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stability field. There is also major pressure
dependence of solidus temperatures inside the
olivine stability field at 10-14GPa and with
0.3-0.5 wt % H
2
O in the system (Figure 2.6).
The maximum H
2
O content in mantle olivines
from kimberlitic xenoliths (0.02-0.04 wt %; e.g.
Matsyuk & Langer, 2004) remains within the
olivine storage capacity at 6GPa and 1300
◦
C,
which is 0.07-0.10 wt % H
2
O. The maximum
H
2
O content of wadsleyite and ringwoodite in the
transition zone is 0.4-0.5 wt % along the mantle
adiabat (Litasov
et al
., 2011b). Considering the
presence of other minerals (up to 40% garnet and
pyroxenes with up to 0.1 wt % H
2
O), the transi-
tion zone may store up to 0.3-0.35 wt % H
2
Oat
average mantle temperature.
The dependence of the H
2
O solubility on
pressure and temperature in eclogite minerals is
poorly studied. At pressures above the stability
field of phlogopite, lawsonite, and phengite,
H
2
O can accumulate in accessory richterite,
phlogopite, or nominally anhydrous phases.
Some studies recorded the appearance of the
K-bearing phase
X
at pressures above 15GPa
when K-richterite disintegrated; this phase could
contain up to 1.5 wt % H
2
O and it may occur
both in eclogitic and peridotitic assemblages if
the system has an elevated K content (Konzett
& Fei, 2000). Aluminous stishovite may also
be a major reservoir for H
2
O in eclogite at
pressures above 20GPa as it can contain up
to 0.3 wt % H
2
O(Litasov
et al
., 2007). Pyrope
and majorite garnet can contain up to 0.1 and
0.13 wt % H
2
O, respectively (Katayama
et al
.,
2003; Mookherjee & Karato, 2010). According
to the available data, the H
2
O solubility in
Na-clinopyroxene is lower than that in garnet.
In the experiments conducted at 600-700
◦
C
(Bromiley & Keppler, 2004), it decreases from
470 ppm at 2GPa to 100 ppm at 10GPa.
The phase boundaries between olivine and wad-
sleyite and ringwoodite and Mg-perovskite
mantle water contents from the depths of the
410 km and 660 km discontinuities, especially in
the regions close to subduction zones.
2.6.2 Systems with CO
2
The peridotite-CO
2
and eclogite-CO
2
systems
were studied in the simplified (CaO-MgO-
Al
2
O
3
-SiO
2
Na
2
O-CO
2
) (Litasov & Ohtani,
2009a; Keshav & Gudfinnsson, 2010; Litasov &
Ohtani, 2010) and in complex, close to natural
compositions (Ghosh
et al
., 2009; Kiseeva
et al
.,
2013) at pressures of up to 21-32GPa, mainly
along the solidus temperatures. The solidi in
these systems are subparallel or have very
gentle PT-slopes above 10GPa. It was found that
Na
2
OandK
2
O play a key role in the melting
of carbonate-bearing peridotite and eclogite.
The addition of 0.1 wt % K
2
O reduces the
solidus temperature by 500
◦
C at 20GPa in both
systems. In addition to the solidus curves, the
thermal stability of carbonate phases (magnesite
and aragonite) is important to constrain the
behavior of carbon in the mantle. Magnesite and
aragonite stability in different systems, including
those with H
2
O
CO
2
(Litasov
et al
., 2011a)
is shown in Figure 2.7. There are three major
regimes of magnesite stability under oxidized
conditions: (i) magnesite-bearing silicate systems
without free silica phase, (ii) magnesite
+
SiO
2
assemblages (for example in eclogite), (iii) H
2
O-
saturated magnesite-bearing systems. In the first
regime magnesite stability is limited by decarbon-
ation and melting reactions involving silicates,
such as MgCO
3
+
+
CO
2
.
In the second regime, magnesite stability is
controlled by melting reactions with silica
phases MgCO
3
+
MgSiO
3
=
Mg
2
SiO
4
+
CO
2
.The
addition of H
2
O to the system causes the thermal
stability limit of magnesite to fall dramatically to
the level of the solidi of K
2
O-containing systems.
The stability lines plot parallel to the pressure
axis. The reason for such a drastic reduction
in magnesite stability is poorly understood at
present. In general, the thermal stability limit of
magnesite in eclogite is lower than in peridotite
(Litasov
et al
., 2011a).
SiO
2
=
MgSiO
3
+
+
periclase shift toward low and high pressures,
respectively, in the presence of H
2
O(Litasov
et al
., 2005; Litasov
et al
., 2006; Frost & Dolejs,
2007) due to the different H
2
O solubility in the
minerals. These data can be used to estimate
