 1975. Biotite zone of the Sanbagawa metamorphic terrain in the
Shiragayama area, central Shikoku. Journal of Metamorphic Geology 8,
413–423.
Higashino, T., 1990. The higher grade metamorphic zonation of the Sanbagawa
metamorphic belt in central Shikoku, Japan. Journal of Metamorphic
Geology 8, 413–423.
Holland, T.J.B., 1983. The experimental determination of activities in
disordered and short range ordered jadeitic pyroxenes. Contributions to
Petrology and Mineralogy 82, 214–220.
Honda, S., 1985. Thermal structure beneath Tohoku, northeast Japan – a case
study for understanding the detailed thermal structure of the subduction
zone. Tectonophysics 112, 69–102.
Honda, S., Uyeda, S., 1983. Thermal process in subduction zones – a review and
preliminary approach on the origin of arc volcanism. In: Shimozuru, D.,
Yokoyama, I. (Eds.), Arc Volcanism: Physics and Tectonics. TERRAPUB,
Tokyo, pp. 117–140.
163
Hower, J., Eslinger, W.V., Hower, M., Perry, E.A., 1976. Mechanism of burial
metamorphism of argillaceous sediments: I. Mineralogical and chemical
evidence. Geological Society of America Bulletin 87, 725–737.
Hsui, A.T., Toksöz, M.N., 1980. The evolution of thermal structures beneath a
subduction zones, in oceanic ridges and arcs. In: Toksöz, M.N., Uyeda, S.,
Francheteau, J. (Eds.), Geodynamic Processes. Development of Geotectonics, vol. 14. Elsevier, New York, pp. 325–342.
Hyndman, R.D., Wang, K., 1993. Thermal constraints on the zone of major
thrust earthquake failure: the Cascadia subduction zone. Journal of
Geophysical Research 98, 2039–2060.
Hyndman, R.D., Yamano, M., Oleskevich, D.A., 1997. The seismogenic zone of
subduction thrust faults. The Island Arc 6, 244–260.
Inoue, T., Yurimoto, H., Kudoh, Y., 1995. Hydrous modified spinel, Mg1.75SiH0.5O4: a new water reservoir in the mantle transition region. Geophysical
Research Letters 22, 117–120.
Irifume, T., Sekine, T., Ringwood, A.E., 1986. The eclogite–garnet transformation at high pressure and some geophysical implications. Earth and
Planetary Sciences Letters 77, 245–256.
Isozaki, Y., Maruyama, S., 1991. Studies on orogeny based on plate tectonics in
Japan and new geotectonic division of the Japanese Islands. Journal of
Geography (Tokyo) 100, 697–761.
Itaya, T., Takasugi, H., 1988. Muscovite K–Ar ages of the Sanbagawa schists,
Japan and argon depletion during cooling and deformation. Contributions to
Mineralogy and Petrology 100, 281–290.
Iwasaki, M., Ichikawa, K., Yao, A., Faure, M., 1984. Fossil age of greenstone
conglomerate in eastern Shikoku. Report of Kansai Branch, vol. 97.
Geological Society of Japan, p. 21.
Johnson, D.M., 1979. Crack distribution in the upper oceanic crust and its
effects upon seismic velocity, seismic structure, formation permeability, and
fluid circulation. In: Donnelly, T., et al.(Eds.) Initial Reports of the Deep Sea
Drilling Project, Washington, DC, pp. 1473–1490.
Kanzaki, M., 1991. Stability of hydrous magnesium silicates in the mantle
transition zone. Physics of the Earth and Planetary Interiors 66, 307–312.
Kawamoto, T., Holloway, J.R., 1997. Melting temperature and partial melt
chemistry of H2O-saturated mantle peridotite to 11 Gigapascals. Science
276, 240–243.
Kawamoto, T., Hervig, R.L., Holloway, J.R., 1996. Experimental evidence for a
hydrous transition zone in the early Earth's mantle. Earth and Planetary
Sciences Letters 142, 587–592.
Kay, R.W., 1978. Aleutian magnesian andesites: melts from subducted Pacific
Ocean crust. Journal of Volcanology and Geothermal Research 4, 117–132.
Kienast, J.R., Lombardo, B., Biino, G., Pinardon, J.L., 1991. Petrology of veryhigh-pressure eclogitic rocks from the Brossasco Isasca Complex, Dora
Maira Massif, Italian Western Alps. Journal of Metamorphic Geology 9,
19–34.
