Oxygen isotope geochemistry is a powerful tool for
investigating rocks that interacted with fluids, to assess fluid sources and
quantify the conditions of fluid–rock interaction. We present an integrated
modelling approach and the computer program PTLoop that combine
thermodynamic and oxygen isotope fractionation modelling for multi-rock open
systems. The strategy involves a robust petrological model performing
on-the-fly Gibbs energy minimizations coupled to an oxygen fractionation
model for a given chemical and isotopic bulk rock composition; both models
are based on internally consistent databases. This approach is applied to
subduction zone metamorphism to predict the possible range of δ18O values for stable phases and aqueous fluids at various pressure
(P) and temperature (T) conditions in the subducting slab. The modelled system
is composed of a mafic oceanic crust with a sedimentary cover of known
initial chemical composition and bulk δ18O. The evolution of
mineral assemblages and δ18O values of each phase is calculated
along a defined P–T path for two typical compositions of basalts and sediments.
In a closed system, the dehydration reactions, fluid loss and mineral
fractionation produce minor to negligible variations (i.e. within 1 ‰) in the bulk δ18O values of the rocks, which
are likely to remain representative of the protolith composition. In an open
system, fluid–rock interaction may occur (1) in the metasediment, as
a consequence of infiltration of the fluid liberated by dehydration reactions
occurring in the metamorphosed mafic oceanic crust, and (2) in the
metabasalt, as a consequence of infiltration of an external fluid originated
by dehydration of underlying serpentinites. In each rock type, the
interaction with external fluids may lead to shifts in δ18O up to 1 order of magnitude larger than those calculated for closed systems. Such
variations can be detected by analysing in situ oxygen isotopes in key
metamorphic minerals such as garnet, white mica and quartz. The simulations
show that when the water released by the slab infiltrates the forearc mantle
wedge, it can cause extensive serpentinization within fractions of 1 Myr and
significant oxygen isotope variation at the interface. The approach
presented here opens new perspectives for tracking fluid pathways in
subduction zones, to distinguish porous from channelled fluid flows, and to
determine the P–T conditions and the extent of fluid–rock interaction.
Introduction
The subducting oceanic slab is composed of a sequence of rock types
corresponding to chemical systems that undergo continuous and discontinuous
reactions in response to pressure (P) and temperature (T) changes. Through its
metamorphic history, the hydrated oceanic lithosphere undergoes extensive
dehydration by the breakdown of low-temperature, volatile-rich minerals
(e.g. Baumgartner and Valley, 2001; Baxter and Caddick, 2013; Hacker, 2008;
Manning, 2004; Page et al., 2013; Poli and Schmidt, 2002). The expelled
aqueous fluid migrates through the slab towards the slab–mantle interface, and it may continue rising to the mantle wedge playing a major role in
triggering mass transfer and melting (Barnicoat and Cartwright, 1995; Bebout
and Penniston-Dorland, 2016). Evidence for fluid circulation in subducted
rocks has been extensively observed in exhumed
high-pressure/ultra-high-pressure (HP/UHP) terrains (e.g. Zack and John,
2007; Baxter and Caddick, 2013; Martin et al., 2014; Rubatto and Angiboust, 2015;
Engi et al., 2018), but a direct link to the primary source production is often
missing and the main source remains matter of debate. The characterization
of fluid pathways in subduction zones has been addressed by using a variety
of methods (i.e. seismicity, thermodynamic modelling, fluid inclusions, HP
veins, trace element and stable isotope studies on metamorphic minerals)
(e.g. Baxter and Caddick, 2013; Hacker, 2008; Hernández-Uribe and
Palin, 2019; Scambelluri and Philippot, 2001; Spandler and Hermann, 2005).
In particular, oxygen isotope composition of metamorphic minerals from
exhumed HP rocks sheds light on the nature of the fluid reacting with those
systems during metamorphism. Thus, oxygen isotope studies of HP rocks have
the potential to make important contributions to the investigation of fluid
sources and pathways in subduction zones (e.g. O'Neil and Taylor, 1967;
Muehlenbachs and Clayton, 1972; Baumgartner and Valley, 2001;
Page et al., 2013; Martin et al., 2014; White and Klein, 2014; Hoefs, 2015; Rubatto and Angiboust,
2015).
Schematic geometry of the subduction models discussed in
the text. The rock column is composed of two rock types (Rock 1 and Rock 2) that can be
infiltrated by an external fluid deriving from a third layer located beneath
them. (a) Example columns used in the calculation along the P–T path shown in
Fig. 2 to produce the results presented in Figs. 3–7. See text for
details. (b) Schematic representation of the three interaction cases
discussed in the text. No-interaction case (NI): the fluid released by the metabasalt does
not interact with the metasediment. The fluid leaving the system is a
mixture of metabasalt-derived and metasediment-derived fluids. Partial-interaction case (PI): 50 % of the metabasalt-derived fluid does not interact with the
metasediment, and 50 % of it equilibrates with the metasediment. The
final fluid released by the system is the mixture between the unmodified
metabasalt-derived fluid and the fluid deriving from the metasediment after
it equilibrates with 50 % of the metabasalt-derived fluid. High-interaction case (HI): all
the fluid released by the metabasalt equilibrates with the metasediment. The
fluid leaving the system exits the metasediment. (c) Possible scenario at
the base of the column. As a consequence of serpentine breakdown,
serpentinite-derived fluids may infiltrate the metabasalt, exchange with it
and affect the fluid infiltrating the metasediment.
The modelling of oxygen isotopic fractionation has been traditionally addressed
as an equilibrium calculation between individual mineral couples. An
alternative approach follows what has been extensively adopted in the last
decades for thermodynamic modelling (see reviews by Lanari and Duesterhoeft,
2019; Powell and Holland, 2008; Spear et al., 2017) and considers an
evolving mineral assemblage. A pioneer model proposed by Kohn (1993) can be
applied to single and closed chemical systems, i.e. for which no
infiltration of external fluids in isotopic disequilibrium is allowed. Such
an approach can simulate how the oxygen isotopic composition changes with
P and T, but it remains too simple for subduction zone settings, where
significant fluid exchange occurs between different lithologies within the
subducting slab. Baumgartner and Valley (2001) proposed a model for stable
isotope fluid–rock exchange based on continuum mechanics, where infiltration
profiles can be calculated, but no information is provided about the
different components (minerals) of the rock, as it is regarded as a
continuum. The fluid / rock (F/R) ratios obtained with this strategy do not
correspond to the physical amount of fluid but rather represent a
measurement of exchange progress.
The subducting oceanic lithosphere is typically composed of a section of
igneous oceanic crust with its sedimentary cover (mostly < 1 km)
above and an ultramafic lithospheric mantle section beneath. The geometry of
the model is illustrated in Fig. 1. The target column represents a
simplified section of the upper part of such oceanic lithosphere. It is
composed of a layer of basaltic composition (Rock 1) overlaid by a layer of
sediments (Rock 2, see below for details). Two different rock columns are
considered: (1) a relatively water-rich system with altered basalts and
terrigenous sediments and (2) a relatively water-poor system with unaltered
basalts (of mid-ocean ridge basalt (MORB) composition) and carbonate sediments. The column has a fixed
section of 1 m2, while the thickness of each rock unit can be set by
the user. The model is conservative with respect to the mass, while the
volume of each rock type changes according to fluid loss and density
variation along the P–T path. The P–T structure of subduction zones depends
on numerous variables, including the age of the incoming lithosphere and the
amount of previously subducted lithosphere (e.g. Peacock, 1990). In this
study, the calculation was performed following the subduction geotherm from
Gerya et al. (2002) (Fig. 2) over a pressure range of 1.3–2.6 GPa,
corresponding to a depth of ∼45 to ∼85 km to
encompass the conditions of interest for the investigated processes. The
modelled temperatures range from a minimum of 350 ∘C to a maximum
of 700 ∘C. The lower limit was chosen to take into account the
large uncertainties in the thermodynamic databases for many common low-T
metamorphic minerals that lead to unsatisfactory models for phase equilibria
and mineral parageneses for low-grade rocks (Frey et al., 1991; Vidal et al.,
2016). The upper limit is fixed by the wet solidus of metasediments – the
present model strictly applies to subsolidus conditions.
P–T diagram showing typical oceanic subduction geotherms and
the intermediate geotherm used in the calculation (red line, Gerya et al.,
2002). The red spots represent the modelling steps. The average D80 geotherm
from Syracuse et al. (2010) (purple dashed line), i.e. the geotherm dominated by a steep T
gradient at 80 km depth, which occurs at the transition from partial to full
coupling, as reported by Penniston-Dorland et al. (2015) and the average slab-top geotherm from Penniston-Dorland et al. (2015)
(green dashed lines) are shown for comparison. Metamorphic facies modified
from Peacock (1993) and Liou et al. (2004). Serpentine breakdown reactions
from Padrón-Navarta et al. (2010) (PN10) and Hermann et al. (2000)
(H00). Mineral abbreviations from Whitney and Evans (2010).
