Supplementary Information
S1. Details of the Mineralogical Model
We chose antigorite as our model serpentine mineral because it is the principal retrograde
metamorphic mineral in low-Al altered ultramafic systems at intermediate temperatures (Evans
et al., 1976; Evans, 2004; O’Hanley, 1996; Deer, Howie and Zussman, 2009), its dehydration
curve is well characterized to high pressures and it has a simple serpentine breakdown reaction
involving no other starting phases. Among the common rock-forming serpentine minerals,
antigorite is also stable to the highest temperatures at a given water pressure. Other associated
minerals are also likely present in the serpentinite assemblage in the cold forearc mantle wedge,
such as lower-temperatures forms of serpentine minerals (chrysotile and lizardite), brucite, talc
and magnetite. Their presence during post-subduction forearc heating and dehydration would
contribute to water release over a broader range of temperatures and spatial distributions than
predicted by our simplified models. This would be especially true if we had adopted a shallower
crust-mantle boundary where these lower-temperature dehydration reactions come into play. We
also ignore the effects of the fayalite component of olivine in peridotites
Table S1: Properties of Antigorite as a Model Serpentine Mineral
Formula
Formula weight
Molar volume,density§
Heat capacity§
Dehydration Reaction
Equilibrium dehydration
curve Te = f(P, GPa),
Slope dTe/dP, K GPa-1
Water loss
V volume change cm3mol-1
Mg48Si34O85(OH)62
4335.97 g
1749.14 cm3 mole-1, 2.596 g/cm3 both at STP
Table 132 in Robinson et al. (1983)§
antigorite 18 forsterite + 4 talc + 27 water
Te (K) = 482.81 + 220.01 P – 121.597P2
+47.138 P3 –7.0138 P4, R = 0.99999, =1 K
dTe/dP=220.01+243.19P+141.415P2-28.0552P3
10.71 weight % (1.60 weight % left in talc)
V = 110.45 – 270.09.(log10 P(GPa))
for antigorite, its variation
based on the phase proportions and the
with P along the dehyd-
equations of state for the phases (from various
ration curve. Calculated
from references in Ref. §.
data sources)
H enthalpy change
H = TV/(dTe/dP)
Data sources
H = 1.645 E6 – 6.0066E5 P in J mol-1 where P is
the water pressure in GPa over the pressure range of 0.5
to 2.25 GPa.
Evans et al (1976); Ulmer and Tromsdorf (1995);
§
Robinson and others (1983)

S2. Details of the Computational Model
Eight-node isoparametric finite elements are used for the two-dimensional thermal model. Thermal
conductivities are assumed to be 2.5 and 2.0 W m-1K-1 for the upper and lower continental crusts (both 15
km thick), respectively, 3.1 W m-1 K-1 for the continental mantle, and 2.9 W m-1K-1 for the subducting
slab. Radiogenic heat production is assumed to be 1.2 and 0.6 W m-3 for the upper and lower continental
crusts, respectively, and 0.02 W m-3 for the rest of the model. Volumetric heat capacity is assumed to be
2.7 x 106 J m-3K-1 and 3.3 x 106 J m-3K-1 for the continental crust and the rest of the model, respectively.
For the stalled-slab scenario, the initial temperature field is calculated for a much larger model domain
assuming a 10-Ma old slab subducting at 4 cm/yr at the steady state, with boundary conditions and mantle
wedge flow introduced in exactly the same way as described in Peacock and Wang (1999). This initial
temperature field is used also for the no-slab-window scenario but after being modified by assigning a
uniform "filling" temperature (1100°C or 1400°C for the model) to all nodes below the slab. Time
marching from the initial condition is performed using a fully implicit algorithm. The upper and lower
surfaces of the model are kept at 0 and 1400°C, respectively, and the seaward boundary is assigned the
geotherms for a 30-Ma old oceanic lithosphere, roughly the age of the present-day Pacific plate off
California. For the landward boundary, the upper 40 km has geotherms consistent with a 80 mW m-2
surface heat flow, and the nodal temperatures of the deeper part are assigned by assuming an adiabatic
mantle thermal gradient of 3 10-4 K m-1. The effect of heat absorption during antigorite breakdown is
included using an "epoch average" approach. A time epoch is chosen to be long enough to allow a
significant number of elements in the old serpentinized mantle wedge to experience antigorite breakdown
(dehydration) and to become heat sinks. The total amount of heat to be absorbed by these elements
divided by the length of the epoch gives an average, constant heat sink rate over the epoch. The size of the
region that experiences antigorite breakdown and the average heat sink rate are nonlinearly interrelated
and are thus determined iteratively for each epoch. The use of larger epochs ensures numerical stability
but degrades time resolution. For our problem, and epoch length of 5 Ma is found to be a good
compromise. We assume that the release of water from the serpentinized wedge into the forearc crust
above is slow enough that significant heat is not advected relative to that of simple heat conduction
(Manning and Ingebritsen, 1999; Fulton and Saffer, 2009).