Kirby, S., Engdahl, E.R., Denlinger, R., 1996. Intermediate-depth intraslab
earthquakes and arc volcanism as physical expressions of crustal and
uppermost mantle metamorphism in subducting slabs. In: Bebout, G.E., et al.
(Eds.) Subduction: Top to Bottom. . Geophysical Monograph, vol. 96. AGU,
pp. 195–214.
Kudoh, Y., Nagase, T., Mizohata, H., Ohtani, E., Sasaki, S., Tanaka, M., 1997.
Structure and crystal chemistry of phase G, a new hydrous magnesium
silicate synthesized at 22 GPa and 1050 °C. Geophysical Research Letters
24, 1051–1054.
Li, X., Jeanloz, R., 1991. Phase and electrical conductivity of a hydrous silicate
assemblage at lower-mantle condition. Nature 350, 332–334.
Liou, J.G., Zhang, R., Ernst, W.G., 1994. An introduction to ultrahigh-pressure
metamorphism. The Island Arc 3, 1–24.
Liu, L., 1986. Phase transformations in serpentine at high pressures and
temperatures and implications for subducting lithosphere. Physics of the
Earth and Planetary Interiors 42, 255–262.
Liu, L., 1987. Effects of H2O on the phase behavior of the forsterite–enstatite
system at high pressures and temperatures and implications for the Earth.
Physics of the Earth and Planetary Interiors 49, 142–167.
Luth, R.W., 1995. Is phase a relevant to the Earth's mantle? Geochimica
Cosmochimica Acta 59, 679–682.
Maede, C., Jeanloz, R., 1991. Deep-focus earthquakes and recycling of water
into the Earth's mantle. Science 252, 68–72.
164
S. Maruyama, K. Okamoto / Gondwana Research 11 (2007) 148–165
Martin, H., 1986. Effect of steeper Archean geothermal gradient on
geochemistry of subduction-zone magmas. Geology 14, 753–756.
Maruyama, S., 1991. Ocean-floor metamorphism: a review on the DSDP, IPOD
and ODP. EOS vol. 72, 552.
Maruyama, S., 1997. Pacific-type orogeny revisited: Miyashiro-type orogeny
proposed. The Island Arc 6, 91–120.
Maruyama, S., Liou, J.G., 1983. The stability of Ca–Na pyroxene in low-grade
metabasites of high-pressure intermediate facies series. American Mineralogist 70, 16–29.
Maruyama, S., Seno, T., 1986. Orogeny and relative plate motions – an example
of the Japanese islands. Tectonophysics 127, 1–25.
Maruyama, S., Liou, J.G., Terabayashi, M., 1996. Blueschists and eclogites of
the world and their exhumation. International Geology Review 38, 485–594.
Miyashiro, A., 1961. Evolution of metamorphic belts. Journal of Petrology 2,
277–311.
Miyashiro, A., 1973. Metamorphism and Metamorphic Belts. John Wiley and
Sons, New York. 492 pp.
Molnar, P., England, P.C., 1990. Temperatures, heat flux, and frictional stress
near major thrust faults. Journal of Geophysical Research 95, 4833–4856.
Moore, C.J., 1992. Fluids in accretionary prisms. Review of Geophysics 30,
113–135.
Nicolls, I.A., Ringwood, A.E., 1973. Effect of water on olivine stability in
tholeiites and production of silica-saturated magmas in the island arc
environment. Journal of Geology 81, 285–300.
Ohtani, E., Shibata, T., Kubo, T., Kato, T., 1995. Stability of hydrous phases in
the transition zone and the uppermost part of the lower mantle. Geophysical
Research Letters 22, 2553–2556.
Ohtani, E., Mizohata, H., Kudoh, Y., Nagase, T., Arashi, H., Yurimoto, H.,
Miyagi, I., 1997. A new hydrous silicate, a water reservoir in the upper part
of the lower mantle. Geophysical Research Letters 24, 1047–1050.
Okamoto, K., Maruyama, S., 1999. The high pressure stability limits of
lawsonite in the MORB + H2O system. American Mineralogist 84,
362–373.
Okamoto, K., Maruyama, S., 2004. The eclogite-garnetite transformation in the
MORB+H2O system. Physics of the Earth and Planetary Interiors 146,
283–296.