During burial and heating the different slab lithologies undergo dehydration
reactions. The produced fluid escapes from the source rock and migrates
upward, likely interacting with the surrounding units of different chemical
and isotopic composition. The effect of an external fluid input on the
δ18O value of growing minerals is strongly dependent on the
isotopic composition of the infiltrating fluid (δ18Ofluid)
and on the degree of fluid–rock interaction. The integrated F/R ratio is
defined here as the total mass of aqueous fluid that has passed through and
interacted with the rock normalized to the mass of the rock. To explore
different scenarios, three models are discussed involving different
associations of fresh or altered oceanic basalts with terrigenous or
carbonate sediments (Fig. 1b): (1) the interaction between the fluid
released from the metamorphosed mafic crust and the overlying metasediment
is negligible and the two rocks evolve independently (no-interaction case, NI); (2) part of
the fluid derived from the metabasalt (50 % when not specified
differently) equilibrates with the metasediment, while the other part leaves
the system (partial-interaction case, PI); and (3) all the fluid released by the metamorphosed
mafic crust equilibrates with the metasediment before escaping the system
(high-interaction case, HI). The fluid released by the entire system is a mixture of fluids
derived from the progressive dehydration of the metabasalt and
metasedimentary layers. In the NI case, both rock types behave like closed
systems and the fluid is liberated from the metabasalt and from the
metasediment separately. In the case of infiltration of fluid derived from the
metabasalt in the metasediment, the amount of fluid released by this latter
includes the fluid produced by dehydration reactions plus the excess fluid
that enters the metasediment and cannot be incorporated into stable hydrous
minerals.
The thickness and the degree of serpentinization of the lithospheric mantle
subducting beneath the oceanic crust can be highly variable. The most
important dehydration reactions in partly or fully serpentinized mantle are
related to antigorite breakdown, which can release up to 12 wt % of water,
playing an important role for water flows in subduction zones.
Deserpentinization is assumed to result in two main subsequent fluid peaks
(Padrón-Navarta et al., 2013; Scambelluri et al., 2004) related to the
reactions
1antigorite+brucite→olivine+chlorite+water∼480∘C,1.7GPa;1to2wt%waterreleased,upto∼50kgm-3,2antigorite→olivine+orthopyroxene+chlorite+water∼660∘C,2.5GPa;≥6.5wt%waterreleased,≥170kgm-3.
The effect on the δ18O of the metabasalts and metasediments of
an external fluid influx, i.e. caused by dehydration of the underlying
serpentinites, was investigated by defining an amount of fluid with a
specific δ18O value that infiltrates the basaltic layer (Rock 1) at
two steps of the model (480 and 660 ∘C along the
chosen P–T path) (Figs. 1c, 2).
Model strategy
The strategy behind PTLoop consists of forward modelling the
evolution of the mineral assemblage and the oxygen isotope composition of a
rock column composed of two lithologies (Fig. 1a) of assigned thickness and
starting bulk chemical and oxygen isotope compositions along a defined P–T path
using a stepwise procedure (Fig. 2). At each P–T step, (1) the equilibrium
mineral assemblage, oxygen isotope composition of stable phases (δ18O ‰ vs. Vienna Standard Mean Ocean Water, VSMOW),
mass (in kg) and isotopic composition of the excess fluid for the metabasalt
are calculated; (2) any fraction of excess fluid deriving from dehydration
reactions in the metabasalt can be transferred to the metasediment or
directly escapes the system; (3) the equilibrium mineral assemblage, δ18O value of stable phases, and amount and δ18O of the excess
fluid for the metasediment are evaluated by accounting for the changes caused by
the fluid input from the metabasalt; (4) the mass and δ18O of
the total fluid leaving the system are calculated. Furthermore, at each step
a chosen amount of external fluid with a given δ18O can be
input into the metabasalt and its contribution is accounted for in the
subsequent steps. This model is based on the assumption of thermodynamic
equilibrium applied to a partially reactive system, whereby phases are
assumed to reach chemical and isotopic equilibrium at all steps within the
reactive part of the system, i.e. removing from the reactive bulk the phases
that are fractionated (Lanari and Engi, 2017). Such petrological models can
account for element sequestration during prograde metamorphism. Mineral
fractionation in relics and fluid input or loss are two processes that are
allowed to modify the reactive bulk composition. No mineral resorption is
permitted. Any fluid liberated during dehydration (excess fluid) does not interact further and leaves the rock. This process is termed Rayleigh volatilization (Rumble, 1982;
Valley, 1986). In natural rocks, it often occurs combined with the opposite
endmember, the batch volatilization,where the produced fluid stays within the system as the
mineral reaction proceeds and remains in isotopic equilibrium with the rock
until the reaction is completed. In most natural cases involving oxygen
isotopes, the difference between the results calculated using the two
processes is negligible (Baumgartner and Valley, 2001). The released fluid
is regarded as pure H2O. Any other effect related to solute transport by the fluid is ignored in the calculation; the potential effect of a
CO2 component in the fluid is discussed in Sect. 3.2.
Governing equations
Equilibrium assemblage calculation for a given bulk rock composition at any
P and T is performed with the program Theriak (de Capitani and Brown, 1987; de
Capitani and Petrakakis, 2010) and is based on Gibbs energy minimization. A
complete description of the minimization algorithm is given by de Capitani
and Brown (1987). The reacting bulk composition may evolve in the course of
the metamorphic history of a rock because of mineral fractionation, fluid
loss or input of external fluids. The effective bulk composition is
recalculated by PTLoop at each subsequent stage following the
strategy of Lanari et al. (2017).
As for phase assemblage determination, equilibrium is a common assumption of
stable isotope transport (Baumgartner and Rumble, 1988; Baumgartner and
Valley, 2001; Bowman et al., 1994; Gerdes et al., 1995a, b). Thus, a
molar equilibrium constant (K) can be defined to describe the thermodynamic
stable isotope equilibrium between two substances i and j (Sharp, 2017):
K=18Oi/16Oi18Oj/16Oj.
The fractionation factor (α) can be related to the equilibrium
constant K as
α=K1/n,
where n is the number of exchanged atoms, normally 1 for simplicity. In
isotope geochemistry, the isotopic composition is commonly expressed in
terms of δ values:
δi=RiRSt-1⋅103(‰),
where Ri and RSt are the isotope ratio measurements for the
compound i and the defined isotope ratio of a standard sample respectively.
For differences in δ values or for δ values of less than
∼10 ‰,
1000lnαi-j=δi-δj
is a valid approximation that is used in most cases (Hoefs, 2015; Sharp,
2017). For oxygen isotope fractionation, the equation that can reproduce
most of the available calibrations describing the stable isotope
fractionation function between two phases is a second-order polynomial of
103/T. Hence the stable isotope fractionation between two phases i (with k
endmembers) and j (pure) as a function of T is described by Eq. (7):
δ18Oi-δ18Oj=∑1kAk,j⋅106T2+Bk,j⋅103T+Ck,j⋅Xk,i⋅Nk,iNi.
The conservation of the bulk δ18O in the system is described by Eq. (8):
δ18Osys⋅Nsys=∑k=1pMk⋅Nk⋅δ18Ok,
where δ18Oi, δ18Oj and δ18Osys are
the isotopic compositions of phase i, phase j and the system (bulk δ18O) respectively; Ak,j, Bk,j and Ck,j are the
fractionation parameters for endmember k of mineral i vs. phase j; Xk,i is the
fraction of endmember k in the phase i; Nk,i, Ni and Nsys are the
total number of moles of oxygen in endmember k, in mineral i and in the
system respectively; p is the number of phases; Mk is the number of moles
of phase k; Nk is the its number of oxygen; and δ18Ok is its
oxygen isotope composition. Given a stable mineral assemblage at any P–T
condition, the oxygen isotope partitioning among the stable phases is
calculated by solving the linear system described by the sets of p-1 equatiomns of type (7) and Eq. (8); Kohn, 1993; Vho et al., 2020). In closed systems,
the first term in Eq. (8) is constant. Open-system behaviour can either
modify the δ18Osys or the number of moles of the phases
(Nsys). The parameters A, B and C between phases were taken from the
internally consistent database for oxygen isotope fractionation
DBOxygen version 2.0.3 (Vho et al., 2020).
Starting assumptions
In order to represent the variability in the basaltic portion of the oceanic
crust two different bulk compositions were used (Table 1): (1) a
representative average MORB basalt (Gale et al., 2013) that has been
hydrated but without any other addition or removal of element
(metabasalt(h)) and (2) a basalt that underwent extensive sea-floor
alteration during hydration (Baxter and Caddick, 2013 after Staudigel et
al., 1996) that will be referred to as metabasalt(a) in the following.
Those compositions are in good agreement with other compilations reported in
the literature (e.g. Sun and McDonough, 1989; Albarède, 2005; Staudigel,
2014; White and Klein, 2014). Oceanic sediments were modelled with two
distinct bulk compositions (Table 1): (1) terrigenous sediment referred to as metasediment(t) (clay from the Mariana Trench; Hacker, 2008, after Plank and
Langmuir, 1998) and (2) nanno ooze carbonate referred to as metasediment(c) (Plank,
2014). Nanno oozes are widespread carbonate sea-floor sediments (Plank,
2014), and they are close in composition to carbonate-rich sediments observed
in HP terrains (e.g. Bebout et al., 2013; Kuhn et al., 2005). Thicknesses
of 1000 m for the basaltic layer, of 175 m for the clay sediment and 75 m
for the carbonate sediment were chosen in order to maintain proportions
between oceanic crust and sediments comparable with the values reported in
various compilations (e.g. Hacker, 2008; Plank, 2014). This results in a
total thickness of the rock column 2 to 3 times smaller than the real
thickness to encompass the assumption of homogeneous temperature over the
whole column within ∼20∘C. To overcome the
effects of possible temperature variations within the column, a
discretization step size of ∼20∘C along the P–T path
was applied.
