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Theses and Dissertations
2013
Carbon Retention in Deeply Subducted
Sedimentary Rocks: Evidence from HP/UHP
Metamorphic Suites in the Italian Alps
Jennie Cook-Kollars
Lehigh University
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Cook-Kollars, Jennie, "Carbon Retention in Deeply Subducted Sedimentary Rocks: Evidence from HP/UHP Metamorphic Suites in
the Italian Alps" (2013). Theses and Dissertations. Paper 1462.
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TITLE PAGE
Carbon Retention in Deeply Subducted Sedimentary Rocks: Evidence from HP/UHP
Metamorphic Suites in the Italian Alps
by
Jennie Cook-Kollars
A Thesis
Presented to the Graduate and Research Committee
of Lehigh University
in Candidacy for the Degree of
Master of Science
in
Earth and Environmental Sciences
Lehigh University
May 2013
COPYRIGHT PAGE
Copyright by Jennie Cook-Kollars
2013
ii
SIGNATURES PAGE
This thesis is accepted and approved in partial fulfillment of the requirements for the
Master of Science.
_______________________
Date
__________________________________________
Gray Bebout, Ph.D., Thesis Advisor
__________________________________________
Steve Peters, Ph.D., Committee Member
__________________________________________
Bruce Idleman, Ph.D., Committee Member
__________________________________________
Frank Pazzaglia, Ph.D., Chairperson of Department
iii
ACKNOWLEDGEMENTS
To:
Gray Bebout, for many hours of editing and patience with a chemist
Bruce Idleman, for keeping everything running
Steve Peters, for clarifying my research question
Nathan Collins, for heavy lifting and answering everything and sundry
Samuel Angiboust and Philippe Agard, for sharing their extensive knowledge of
Alpine geology
Funding for this research came largely from National Science Foundation grant
EAR-1119264 (to G. Bebout), with smaller amounts of support coming from the
Lehigh University Department of Earth and Environmental Sciences and the
College of Arts and Sciences. A pilot dataset for this study was generated by the
students involved in a group project for course credit at Lehigh University (see
Bebout et al., 2010).
iv
TABLE OF CONTENTS
Title Page ...................................................................................................................... i
Copyright Page ............................................................................................................ ii
Signatures Page .......................................................................................................... iii
Acknowledgements .................................................................................................... iv
Table of Contents .........................................................................................................v
List of Figures............................................................................................................. vi
Abstract.........................................................................................................................1
Context of This Work in Larger C Subduction Project ................................................3
1.
Introduction ...........................................................................................................3
2.
Geologic Setting ..................................................................................................11
3.
Stable Isotope Methods .......................................................................................14
4.
Results .................................................................................................................16
4.1 Petrography .......................................................................................................16
4.2 Stable Isotope Analyses ....................................................................................18
5.
Possible C Loss in the W. Alps Metasedimentary Section .................................23
5.1 Possible Effects of Devolatilization on Isotope Compositions .........................23
5.2 Open vs. Closed System Decarbonation ...........................................................26
5.3 Dissolution of Carbonate as a Mechanism for Mobilizing CO2 .......................28
6.
Closed-System Isotopic Exchange Between Mineral Phases .............................29
6.1 Carbon ...............................................................................................................29
6.2 Oxygen ..............................................................................................................32
6.3 Isotopic Exchange Among Adjacent Layers ....................................................32
7.
Open-System Behavior to Explain the O Isotope Systematics ...........................34
7.1 Fluid Characteristics .........................................................................................35
7.2 Timing of Fluid Infiltration...............................................................................35
7.3 Consistency with Thermodynamic Calculations of Devolatilization in
Metapelites and Metacarbonates .............................................................................37
7.4 Comparison with C Isotope Behavior ...............................................................39
8.