Okamoto, K., Shinjoe, H., Katayama, I., Terada, K., Sano, Y., Johnson, S., 2004.
SHRIMP U-Pb dating of quartz-bearing eclogite from the Sanbagawa Belt,
SW Japan; Implications for Protolith and Metamorphic evolution. Terra
Nova 16, 81–89 (Oleskevich).
Oleskevich, D.A., Hyndman, R.D., Wang, K., 1999. The updip and downdip
limits to great subduction earthquakes: thermal and structural models of
Cascadia, south Alaska, SW Japan, and Chile. Journal of Geophysical
Research 104, 14,965–14,991.
Otsuki, M., Banno, S., 1990. Prograde and retrograde metamorphism of
hematite-bearing basic schists in the Sanbagawa belt in central Shikoku.
Journal of Metamorphic Geology 8, 425–439.
Oxburgh, E.R., Turcotte, D.L., 1970. Thermal structure of island arcs.
Geological Society of America Bulletin 81, 1665–1688.
Pacalo, R.E.G., Parise, J.B., 1992. Crystal structure of superhydrous B, a
hydrous magnesium silicate synthesized at 1400 °C and 20 GPa. American
Mineralogist 77, 681–684.
Pawley, A.R., Wood, B.J., 1996. The low-pressure stability of phase A,
Mg7Si2O8(OH)6. Contributions to Mineralogy and Petrology 124, 90–97.
Peacock, S.M., 1987. Thermal effects of metamorphic fluids in subduction
zones. Geology 15, 1057–1060.
Peacock, S.M., 1990. Fluid processes in subduction zones. Science 248,
329–337.
Peacock, S.M., 1991. Numerical simulation of subduction zone pressure–
temperature–time paths: constraints on fluid production and arc magmatism. Philosophical Transaction of Royal Society of London A 335,
341–353.
Peacock, S.M., 1992. Blueschist–facies metamorphism, shear heating and P–
T–t paths in subduction shear zone. Journal of Geophysical Research 97,
17,693–17,707.
Peacock, S.M., 1993. The importance of blueschist – eclogite dehydration
reactions in subducting oceanic crust. Geological Society of America
Bulletin 105, 684–694.
Peacock, S.M., 1996. Thermal and petrological structure of subduction zones. In:
Bebout, G.E., et al.(Eds.) Subduction: Top to Bottom. Monograph, vol. 96.
American Geophysical Union, pp. 119–133.
Peacock, S.M., Rushmer, T., Thompson, A.B., 1994. Partial melting of
subducting oceanic crust. Earth and Planetary Science Letters 121,
227–244.
Poli, S., 1993. The amphibolite–eclogite transformation: an experimental study
on basalt. American Journal of Science 293, 1061–1107.
Poli, S., Schmidt, M.W., 1995. H2O transport and release in subduction zones:
experimental constraints on basaltic and andesitic systems. Journal of
Geophysical Research 100, 22,299–22,314.
Ringwood, A.E., 1975. Composition and Petrology of the Earth's Mantle.
McGraw-Hill, New York. 618 pp.
Ringwood, A.E., Major, A., 1967. A high pressure transformations in
pyroxenes. Earth and Planetary Science Letters 1, 351–357.
Schmidt, M.W., 1996. Experimental constraints on recycling of potassium from
subducting oceanic crust. Science 272, 1927–1930.
Scholz, C.H., 1990. The Mechanics of Earthquakes and Faulting. Cambridge
Univ. Press, Cambridge. 439 pp.
Schreyer, W., 1988. Experimental studies on metamorphism of crustal rocks
under mantle pressures. Mineralogical Magazine 52, 11–26.
Schreyer, W., 1995. Ultradeep metamorphic rocks: the retrospective viewpoint.
Journal of Geophysical Research 100, 8353–8366.
Seno, T., Yamanaka, Y., 1996. Double seismic zones, compressional deep trench–
outer rise events, and superplumes. In: Bebout, G.E., et al.(Eds.) Subduction:
Top to Bottom. Geophysical Monograph, vol. 96. AGU, pp. 347–355.
Shieh, S.R., Ming, L., Mao, H., Hemley, R.J., 1998. Decomposition of phase D
in the lower mantle and the fate of dense hydrous silicates in subducting
slabs. Earth and Planetary Science Letters 159, 13–23.