1 Gale et al. (2013).
2 Baxter and Caddick (2013) after Staudigel et al. (1996).
3 Mariana clay from Plank and Langmuir (1998).
4 Nanno ooze from Plank (2014).
5 Used for the thermodynamic modelling.
6 Water content at saturation at 350 ∘C and 1.3 GPa. The calculated initial water content is consistent with values from the
literature at these conditions (e.g. Poli and Schmidt, 1998; Hacker et al.,
2003). NA: not available.
The bulk compositions were simplified to the
Na2O–CaO–K2O–MgO–FeO–Al2O3–TiO2–SiO2–H2O
system; C is present in the initial bulk composition of the
metabasalt(a) and the metasediment(c) (Table 1). MnO was excluded
because it overemphasizes the stability of garnet at low metamorphic
conditions (T≤350∘C). The conditions of the garnet-in
reaction in Mn-absent systems match the results obtained for garnet
nucleation in natural rocks (e.g. Laurent et al., 2018). Thermodynamic
modelling was performed using the internally consistent dataset of Holland
and Powell (1998) and subsequent updates (tc55, distributed with
Theriak-Domino 4 February 2017; see Supplement S1). The
following activity models where used for the solid solutions: Holland and
Powell (2003) for calcite–dolomite–magnesite; Holland and Powell (1998) for
garnet, white mica and talc; Holland and Powell (1996) for omphacite;
Holland et al. (1998) for chlorite; Diener et al. (2007) for amphibole. In
the presented model garnet undergoes fractional crystallization both in
Rock 1 and Rock 2 fractionating from the reactive bulk for the subsequent steps. The amount
of initial H2O in each rock was set at saturation and is reported in
Table 1. No pore fluid expulsion, diagenetic and low-grade (T < 350 ∘C) devolatilization reactions are considered in this study (see
above).
A starting bulk δ18O for the metabasalt(h) of 5.7 ‰ was chosen and represents the reference value for an
unaltered MORB (e.g. Cartwright and Barnicoat, 1999; Eiler, 2001;
Staudigel, 2014; White and Klein, 2014), while a starting bulk δ18O of 9.0 ‰ is used for metabasalt(a),
representative of basaltic material that underwent sea-floor alteration at
T≤400∘C (e.g. Alt et al., 1986; Cartwright and Barnicoat,
1999; Eiler, 2001; Gregory and Taylor Jr., 1981; Miller and Cartwright, 2000;
Staudigel, 2014; White and Klein, 2014). The starting bulk of the
terrigenous sediment of 15 ‰ represents the average for
the δ18O of clastic sediments reported by Eiler (2001). The
chosen δ18O starting bulk of the carbonate sediment is 25 ‰, which represents a conservative estimate of marine
carbonate δ18O (typically 25 ‰–35 ‰; Eiler, 2001). It is ∼5 ‰ higher than the
values for metasedimentary carbonates in the Italian Alps (e.g.
Cook-Kollars et al., 2014) that are likely to have interacted with lower-
δ18O fluids during subduction.
In order to define the contribution of an external fluid originating in the
lithospheric mantle by serpentine breakdown, a layer of 150 m of pure
serpentine containing 12 wt % bulk H2O was considered and the mass of
water released at each reaction was calculated by mass balance, resulting in
an input of 7800 kg of water at 480 ∘C and of 25 350 kg at 660 ∘C to satisfy the reactions in Eqs. (1) and (2) respectively. In order to fit
the thicknesses chosen for the oceanic crust and the sedimentary layer (2 to
3 times thinner than an average lithospheric section), the 150 m of pure
serpentine correspond to a conservative estimate of 3000 m of serpentinized
peridotite with an average serpentine content of 5 % in volume. This is
in agreement with the values used by Barnes and Straub (2010) and John et
al. (2011) based on the estimate by Sharp and Barnes (2004). Serpentine
oxygen isotope compositions reported in the literature are highly variable
(Cartwright and Barnicoat, 1999, 2003; Früh-Green et al., 2001;
Mével, 2003; Miller et al., 2001), typically ranging from 1 to 10 ‰. In mid-oceanic ridge environments, the distribution
has a peak between 2 ‰ and 5 ‰ (Mével, 2003). A value
of 2.5 ‰ was chosen from the lower-δ18O
side of this peak. This results in the liberation of a fluid with a
characteristic low-δ18O value, but one which is still feasible for natural
serpentinites; this value is clearly distinct with respect to the overlying
lithologies. The effect of infiltration into the metabasalts and metasediments
of serpentinite-derived fluids will become smaller as the fluid δ18O becomes higher, approaching the equilibrium with the overlying
lithologies. The δ18O value of the released fluid is
∼4.5 ‰ at T>550∘C
(serpentine–water oxygen isotope fractionation factors from Vho et al., 2020).
Further details on the modelling input data are given in Supplement S2.
ResultsStable mineral assemblage
The evolving stable mineral assemblages and bulk water contents of each
lithology, without external fluid input, were calculated for each rock
composition along the prograde P–T path. Results are provided as mode-box
diagrams in Fig. 3. The H2O field represents the volume fraction of
excess water in each rock type. The fluid is progressively extracted
becoming isolated from the reactive part of the system. Garnet is the only
phase prevented from re-equilibrating in the model, thus fractionating from
the reactive bulk composition. Below 450 ∘C and 1.80 GPa, in the
metabasalts glaucophane, actinolite and lawsonite comprise ∼80 vol. % of the paragenesis, with minor phengite (Siapfu=3.67–3.63, XMg=0.62–0.56 in metabasalt(h); Siapfu=3.68–3.65, XMg=0.67–0.58 in metabasalt(a)), omphacite
(XNa=0.45–0.42, XMg=0.81–0.72), chlorite and
titanite. Metabasalt(h) is richer in SiO2, FeO and MgO with
respect to the metabasalt(a), and chlorite is stable up to 480 ∘C. Metabasalt(a) contains ∼5 vol. % of
Ca carbonate that remains stable over the entire P–T path. For either
composition, the volume of glaucophane, actinolite and lawsonite gradually
decreases from 480 ∘C and ∼1.90 GPa until complete
consumption at 600–620 ∘C and 2.30–2.36 GPa. Those
represent the major hydrous phases contributing to the dehydration, while a
secondary role is played by talc and zoisite at higher P–T conditions (T≥580∘C, P≥2.24 GPa). Most of the water still retained in
the rocks is stored in phengite, the abundance of which is primarily
controlled by bulk K2O content, higher in the metabasalt(a), and
that remains stable beyond the model conditions. Garnet production starts at
∼500∘C and ∼2.00 GPa (Xalm=0.60, Xgrs=0.35), and garnet grows continuously until
constituting ∼20 vol. % of the metabasalt(a) and
∼35 % of the metabasalt(h).
Mode-box diagrams showing the evolution of the mineral
assemblages and fluid during subduction of the different rock types along
the geotherm shown in Fig. 2. The initial H2O content is < 1 vol % in the metabasalts and in the metasediment(c) and ∼3 vol. % in the metasediment(t) (see Table 1 for details). Garnet
fractionation is applied to all the lithologies. The volume fraction of
garnet shown at each step represents the sum of the fractionated and newly
grown garnet. The phase proportions refer to the NI case, where the H2O
content is the excess H2O produced by each rock type evolving
independently. The excess water is fractionated at each step and the volume
fraction displayed represents the sum of the fractionated and the newly
produced water. Mineral abbreviations from Whitney and Evans (2010).
In the metasediment(c) calcium carbonate, quartz, phengite (Siapfu=3.42, XMg=0.53) and jadeite (jadeite content Xjd=0.95,
XNa=0.96, XMg=0.39) compose ∼80 vol. % of
the solids at 350 ∘C and 1.30 GPa. Glaucophane and lawsonite are
stable up to 460 and 560 ∘C respectively. Jadeite
abundance increases from ∼10 vol. % at 440 ∘C to
∼16 vol. % at 460 ∘C; ankerite is stable in minor
amounts (≤3 vol. %) at 440–560 ∘C. Garnet (Xalm=0.57, Xgrs=0.41) is stable only at 540–580 ∘C and
2.12–2.24 GPa, reaching ∼5 vol. %, and is then preserved
because of the assumption of fractionation from the bulk in the model. The
metasediment(t) at T<500∘C shows a paragenesis of
phengite (Siapfu=3.68–3.55, XMg=0.53–0.44),
glaucophane, lawsonite, quartz and omphacite (XNa=0.50–0.48,
XMg=0.87–0.70), with minor titanite. At 520 ∘C and
2.06 GPa, lawsonite is consumed and the amphibole proportion reduces from
∼30 vol. % to < 20 vol. %, while garnet (Xalm=0.63, Xgrs=0.34) is produced and reaches ∼10 vol. %. At these conditions, the omphacite content decreases and a
clinopyroxene of more jadeitic composition (Xjd=0.72, XNa=0.76, XMg=0.45) becomes stable. These models are in line with the
first-order mineralogical changes observed in subducted (and exhumed)
crustal material. Thermodynamic calculations predict the coexistence of a
calcic amphibole and a sodic amphibole in the metabasalts and of jadeite and
omphacite in the metasediment(t). From an oxygen isotope partitioning
perspective, the interpretation of the modelled coexistence of a sodic and a
calcic amphibole either as two endmembers of a solid solution or as
coexisting minerals is equivalent. The same applies to pyroxenes, for which
the modelled coexistence of two pyroxenes can be interpreted as a continuous
solid solution. Therefore, this does not affect the oxygen isotope partition
model final results for the other phases and the bulk.