Implications for Carbon Cycling Models ............................................................39
9. Conclusions .........................................................................................................42
References ..................................................................................................................44
Appendices .................................................................................................................52
Vita .............................................................................................................................74
v
LIST OF FIGURES
Figure 1. Cartoon of deep Earth C cycle, showing current estimates of C fluxes in
blue and reservoir volumes in red (in moles of CO2). Fluxes shown suggest a
return efficiency of 23-100% (and possibly higher, although that is less
plausible). After Dasgupta and Hirschmann (2010) with flux updates from
Dasgupta (2013). ...................................................................................................6
Figure 2. Simplified geologic map of the Western Alps, showing location of Lago di
Cignana (enlarged in C), Zermatt-Saas sampling localities, and Cottian Alps
Schistes Lustres traverse (box in A; enlarged in B), in addition to relevant
lithologies. Inset (B) shows P-T data from Agard et al. (2001a) and Angiboust et
al. (2009). Fr: Fraiteve, As: Assietta, Alb: Albergian, Fin:Finestre, Z-S: ZermattSaas, Cig: Lago di Cignana. Figure after Bebout et al. (2013). See detailed
sampling localities in Appendices 3 and 4. .........................................................11
Figure 3. Representative field photos and photomicrographs (long dimension = .164
mm, carb: carbonate layer, pel: pelitic layer, gt: garnet, clz: clinozoisite). (A)
Mostly pelitic outcrop at Finestre locality with thin carbonate layers. Lens cap
for scale in A and B. (B) Carbonate-rich outcrop at Lago di Cignana locality,
showing cm-scale interlayering. (C) Lower grade pelitic sample (SL99-23A,
Albergian). Contains fine-grained carbonate, quartz, organic matter, chlorite,
and oxides. (D) Lower grade carbonate-rich sample (SL99-19D, Albergian).
Contains white mica, quartz, organic matter, and trace biotite. (E) Siliceous
marble from Zermatt-Saas (CV12-3J), consistent with high-grade
metamorphism. Contains coarser-grained garnet, carbonate, white mica, and
clinozoisite. (F) Metapelite from Lago di Cignana (02-LDCS-6), showing major
recrystallization, quartz zoning, and high-grade mineral assemblage. Contains
garnet, quartz, white mica, rutile, and oxides. ....................................................15
Figure 4. Percent of samples from each grade containing various mineral phases.
Although most samples contain quartz and carbonate, compositions range from
quartz-dominated metapelites to marbles. Note increase in calc-silicate phases
between Albergian and Finestre. For P-T estimates see Fig. 2B, and for modal
abundance information see Appendix 1. .............................................................19
Figure 5. Stable isotope compositions of carbonate in the W. Alps metasedimentary
suite. Data are ordered by increasing grade (consistent with other figures). For
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the Schistes Lustres localities, samples are ordered west to east (which
corresponds to grade as a general rule, see Fig. 2B). For the Zermatt-Saas and
Cignana localities samples are not organized within grades. (A) δ13CVPDB of
carbonate in samples, ordered by increasing grade and shaded/sized according to
ratio of carbonate:reduced C weight percents. In general, samples with smaller
proportional amounts of carbonate exhibit larger decreases from standard marine
values (-1.5‰ to +1‰ VPDB). (B) δ18OVSMOW of carbonate in samples, ordered
by increasing grade. Values are uniformly lower than values for marine
carbonate (for the latter, +28 to +30‰) with slightly lighter values at the
Zermatt-Saas and Lago di Cignana localities consistent with greater exchange
with silicates. Shaded/sized according to ratio of carbonate:silicate weight
percents. Highest-grade samples may show a slight correlation with ratio.
Otherwise, shifts are not seen to correlate with carbonate weight percent, silicate
weight percent, or δ13C........................................................................................21
Figure 6. (A) δ13CVPDB of organic matter, ordered by grade as in Figs. 2B, 5 and
shaded according to the ratio of carbonate weight percents to reduced C.