Smyth, J.R., Kawamoto, T., 1997. Wadsleyite: II. A new high pressure hydrous
phase in the peridotite–H2O system. Earth and Planetary Science Letters
146, E9–E16.
Staudigel, H., King, S.D., 1992. Ultrafast subduction: the key to slab recycling
efficiency and mantle differentiation? Earth and Planetary Science Letters
109, 517–530.
Staudigel, H., Davies, G.R., Hart, S.R., Marchant, K.M., Smith, B.M., 1995.
Large scale isotopic Sr, Nd and O isotopic anatomy of altered oceanic crust:
DSDP/ODP sites 417/418. Earth and Planetary Science Letters 130,
169–185.
Stein, S., 1995. Deep earthquakes: a fault too big? Science 268, 49–50.
Taira, A., 1985. Sedimentary evolution of Shikoku subduction zone. The
Shimanto Belt and Nankai Trough. In: Nasu, N., et al.(Eds.) Formation of
Active Ocean Margins. Terra Science Publication, Tokyo, pp. 273–290.
Takasu, A., Dallmeyer, R.D., 1990. 40Ar/39Ar age constraints for the
tectonothermal evolution of the Sanbagawa metamorphic belt, central
Shikoku, Japan: a Cretaceous accretionary prism. Tectonophysics 185,
303–324.
Takasu, A., Wallis, S., Banno, S., Dallmeyer, R.D., 1994. Evolution of the
Sanbagawa metamorphic belt, Japan. Lithos 33, 119–133.
Tatsumi, Y., Hamilton, D.L., Nesbitt, R.W., 1986. Chemical characteristics of
fluid phase released from a subducted lithosphere and origin of arc magmas:
evidence from high-pressure experiments and natural rocks. Journal of
Volcanology and Geothermal Research 29, 293–309.
Thompson, A.B., 1992. Water in the Earth's upper mantle. Nature 358,
295–302.
Titchelaar, B.W., Ruff, L.J., 1993. Depth of seismic coupling along subduction
zones. Journal of Geophysical Research 98, 2017–2037.
Toksöz, M.N., Hsui, A.T., 1989. A review of the thermal–mechanical structures
at convergent plate boundaries and their implications for crust/mantle
recycling. In: Hart, S.R., Gulen, L. (Eds.), Crust/Mantle Recycling at
Convergent Zones. Kluwer Academic, Norwell, Mass, pp. 75–80.
Toksöz, M.N., Minear, J.W., Julian, B.R., 1971. Temperature field and
geophysical effects of a downgoing slab. Journal of Geophysical Research
76, 1113–1138.
Turcotte, D.L., Schubert, G., 1973. Frictional heating of the descending
lithosphere. Journal of Geophysical Research 78, 5876–5886.
Ulmer, P., Trommsdorff, V., 1995. Serpentine stability to mantle depths and
subduction-related magmatism. Science 268, 858–861.
S. Maruyama, K. Okamoto / Gondwana Research 11 (2007) 148–165
van den Beukel, J., Wortel, R., 1988. Thermo-mechanical modeling of arc–
trench regions. Tectonophysics 154, 177–193.
Wang, C.-Y., Shi, Y.-L., 1984. On the thermal structure of subduction complexes:
a preliminary study. Journal of Geophysical Research 89, 7709–7718.
Wyllie, P.J., 1988. Magma genesis, plate tectonics, and chemical differentiation
of the Earth. Review of Geophysics 26, 370–404.
Wyllie, P.J., Sekine, T., 1982. The formation of mantle phlogopite in subduction
hybridization. Contributions to Mineralogy and Petrology 79, 375–380.
Yamamoto, K., Akimoto, S., 1977. The system MgO–SiO2–H2O at high
pressures and temperatures – stability field for hydroxyl–chondrodite,
165
hydroxyl–clinohumite and 10 Å-phase. American Journal of Science 277,
288–312.
Yang, H., Prewitt, C.T., Frost, D.J., 1997. Crystal structure of the dense hydrous
magnesium silicate, Phase D. American Mineralogist 82, 651–654.
Zhang, J., Li, B., Utsumi, W., Liebermann, R.C., 1996. In situ X-ray
observations of the coesite–stishovite transition: reversed phase boundary
and kinetics. Physics of Chemistry and Mineralogy 23, 1–10.