Production of aqueous fluid
At the initial conditions, all the lithologies are saturated in H2O
(Table 1). Up to 500 ∘C, lawsonite, actinolite and glaucophane
are the main repositories of H2O in the metabasalts, followed by
chlorite and minor phengite. A significant pulse of water is modelled at 500–520 ∘C and 2.00–2.06 GPa for both the metabasalts (Fig. 3a, b). This pulse is caused by a decreasing abundance of lawsonite and
amphibole and a breakdown of chlorite followed by growth of garnet and omphacite. This
first dehydration stage releases ∼25 % of the total water
loss from the metabasalt(a) (∼4.0 vol. % H2O
liberated) and ∼45 % from the metabasalt(h)
(∼6.5 vol. % H2O liberated). The second significant
pulse in the metabasalts occurs at 620–640 ∘C and 2.36–2.42 GPa, releasing ∼40 % of the total water loss from the
metabasalt(a) and ∼15 % from the metabasalt(h).
This pulse is caused by the final breakdown of lawsonite and of amphibole in
the metabasalt(a) (Fig. 3b). In the metabasalt(h), glaucophane and
actinolite breakdown takes place at 600 ∘C and 2.30 GPa. Growth
of talc only incorporates half of the water released by amphibole breakdown,
causing an intermediate fluid pulse of minor magnitude at these P–T conditions
(Fig. 3a). If an initial undersaturated basaltic composition was considered,
the amount of released fluid due to the breakdown of hydrous phases would be
smaller.
The metasediment(c) is the rock type that dehydrates the least: the two
main pulses of fluid production are at 480 ∘C and 1.92 GPa
(∼0.4 vol. % H2O liberated) and from 540 ∘C
and 2.12 GPa to 560 ∘C and 2.18 GPa (∼1.7 vol. %
H2O liberated), caused by a breakdown of glaucophane and lawsonite
respectively (Fig. 3c). The water produced from these two dehydration stages
represents < 0.02 wt % of the total water released by the system
composed of metabasalt(h) and metasediment(c). In the
metasediment(t), the main fluid pulse occurs at 520 ∘C and
2.06 GPa (∼3.0 vol. % of H2O liberated), caused by the
breakdown of lawsonite and a decrease in glaucophane abundance (Fig. 3d).
The water produced from this dehydration stage represents ∼0.07 wt % of the total water released by the system composed of
metabasalt(a) and metasediment(t).
As specified in Sect. 2.2, the aqueous fluid used for the calculation
considers only the water component (aH2O=1). The release of CO2
occurs during progressive metamorphism of metabasalt(a) and
metasediment(c). CO2 is absent or present in negligible amounts (< 1 mol % of the total fluid) up to 440 ∘C and 1.76 GPa. At higher temperatures, the X(CO2) content increases significantly
up to ∼10 mol % in the fluid released by metabasalt(a)
and ∼30 mol % in the fluid released by metasediment(c) at 700 ∘C and 2.60 GPa. However, the total amount of fluid
produced at these conditions is negligible (< 0.01 vol. %). The
oxygen isotope fractionation database used for this study (Vho et al., 2020)
does not include fractionation data for CO2. Limited calibrations are
available for oxygen isotope fractionation between H2O and CO2
(Friedman and O'Neil, 1977; O'Neil and Adami, 1969; Zheng, 1994). The
fractionation values diverge significantly, for example between -4.9 ‰ and -8.3 ‰ at 450 ∘C and between 1.5 ‰ and -4.4 ‰ at 700 ∘C. At 520 ∘C, the fluid released by
metabasalt(a) contains 2 mol % of CO2 and 6 mol %–7 mol % at 620–640 ∘C. The consideration of CO2 would produce a
negligible shift on the fluid δ18O (0.1 ‰–0.2 ‰ at 520 ∘C and fractionation used). The amount of CO2 in metasediment(c) derived-fluid is 2 mol % at 480 ∘C and 5 mol % at 540 ∘C. The
consideration of CO2 would also produce negligible to minor shifts in
the fluid δ18O (0.1 ‰–0.3 ‰ at 480 ∘C and 0.2 ‰–1.1 ‰ at 620 ∘C,
depending on the fractionation).
Oxygen isotope composition
The largest initial bulk δ18O difference occurs between
metabasalt(h) and metasediment(c) (14.3 ‰,
the relatively water-poor system), while the smallest initial bulk δ18O difference is observed between metabasalt(a) and
metasediment(t) (6.0 ‰, the relatively water-rich
system). In the following, the results are presented in detail for a
selection of two endmember scenarios (Figs. 4, 5): (1) metasediment(c)
associated with metabasalt(h) and (2) metasediment(t) associated with
metabasalt(a) when not specified differently. Other scenarios (i.e.
metabasalt(h) associated with metasediment(t) and metabasalt(a)
associated with metasediment(c)) give intermediate results in terms of
oxygen isotope composition variations as a consequence of fluid–rock
interaction. Further details and the results for the intermediate scenarios
are given in Supplement S3.
Bulk oxygen isotope compositions
For rocks that undergo only dehydration reactions, the starting bulk δ18O evolves as a consequence of garnet and fluid fractionation (Fig. 4).
The bulk δ18O shift related to water fractionation is within
0.2 ‰, while the shift due to garnet fractionation is
within 0.5 ‰ (Fig. 4c, d). Since water has typically
heavier δ18O with respect to the bulk and the garnet a lighter
one, the two effects produce opposite trends. The combination of both
effects results in a shift in the bulk δ18O in the considered
lithologies restricted to < 0.3 ‰. This in turn
leads to negligible (< 0.2 ‰) variations in the
δ18O values of the stable phases.
Calculated bulk and mineral δ18O values
along the geotherm shown in Fig. 2. Bulk δ18O: black solid
line. Hydrous mineral δ18O: coloured dashed lines. Anhydrous
mineral δ18O: coloured solid lines. Released H2O: thick
blue dashed lines. (a, b) Modelled mineral, bulk and released fluid δ18O values for the metabasalt(h) and the metabasalt(a)
considering garnet fractionation and excess fluid loss and in the absence of
external fluid input. (c, d) Quantification of the effects of garnet
fractionation and fluid loss on the bulk δ18O of the
metabasalts. (e, f) Modelled mineral, bulk and released fluid δ18O values for the metasediment(c) and the metasediment(t)
considering garnet fractionation and excess fluid loss and in the absence of
external fluid input. Scale on the top x axis indicates P (GPa). Mineral abbreviations from Whitney and Evans (2010).
In the metasediments, the progressive interaction with the fluid from the
metabasalts causes a decrease in the bulk δ18O (Fig. 5) that is
controlled by the amount and δ18O of the incoming fluid. A
significant decrease starts at 480 ∘C, where the amount of water
released by the metabasalts increases of about 1 order of magnitude from
< 0.05 vol. % to ∼0.3 vol. % due to partial
consumption of amphibole and lawsonite. The maximum shift in bulk δ18O was calculated for the metasediment(c) interacting with the
metabasalt(h)-derived fluid at -12.9 ‰ for the HI case (integrated
F/R ratio = 0.75 kg kg-1), while it is -8.7 ‰ for the
PI case (integrated F/R ratio = 0.38). The shift in the bulk δ18O of the metasediment(t) interacting with the
metabasalt(a)-derived fluid is -1.5 ‰ for the HI case (integrated
F/R ratio = 0.35) and -2.7 ‰ for the PI case (integrated
F/R ratio = 0.18).
Calculated bulk and mineral δ18O values in
the metasediments along the geotherm shown in Fig. 2 in the case of interaction
with the metabasalt-derived fluid. Only bulk, released H2O and
representative mineral δ18O values are shown for clarity.
Garnet fractionation and excess fluid loss are considered. (a) Metasediment(c): evolution of the δ18O values in the case of
NI (continuous lines) and PI (dashed lines). (b) Metasediment(t): evolution
of the δ18O values in the case of NI (continuous lines) and PI (dashed
lines). (c) Metasediment(c): evolution of the δ18O values
in the case of NI (continuous lines) and HI (dashed lines). (d) Metasediment(t): evolution of the δ18O values in the case of
NI (continuous lines) and HI (dashed lines). Scale on the top x axis indicates P (GPa). Mineral abbreviations are from
Whitney and Evans (2010).
Oxygen isotope composition of mineral phases
Since at infinite temperature the fractionation between any two phases
approaches 0 ‰, a general trend in the reduction of oxygen
isotope fractionation between the stable phases with increasing metamorphic
grade is observed in all lithologies and is a result of the temperature
increase (Fig. 4a, b, e, f). As a consequence, mineral phases typically heavier
than the bulk (i.e. quartz and carbonates) become isotopically lighter with
increasing metamorphic conditions, and the mineral phases typically lighter
than the bulk (i.e. rutile, garnet and titanite) become isotopically
heavier. Such variations are limited (i.e. within 1.0 ‰) for most of the phases, with the exception of quartz, calcite and rutile
that may vary up to 3.0 ‰ in response to temperature
variation only in the considered range.
In the case of ingress of the low-δ18O fluid from the
metabasalts in the metasedimentary rocks (PI and HI cases), the mineral δ18O values decrease progressively with respect to the NI case following
the trend of the bulk δ18O (Fig. 5). For instance, in the case of
NI the δ18O of quartz in the metasediment(t) decreases from
19.4 ‰ to 17.3 ‰ (-2.1 ‰) and that of quartz in the metasediment(c) from 28.0 ‰ to 26.5 ‰ (-1.5 ‰) over the total temperature range modelled. In the
PI and HI cases, the δ18O shift in quartz is respectively -3.8 ‰ and -5.0 ‰ in the
metasediment(t) and -10.0 ‰ and -13.9 ‰ in the metasediment(c). Hence, the final quartz
δ18O values (at 700 ∘C, 2.60 GPa) for the HI case in
the metasediment(t) and in the metasediment(c) are respectively
3.0 ‰ and 13.0 ‰ isotopically lower
than the expected values in the case of NI (Fig. 5). The maximum shift in δ18O (i.e. between NI and HI cases) for the other stable phases is within
those values. In the metasediment(t), the δ18O values of
phengite, glaucophane, jadeite, rutile and garnet decrease by 3.0 ‰ and those of omphacite by 2.4 ‰ from the
NI to the HI case. Lawsonite and titanite are not stable after the first
significant input of metabasalt-derived fluid (500 ∘C, 2.00 GPa),
and their δ18O values decrease by 1.1 ‰ and
0.3 ‰ respectively from the NI to the HI case. In the
metasediment(c), the δ18O values of dolomite, jadeite,
phengite, rutile and aragonite decrease by a maximum of 13.1 ‰, of lawsonite and ankerite by a maximum of 10.3 ‰ and 9.5 ‰ respectively. Garnet
crystallizes only between 540 and 580 ∘C and its
δ18O value in the HI case is 5.9 ‰ lower than
the one in the NI case.
Oxygen isotope composition of the fractionated fluids
The δ18O of the fractionated water from each rock type at each
step is in isotopic equilibrium with the stable mineral assemblage at the
given conditions. In the temperature range where most of the fluid is
released (i.e. T>480∘C), the δ18O of
the fluid is 7.0±0.5 ‰ for the
metabasalt(h) and 10.0±0.5 ‰ for the
metabasalt(a). At T>480∘C, the water released
in the NI case by the metasediment(c) has a δ18O= 24.2–25.6 ‰ and that released from the metasediment(t) has a value of 15.4 ‰–16.4 ‰ (Fig. 6). The δ18O of the fluid leaving
the system, e.g. infiltrating an upper layer or the mantle wedge, results
from the mixing of the water released from the metabasalts and the overlying
metasediments, and it derives from the balance between the amount of fluid
released by each rock type and its δ18O value (Fig. 6a, b).
Double plot diagrams showing the oxygen isotope
composition of the released fluids (left axis, coloured lines) and the
amount (in wt % of the total rock) of the total fluid released by the
systems (right axis, grey field) for each interaction case in the absence of
serpentinite-derived fluid input. (a) Modelled fluid δ18O
values and amount for the system metabasalt(h)+ metasediment(c). (b) Modelled fluid δ18O values and amount
for the system metabasalt(a)+ metasediment(t). Dashed lines
show the δ18O values of the fluids released by each rock type and solid lines the δ18O values of the final fluids released by
each system. In the case of HI, all the metabasalt-derived fluid infiltrates the
metasediment. Hence the final fluid released overlaps with the fluid
expelled by the metasediment and only one line is represented (light blue,
marked as HI). Because the δ18O values of the fluids released by
the metabasalts are not affected by the degree of interaction, all three
cases are represented by one line (marked as NI).
In the NI case, the δ18O of water leaving the system is up to 1 ‰ higher than the composition of the fluid released by
the metabasalt because of the minor input from the metasediment at around
500–550 ∘C. The only exceptions are for the interaction
between the metabasalt(a) and the metasediment(t) at T of
∼450 and ∼700∘C. At
these conditions, the δ18O values of the final fluid are up to
5 ‰ higher than the metabasalt(h)-derived fluid
(Fig. 6b). This increase is caused by a predominance of the
metasediment-derived fluid at those conditions; however, the amount of
high-δ18O fluid represents < 0.1 wt % of the rock
column and ∼1 wt % of the total released fluid (Fig. 6b).
In the case of interaction between the metasediments and the
metabasalt-derived fluid, part of or all this fluid reacts with the
overlying metasediment before leaving the system. The final δ18O of the fluid is controlled by the integrated F/R ratio in the
metasediments and their buffering capacity. In the HI case, the δ18O of the released fluid has a dominant sedimentary signature at T
< 500–520 ∘C, before the first fluid pulse from the
metabasalt (14.5 ‰–15.5 ‰ for the metasediment(t)
and 23.0 ‰–24.0 ‰ for the metasediment(c), Fig. 6). The first metabasalt-derived fluid pulse (500–520 ∘C, see
above) causes a drop in the bulk δ18O values of the total
released fluids of 0.7 ‰ for the
metabasalt(a)–metasediment(t) association and of 6.3 ‰ for the metabasalt(h)–metasediment(c)
association. The second metabasalt-derived fluid pulse (620–640 ∘C, see above) causes a second decrease in the δ18O
values of the total released fluid equal to 1.0 ‰ for
both the lithological associations.
Input of serpentinite-derived fluid
Ultramafic rocks tend to undergo episodic dehydration (see above). In the
following, the effects caused by the input of a serpentinite-derived fluid
at the base of the rock column in the case of HI are described (Fig. 7). The input
of the amount of water with a δ18O of 4.5 ‰ (see above), corresponding to the dehydration of 150 m of pure
serpentine at 480 ∘C (2.0 wt % H2O) and 660 ∘C
(6.5 wt % H2O), has a limited impact on the δ18O of the system. It produces a decrease of < 0.2 ‰ in the final bulk δ18O of the
metabasalts, up to ∼1.0 ‰ in the
metasediment(c) (Fig. 7a) and < 0.5 ‰ in
the metasediment(t) with respect to the HI case with no
serpentinite-derived fluid (Fig. 7b). The largest decrease occurs at 660 ∘C, where the second pulse of serpentinite-derived fluid enters the system,
resulting in ca. -0.1 ‰, -0.3 ‰ and
-1.0 ‰ for the metabasalts, metasediment(t) and
metasediment(c)δ18O bulk values respectively (Fig. 7).
Even by increasing the thickness of the serpentinite by a factor of 2, the
variations in bulk δ18O values with respect to the HI case with
no serpentinite-derived fluid are < 1.0 ‰ for
any rock type with the exception of the metasediment(c), for which the
bulk δ18O decreases by 2.0 ‰ with respect
to the HI case with no serpentinite-derived fluid (Supplement S3).
The effect of the serpentinite-derived fluid on the metabasalts remains the
same for any interaction case (NI, PI and HI), while the variations in the δ18O of metasedimentary rocks decreases to zero with decreasing
infiltration of external fluid (Supplement S3).
DiscussionEffect of stable assemblage evolution and phase fractionation on the δ18O
The changes in mineral assemblage, modes and compositions along a prograde
P–T path control (1) the oxygen isotope partitioning between the stable phases
and (2) the amount and δ18O of the fluid released by the
system. At the same time, oxygen isotope fractionation between the stable
phases is controlled by temperature. Thus, the effects of evolving
paragenesis and increasing temperature are systematically overlapping in
nature. In the case of a closed system, the bulk concentrations of 18O
and 16O remain constant and a change in one phase is compensated for exactly by adjustments in other phases (Baumgartner and Valley, 2001; Kohn,
1993). In this situation, major changes in mineral assemblage do not play a
significant role in shifting the δ18O of stable phases: this is
demonstrated by the limited (< 0.5 ‰) shift in
δ18O values of quartz, garnet, phengite, omphacite and rutile
in the metabasalt(h) after (1) the breakdown of amphibole and lawsonite
and (2) the crystallization of talc and kyanite over a narrow temperature
range between 500 and 580 ∘C (Fig. 4a, b). Similar effects can also be
anticipated for rocks with different chemical composition that undergo
major changes in the mineral assemblage (see Supplement S4).
In a closed system evolving at equilibrium, the initial chemical bulk
composition and bulk δ18O do not change along the P–T evolution.
However, in metamorphic rocks, compositional zoning and metamorphic
overgrowths are often preserved in refractory minerals (Lanari and Engi,
2017) indicating that parts of the minerals have become isolated from the
reactive volume of the rock. This scenario is commonly referred to as
“partially re-equilibrated open systems”, because the chemical and the isotopic compositions vary as a consequence
of the fractionation of solid and fluid phases (i.e. garnet fractionation and
excess fluid removal) even in the absence of external fluid input. Phase
fractionation is expected to affect the bulk δ18O as function
of both the amount of fractionated or expelled phases and their isotopic
composition. Fractionation of a phase lighter than the bulk in δ18O leads to an increase in the reactive bulk δ18O value,
while fractionation of an isotopic heavier phase leads to a decrease in the
reactive bulk δ18O value.
The most common example of
fractionating metamorphic mineral is garnet, which systematically records
compositional zonation at low- to medium-grade (Evans, 2004; Giuntoli et
al., 2018; Konrad-Schmolke et al., 2008; Spear, 1988; Tracy, 1982).
Therefore, garnet fractionation was incorporated into the model in order to
better approximate the behaviour of natural systems. Note that this effect
is reduced at higher grade where intra-crystalline diffusion becomes
efficient to partially re-equilibrate garnet (Caddick et al., 2010; Lanari
and Duesterhoeft, 2019). As already documented by Konrad-Schmolke et al. (2008), garnet fractionation controls the extent of the garnet stability
field. Garnet crystallization is not systematically expected to occur near
the peak conditions, if the matrix was strongly depleted due to garnet
fractionation and the volume of garnet remains constant (i.e. for the
metabasalt(a), Fig. 3b). While garnet fractionation is recognized to
significantly affect isopleth thermobarometry and volume fractions (Lanari
and Engi, 2017), its effect on oxygen isotope bulk composition and
partitioning is negligible (< 0.5 ‰) in all the
studied lithologies. In the model, the garnet fraction varies from
∼5 vol. % in the metasediment(c) to ∼35 vol % in the metabasalt(h) (Fig. 3), and its δ18O is 0.8 ‰
to 1.7 ‰ lower than the bulk (Fig. 4a, b, e, f).
Double plot diagrams showing the effect of the fluid deriving from a layer of 150 m of pure
serpentine (see text for details) on the δ18O of the rock types
and of the of the total released H2O (left axis) and on the amount and
distribution of the H2O released by the systems (right axis). All the
values are calculated assuming HI between the metabasalts and the
metasediments. Black dashed lines represent the bulk δ18O of
the different lithologies and the blue lines the δ18O of the
final fluid released by the systems. The final H2O released by each
system is represented with a red line. The amount and distribution of the
final fluids in the case of no serpentinite-derived fluid input (grey fields)
are shown for comparison. (a) Modelled δ18O values and fluid amount for the system metabasalt(h)+ metasediment(c). (b) Modelled δ18O values and fluid amount for the system
metabasalt(a)+ metasediment(t).
Beside
garnet fractionation, dehydration due to hydrous mineral breakdown and
expulsion of excess water may contribute to changing the starting chemical
and isotopic bulk compositions. Baumgartner and Valley (2001) postulated
that the liberation of metamorphic fluids might have a profound effect on
the stable isotope composition of the residual rock. In the present study,
the maximum fluid loss is from the metabasalts that release ∼15 vol. % (∼5 wt %) of H2O with δ18O
values 0.9 ‰ to 1.5 ‰ higher than the bulk rock (Fig. 4a, b) at T≥500∘C. This significant fluid flux produces a
decrease in the bulk δ18O of less than 0.2 ‰ (Fig. 4c, d). Even if more extensive dehydration
occurs, the effect on the bulk δ18O value will be typically
lower than 1.0 ‰. No significant differences in the
effect of stable assemblage evolution and phase fractionation are observed
between the four lithologies. Therefore, the bulk δ18O of a
rock that experienced a succession of dehydration reactions, without
rehydration by external fluids or major compositional changes through
decarbonation or mineral dissolution, is likely to be representative of its
protolith composition. In this regard, integrated thermodynamics and oxygen
isotope modelling represent a key tool for quantifying the potential
effects of different processes and for assessing closed- or open-system
behaviours.
Mineral δ18O zoning as indicator of open-system behaviour
In the last decades, the significant advances of oxygen stable isotope
analyses by SIMS (secondary ion mass spectrometer) have allowed zoned metamorphic
minerals to be analysed in situ with a precision down to 0.2 ‰–0.3 ‰ (2σ) (e.g. Ferry et al., 2014; Martin et al.,
2014; Page et al., 2010). The magnitude of the intra-crystalline δ18O variation in key metamorphic minerals has been widely used to
establish whether a metasomatic stage is related to an internal fluid deriving
from the breakdown of hydrous phases or if it reflects equilibration with an
external fluid of different isotopic composition (e.g. Putlitz et al.,
2000; Errico et al., 2013; Page et al., 2013; Russell et al., 2013; Martin
et al., 2014; Rubatto and Angiboust, 2015; Engi et al., 2018). Understanding
the scale of fluid migration at depth and the magnitude of the interaction
between fluids and minerals is of special interest and can be enhanced by
modelling of such fluid flow and isotopic exchange (Baumgartner and Valley,
2001). The definition of different interaction cases (NI, PI, HI) is useful to represent various
degrees of isotopic exchange between the fluid and the rock. If the flow is
channelled, all the interaction cases can possibly occur in close proximity and the
modelled scenarios can be linked to the evolution of different domains
around the vein (Fig. 8). The flow of a pervasive fluid leads to
the homogenization of the chemical potential of all components, including stable
isotopes (Baumgartner and Valley, 2001), and it is represented by the HI case,
as long as integrated F/R ratios are high. In contrast, the flow of a
channelled fluid results in local chemical heterogeneities, allowing some
portions of the rock and of the fluid to remain unaffected (NI case) and some
others to be only partially affected (PI case).
Schematic section of a channelled fluid flow where
different degrees of exchange between the fluid and the rock may occur in
spatial proximity. From the host rock perspective, the NI case describes the
distal portion of the rock walls where no fluid infiltrates, the PI case the
intermediate portion where a limited amount of external fluid is available
and the HI case the pervasively infiltrated rock proximal to the vein. From
the fluid perspective, the NI case describes the fluid flow in the centre of
the channel, for which the exchange with the rock walls is negligible; the
PI case the situation in which part of the fluid does not react with the wall rock
and part equilibrates with it, and the HI case the situation in which no fluid
flow occurs without equilibration with the host rock.
The first step for a meaningful interpretation of an observed intra-grain
variation in δ18O value is the quantification of the possible
effects of changes in T and mineral assemblage. Such effects are
characteristic of each phase (Fig. 4a, b, e, f). Quartz, calcite and rutile are
the minerals most sensitive to temperature changes. Their composition is
expected to vary up to 1 ‰ per 100 ∘C, and
they are stable over a wide range of temperature. For such phases, care is
required in interpreting significant intra-grain δ18O
variations (i.e. up to 3.0 ‰) since it does not
necessarily reflect interaction with an external fluid having a different
isotopic composition. However, quartz and calcite are not necessarily robust
minerals that can preserve chemical and isotopic zoning during metamorphism.
On the other hand, the variation in δ18O value over 150 ∘C in a mineral such as garnet that commonly retains growth
zoning is typically within 0.5 ‰ when no phase
fractionation is involved and still less than 1.0 ‰ when
considering the effect of previous garnet and/or excess fluid fractionation
(Fig. 4a, b, e, f). Thus, any larger variation has to be linked to a
significant change in bulk δ18O. Similar behaviour is observed
for other key metamorphic minerals such as phengite, amphibole and
clinopyroxene. These minerals have been widely used in metamorphic petrology
as thermometers and geobarometers (Dubacq et al., 2010; Ferry and Spear,
1978; Parra et al., 2002) and are expected to be robust targets to link the
fluid evolution along the P–T path, especially when mineral relics are
preserved. Due to its large capacity to preserve growth chemistry, garnet
has been a primary target for microscale measurement of oxygen isotopes.
Protocols and reference materials for SIMS measurements for a range of
garnet compositions are well established (e.g. Martin et al., 2014; Page et
al., 2010; Vielzeuf et al., 2005b) and its retentivity to high T resetting by
diffusion has been investigated (Higashino et al., 2018; Vielzeuf et al.,
2005a). In HP rocks, various degrees of intra-grain variations in garnet
δ18O associated with external fluid infiltration have been
reported in the literature. Where field constraints on the fluid source (i.e.
oxygen isotope measurements on the feasible source lithologies) are
available, the intra-grain variation can assist in the calculation of F/R ratios. Martin et al. (2014) describe a shift of -2.5 ‰
associated with infiltration of serpentinite- or altered gabbro-derived fluids
in metasediments from the continental basement in Corsica. Rubatto and
Angiboust (2015) observed a shift of 3.5 ‰–4.0 ‰ in
garnet from an eclogite from Monviso that they attributed to
sediment-derived fluid infiltration. Vielzeuf et al. (2005b) measured a
decrease of 2.5 ‰–3.0 ‰ between garnet core and rim in
the Dora Maira gneiss probably caused by the interaction with a fluid
derived from the whiteschists (Gauthiez-Putallaz et al., 2016). Studies that
used in situ measurement of micas are limited (Bulle et al., 2019; Siron et
al., 2017), and thus the potential of mica to trace fluid–rock interaction is
still underexplored. The matrix complexity of pyroxenes and amphiboles
remains a challenge for SIMS measurements.
Interaction with fluids from an ultramafic source
The effect of pervasive fluid flow deriving from the breakdown of serpentine
(δ18O= 4.5 ‰) in a serpentine layer of
different thicknesses (150, 300, 600 m, see above) on the bulk δ18O of the two metabasalts is negligible (< 0.5 ‰, Fig. 7, Supplement S3). This is mainly
due to (1) the minor difference in δ18O between the
serpentinite-derived fluid and the metabasalts (1.2 ‰
for the metabasalt(h) and 4.5 ‰ for the
metabasalt(a)) and (2) the very low integrated F/R ratio (0.01, 0.02
and 0.04 for the three cases). In this case, the integrated F/R ratio seems to
play a bigger role, since the limited change in δ18O is similar
for both the metabasalts even if the initial Δ18O between fluid
and rock is larger for the metabasalt(a). With the same total volume of
fluid and rock, a channelled fluid flow would imply larger volumes of fluid
interacting with smaller volumes of rock (higher F/R ratios) and would thus
be expected to drive larger variations in isotopic composition. For
instance, by increasing the F/R ratio by a factor of 10 (from 0.01 to 0.1)
the bulk δ18O decreases by 0.6 ‰
(metabasalt(h)) and 1.1 ‰ (metabasalt(a)) upon
infiltration of the serpentine-derived fluids.
In contrast to the metabasaltic layers, the overlying metasediments have (1) larger compositional differences with the serpentinite-derived fluid and (2) a smaller mass. The effect of the serpentinite-derived fluid input at the
base of the column on the metasediment bulk δ18O compositions
can be up to 10 times larger than the effect that the same amount of fluid
has on basaltic compositions. This is the case even when the
serpentinite-derived fluid completely equilibrates with the metabasalt
before interacting with the overlying metasediment. In our model, the fluid
infiltrating the metasediment is always a mixture of metabasalt-derived and
serpentinite-derived fluid. Considering instead an input of
serpentinite-derived fluid directly in the metasediment and applying an F/R ratio of 0.1, the final bulk δ18O of the metasediment(c)
decreases by 4.6 ‰ and that of the metasediment(t) by
2.6 ‰. These values are significant and are comparable
with variations observed in natural samples. For instance, Martin et al. (2014) describe a shift in δ18O of -2.5 ‰
among different generations of HP garnet in a sample from the Corsica
continental basement (garnet mantle δ18O=7.2±0.4 ‰, garnet rim δ18O=4.7±0.5 ‰). The authors associate this shift with an infiltration
of serpentinite-derived fluids and, to a lesser extent, altered
gabbro-derived fluid. Williams (2019) describe an extreme δ18O
shift of -15 ‰ between garnet core and rim in a
metasediment from the Lago di Cignana Unit. Such an oxygen isotope
composition variation has been related to a strongly channelled fluid
influx originating from the dehydration of serpentinites. These results
demonstrate that metasediments can be a good target to detect fluids from an
ultramafic source migrating upward through the subducting slab or along the
subduction interface, even though the two lithologies may not be in direct
contact in the field.
It is important to note that the relatively water-poor system composed of
the metabasalt(h) and the metasediment(c) is more sensitive to
external fluid infiltration and thus affected by the highest changes in
δ18O, according to observations in natural systems (e.g. Page
et al., 2019).
Effect of the subduction geotherm
As discussed in detail by previous studies (Baxter and Caddick, 2013;
Hacker, 2008; Hernández-Uribe and Palin, 2019; Syracuse et al., 2010),
the subduction geotherm has an important effect on hydrous phase stability
and P–T conditions of fluid release into the mantle wedge. Along the average
D80 geotherm by Syracuse et al. (2010) (Fig. 2), the top of the slab crust
releases ∼95 % of the water at ∼80 km,
during the transition from partial to full coupling. Along the cold geotherm
by Penniston-Dorland et al. (2015), the first significant fluid pulse (20 %–40 % of water released) occurs as a consequence of the breakdown of
glaucophane and actinolite at greater depths than predicted by our model
(∼500∘C and 2.7–2.8 GPa, 90–100 km depth),
and the remaining water is released at depth > 100 km. Both those
models imply a relatively dry mantle forearc region, contradicting what has
been described by Bostock et al. (2002). By contrast, along the warm
geotherm by Penniston-Dorland et al. (2015), the breakdown of chlorite,
epidote and actinolite releases 40 %–50 % of the water at 460–470 ∘C and ∼0.6 GPa (∼20 km depth);
after this stage, ∼95 % of the water is released at depth
< 70 km. This implies that little to no water is available at subarc
depths. The average geotherm by Penniston-Dorland et al. (2015) is hotter
than the one by Gerya et al. (2002) used in our model for T<650∘C and P<2.5 GPa (Fig. 2). Along this P–T gradient,
chlorite is the main water carrier in low-grade conditions and the first two
main dehydration pulses occur at depths of 30–40 km (significant decrease in chlorite, 20 %–25 % of water released) and of ∼50 km
(complete breakdown of chlorite, 20 %–25 % of water released). Along
this geotherm, the most water is released at shallower depths than in our
model. Differences can be investigated in detail by modelling each case with
PTLoop. Nevertheless, the effects of fluid–rock interaction on the
bulk and mineral δ18O compositions follow the general trends
described above. Different parageneses are expected to form during
hydration, but the shift in bulk δ18O remains constrained by
the F/R ratio and the isotopic composition of the incoming fluid.
Implication for mantle wedge hydration
Infiltration of the slab-derived fluid into the mantle wedge is important
for subduction zone settings because mantle minerals are strongly depleted
in volatiles. At equilibrium, a free aqueous fluid is not stable in the
mantle wedge at T<650∘C until a fully hydrated mineral
assemblage has formed (i.e. serpentine, chlorite, talc, and amphibole)
(Manning, 2004). As shown above, the P–T conditions of H2O release from
the subducting slab, as well as the volume and the δ18O of the
liberated H2O, can vary according to the geometry of the subduction zone
and the composition of the subducting lithosphere (e.g. Hacker, 2008;
Hernández-Uribe and Palin, 2019; Poli and Schmidt, 2002). The program
PTLoop allows the P–T conditions, amount and oxygen isotope
composition of the released fluid to be calculated for the system of
interest. The presented model uses an intermediate P–T gradient that stabilizes
lawsonite and results in a first significant fluid pulse at 65–70 km
depth and a second pulse at ∼80 km depth. Below that depth,
phengite is the main carrier of H2O. This implies that the majority of
fluid is released in the forearc region, in agreement with previous studies
investigating the dehydration of basaltic and sedimentary components of the
slab (e.g. Baxter and Caddick, 2013; Kerrick and Connolly, 2001; Schmidt
and Poli, 1994).
The influence of the slab-derived fluid on (1) the degree of hydration and
(2) the δ18O modification of the overlying mantle rocks was
estimated on the basis of the results presented above. A slab composed by
the metabasalt(a) and the metasediment(t) (left column in Fig. 1a,
assuming the real thickness of the slab to be 3 times the modelled one)
subducting at 1 cm yr-1 was considered (Fig. 9a). This subduction rate
represents a conservative estimate, considering that previous averages have
proposed 4–5 cm yr-1 (Stern, 2002). Any faster subducting rate may shorten
the timescale of the processes discussed below, allowing larger fluid fluxes
in the mantle wedge over the same interval. Mechanical decoupling between
the slab and the wedge and a steady-state cold mantle wedge are assumed
(e.g. Abers et al., 2006; Hirauchi and Katayama, 2013; Wada et al., 2008).
The fluid released by the slab at 500–520 ∘C with a
characteristic δ18O= 15.0 ‰ (Fig. 6a)
was allowed to infiltrate into an initially dry peridotite (composition KLB-1 from
Walter, 1998, simplified to the Fe–Mg–Al–Si system; Table S2.4 in Supplement S2) with T=570∘C and an initial δ18O= 5.5 ‰ (Eiler et al., 1997; Mattey et al., 1994).
This simplified model ignores the dynamics of fluid infiltration and assumes
pervasive flow. Given the column length of 1 m, a subduction rate of 1 cm yr-1
implies that, in 100 years, any fixed point (i.e. fixed P–T conditions) at the slab–mantle
interface receives the total amount of fluid that a single
column can liberate in those conditions. Hence, in this example 4892.6 kg of
water per 100 years (i.e. the amount released by the considered column at the
chosen conditions, see above) infiltrate the mantle wedge. The released
fluid first interacts with a small volume at the slab–mantle interface (V1 1000×1000×1 m; Fig. 9b) and once it has equilibrated and saturated V1,
it infiltrates volume V2 (3000×2000×1 m; Fig. 9b). Both the volumes were
scaled to 1:3 for the calculation in order to maintain the volume proportion
with the modelled slab. The slab-derived fluid, which is in continuous
supply during subduction of new material, infiltrates V1 and causes a
progressive change in mineralogy from olivine + orthopyroxene + garnet
to serpentine + chlorite + minor olivine until the rock reaches
saturation after 0.35 Myr of subduction. At this stage, V1 has a bulk
δ18O of ∼8 ‰ that is
significantly higher with respect to the initial mantle value (Fig. 9c).
With ongoing subduction, the continuing release of water from new slab
material under a static mantle drives the δ18O of the volume of
mantle wedge toward higher values. The water that will infiltrate V2 has a
δ18O that depends on (1) the δ18O of the
slab-derived water and (2) the buffering capacity of V1. The same changes in
mineral assemblage described for V1 occur also in V2, while the bulk δ18O of V2 increases more moderately than the one of V1 and reaches a
δ18O of ∼6 ‰ after 0.75 Myr
of subduction (Supplement S4).
Case model for mantle wedge hydration. (a) Sketch of a
subduction zone (crustal thickness and serpentinized mantle wedge dimensions are from Bostock et al., 2002). The subducting slab
is composed of metabasalt(a) and metasediment(t) as shown in Fig. 1a, left column. (b) Geometry of the model. The blue arrow represents slab
dehydration at 500–520 ∘C. Abbreviations: V1 represents the
volume of mantle rocks at the interface and V2 the surrounding volume (see
text for details). (c) Plot of the bulk δ18O variations of V1
and V2 as a consequence of continuous slab dehydration over 1 Myr.
In the proposed model, most of the fluid is released by the slab at forearc
depths. However, in most subduction zones no melting appears to occur in the
forearc region and the serpentinite acts as the effective H2O absorber
(Iwamori, 1998), recording the possible variation in δ18O
induced by the slab-derived fluid. Progressive oceanward migration of the
slab (“slab rollback”) has been regarded as an important mechanism
acting in most active subduction zones (e.g. Heuret and Lallemand, 2005;
Nakakuki and Mura, 2013). The rollback of the slab results in a lateral
extension of the serpentinized wedge. As a consequence, the melt ascending
below the arc can interact with serpentinized, high-δ18O mantle
portions that were originally part of the forearc mantle and modify its
original isotopic composition. High-δ18O arc lavas and melt in
arc-setting peridotites have been described (e.g. phenocrysts in lavas from
central Kamchatka: olivine δ18O= 5.8 ‰–7.1 ‰ and clinopyroxenes δ18O= 6.2 ‰–7.5 ‰, Dorendorf et al., 2000; New Guinea: silicate glass
inclusions in olivine δ18O= 8.8 ‰–12.2 ‰, clinopyroxenes in metasomatized lehrzolite δ18O= 6.2 ‰–10.3 ‰, Eiler et al., 1998), but
the mechanism of crustal contamination is still debated. Our results support
the model proposed by Auer et al. (2009) that relates such high-δ18O lavas to the interaction between primitive basaltic melts with an
uppermost mantle that was hydrated and enriched as part of the forearc
mantle prior to trench migration.
Model applications and future directions
The presented approach has a broad range of applications for modelling
fluid–rock interaction in different tectonic settings. We have presented
here an example of subducted crust but the same principles apply also for
regional metamorphism or hydrothermal systems. The model also provides new
ways to quantify the degree of interaction of an external fluid within the
same rock unit. We have shown that the observed effect on the δ18O of a rock of channelled vs. pervasive hydration is strictly coupled
with the composition of the fluid source. Nevertheless, important insights
can be given by linking observations of ideal cases with modelling even if
the composition of the infiltrating fluid is not known a priori. For
instance, the oxygen isotope composition of a fluid source can be retrieved
when a variation in δ18O is observed within the same rock type
from the more hydrated to the less hydrated portions, even in the absence of a
clear presence of a vein or vein system. Alternatively, the degree of
equilibration of the host rock around a vein can be calculated when the
isotopic composition of the fluid is known. Therefore, the combination of
mineral-scale in situ oxygen isotope analyses with major and trace element
mapping will provide much more detailed information on quantitative element
mobility during fluid–rock interaction.
Fluids play a central role as a catalyst for chemical reactions in rocks.
Generally re-equilibration reactions occur only in the presence of fluids that
either derive from breakdown of hydrous phases or from external sources
(e.g. Airaghi et al., 2017; Cartwright and Barnicoat, 2003; Engi et al.,
2018; Konrad-Schmolke et al., 2011; Rubatto and Angiboust, 2015). The
program PTLoop calculates at which P–T conditions the breakdown of hydrous
phases occurs, and, consequently, metamorphic reactions and free fluid are
expected. If fluid-driven reactions occurred in a rock at P–T conditions where
no release of internal fluid is expected, the role of an external fluid
should be considered, and its amount and isotopic composition can be
retrieved using the approach outlined in this paper.
Fluid–rock reactions in subducted lithosphere are likely to involve more
complex open-system processes than the dehydration and rehydration
considered for this model, driving silicate and carbonate dissolution,
transport and re-precipitation (e.g. Ague and Nicolescu, 2013; Piccoli et
al., 2016). Carbon release via the dissolution of calcium carbonate has been
recognized to have important implications for CO2 release from
subduction zones, and it is controlled by H2O-rich fluid infiltration
(e.g. Ague and Nicolescu, 2013; Frezzotti et al., 2011; Gorman et al., 2006;
Kerrick and Connolly, 2001). Future models may account for such variations
in reactive major element bulk composition of the rocks along the P–T evolution
as a consequence of mineral net transfer reactions occurring simultaneously
with fluid–rock exchange (Baumgartner and Valley, 2001), in addition to
water liberation and mineral fractionation (see above). This would require
additional considerations on the changes in the fluid isotopic composition
during transfer of solute species through the fluid. The implementation of other
fluid species beside H2O, such as CO2, could be assessed, provided
that (1) reliable constraints on the oxygen isotope fractionation between
these species and water or minerals are determined and (2) their consistency
with other available data is established. Other species such as CH4 and
H2 do not contain any oxygen, and thus they are likely to be less
relevant to this model.
Conclusions
We developed a user-friendly tool that combines equilibrium thermodynamic
with oxygen isotope fractionation modelling for investigating the
interaction between fluids and minerals in rocks during their metamorphic
evolution. The program simulates along any given P–T path the stable mineral
assemblages, bulk δ18O and δ18O of stable phases
and the amount and oxygen isotope composition of the fluid released.
The capabilities of the program PTLoop are illustrated by an
application to subduction zones, but the presented modelling strategy can be
applied to various metamorphic and tectonic settings. In this study, the
chosen system represents a section of subducting oceanic crust composed by a
lower layer of metabasalt and an upper layer of metasediments of
carbonaceous or pelitic composition. The calculation follows a step-wise
procedure along the chosen P–T path. During the prograde evolution, any mineral
and excess fluid can be fractionated from the reactive bulk composition.
Mineral fractionation and/or excess fluid loss produce only minor (i.e.
< 1.0 ‰) shifts in the bulk δ18O
of any lithology. Hence, the bulk δ18O of a rock that
experienced a succession of such processes without interaction with external
fluids is likely to be representative of its protolith composition.
Variations in δ18O of stable phases due to mineral
fractionation and/or excess fluid loss are also negligible (i.e.
< 0.5 ‰), while the effect of temperature
variation over a range of ∼150∘C on the mineral
δ18O is phase dependent and may be significant (> 1.0 ‰).
Interaction with an external fluid of different
oxygen isotope composition leads to shifts in bulk and mineral δ18O values according to the degree of fluid–rock interaction and
δ18O difference between the rock and the fluid. Extremely large
variations in bulk δ18O of ∼12 ‰ are calculated for the carbonate metasediment
equilibrating with a fluid derived from a metabasalt with an initial
hydrated MORB composition, while small variations of ∼3 ‰ are calculated for the terrigenous metasediment
equilibrating with a fluid from a metabasalt that derives from an altered
oceanic crust. When 50 % or more of the fluid deriving from dehydration
of the metabasalts equilibrates with any of the overlying metasediments, the
final δ18O of the fluid released by the system has a dominant
sedimentary signature, with values between 12 ‰ and 18 ‰.
Such fluids have δ18O values significantly higher than the
mantle value (5.5 ‰) and have a great potential to
modify the oxygen isotope composition of the mantle wedge at the slab–mantle
interface. Extensive serpentinization and a δ18O increase of
∼2.5 ‰ are modelled at the interface
already after 0.35 Myr of ongoing subduction.
PTLoop provides a powerful way to evaluate the effect of closed-system vs. open-system behaviour with respect to oxygen isotopes during the
evolution of the rocks. Different degrees of interaction between the
external fluids and the sink lithology can be simulated and the effects of
internally vs. externally buffered fluids on the mineral paragenesis and on the
mineral isotopic composition investigated.
Measured oxygen isotope compositions in minerals, intra-grain or bulk
δ18O variations at different scales can be compared with the
results of the model for specific scenarios. If the measured isotopic
compositions are not consistent with the behaviour of a closed system, the
presented approach can be used to determine feasible external fluid sources,
to estimate the degree of fluid–rock interaction and the metamorphic
conditions at which this happened. This modelling strategy can also assist
in retrieving the oxygen isotope composition of a fluid source when a
variation in δ18O is observed within the same rock type from
the more hydrated to the less hydrated portions, even in the absence of a clear
presence of a vein or vein system. Our model thus opens new avenues for
mapping fluid pathways related to external fluid infiltration during the
metamorphic evolution of the crust, with important consequences for element
recycling in subduction zones and the investigation of fluid-induced
earthquakes.
Code availability
A compiled version of the program PTLoop, the oxygen isotope fractionation
database and the thermodynamic database used for this study are available at
http://oxygen.petrochronology.org (Lanari and Vho, 2020).
The supplement related to this article is available online at: https://doi.org/10.5194/se-11-307-2020-supplement.
Author contributions
AV developed the model and performed the calculations. PL
supervised the software development and contributed to the code
implementation. DR and JH contributed to
formulate and design the model and to the interpretation of the results.
AV prepared the paper with contributions from all co-authors.
DR conceived the project and secured funding.
Competing interests
The authors declare that they have no conflict of interest.
Special issue statement
This article is part of the special issue “Exploring new frontiers in fluids processes in subduction zones”. It is a result of the EGU Galileo conference “Exploring new frontiers in fluids processes in subduction zones”, Leibnitz, Austria, 24–29 June 2018.
Acknowledgements
We thank the conveners and the participants of the EGU Galileo conference
“Exploring new frontiers in fluid processes in subduction zones” for
constructive discussions. We also thank Ulrich Linden for the technical
support and the reviewers Ralf Halama and Alberto Vitale Bovarone for the constructive comments.
Financial support
This research has been supported by the Swiss National Science Foundation to Daniela Rubatto (grant no. 200021_166280) and Jörg Hermann (grant no. 200021_169062).
Review statement
This paper was edited by Nadia Malaspina and reviewed by Ralf Halama and Alberto Vitale Brovarone.
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