Carbonate-dominated samples exhibit large increase with grade, whereas
samples with lower ratios remain near typical values for marine organic matter
values (-25 to -20‰). (B) Reduced C weight percents also ordered by grade,
showing significant scatter and sharp decrease at Lago di Cignana. ..................22
Figure 7. Demonstration of C isotope shifts resulting from Rayleigh and batch-loss
for (A) Carbonate (assuming CO2 as the fluid C species), and (B) Graphite
(assuming CH4 as the fluid species). Fractionation factors from Chacko et al.
(1991) and Bottinga (1969). ................................................................................26
Figure 8. Carbon isotope reequilibration as a function of metamorphic grade in the
W. Alps metasedimentary suite. Measured Δ13Ccarb-red values are plotted in order
of increasing grade (see Figs. 2B and 5) and colored/sized according to
carbonate:reduced weight percents. Expected equilibrium values from previous
work are shown as dotted lines, sourced from Beaudoin and Therrien (2009).
Note that few samples exhibit complete equilibration. Higher carbonate:reduced
ratios generally correspond to smaller Δ13C values. (Some data could not be
plotted due to very low weight percents, resulting in difficulty determining
isotope ratios for one phase.) ..............................................................................31
Figure 9. Carbonate O isotope data ordered by carbonate weight %, as compared
with modeled whole-rock equilibration after Henry et al. (1996). The low 18O
vii
values for most of the carbonate-rich samples appear to require control by an
alternative reservoir, likely via infiltration by fluids, whereas at low grades, 18O
of the carbonate is consistent with that expected for closed-system
reequilibration with silicate phases. Fractionation factors from Chacko et al.
(1991) and Zheng (1999). ...................................................................................33
Appendix 1. Petrography tables, where x: <10 modal percent, xx: 10-50 modal
percent, and xxx: >50 modal percent. .................................................................52
Appendix 2. Stable isotope data. ................................................................................56
Appendix 3. Sampling localities and sample numbers for Cottian Alps Samples.
Samples from Val d’Aosta region were sampled from single outcrop or streams.
See Appendix 4 for coordinates. Place names ©2013 Esri, DeLorme, NAVTEQ,
TomTom. .............................................................................................................60
Appendix 4. Sample coordinates. ...............................................................................61
Appendix 5. Data from microdrilled traverses, plotted as isotope ratio/carbonate
weight percent compared with distance across layer. Sample photos are also
shown for context. ...............................................................................................62
Appendix 6. Comparison of observed bulk δ13C values with values predicted for a
closed-system model, where isotope ratios are controlled by exchange within the
system rather than by either loss or gain of C during metamorphism (the latter
loss or gain presumably involving metamorphic fluids). Calculated with initial
δ13C carbonate = 0‰ and initial δ13C reduced C = -22.5‰. Line shown is for
predicted = observed bulk-rock values. Many samples could not be shown due
to difficulty determing isotope ratios for samples with very low carbonate
weight percents....................................................................................................73
viii
ABSTRACT
Metamorphism of subducting oceanic crust and carbonate-rich seafloor
sediments likely plays an important regulatory role in the long-term global C cycle
by controlling the fraction of subducting C emitted into the atmosphere via arc
volcanic gases as compared with the fraction of C continuing into the deep mantle.
Metasedimentary suites representing high-pressure (HP)- and ultrahigh-pressure
(UHP)-metamorphosed Jurassic oceanic sediment cover were sampled from
locations across the Italian Alps (the Schistes Lustres, Zermatt-Saas ophiolite, and at
Lago di Cignana). These samples were selected for the broad range of peak P-T (1.53.0 GPa; 330-550°C) conditions that they experienced during subduction (Agard et
al., 2001a; Bebout et al., 2013), corresponding to depths approaching those beneath
volcanic fronts.
Reported in this thesis are field, petrographic, and geochemical evidence for the
efficient retention of C in subducting shale-carbonate sequences through forearcs to
depths approaching those beneath volcanic fronts. This information can inform
models of the global C budget and long-term evolution of the atmosphere. Calcsilicate minerals occur only at the higher grades, in minor abundance, indicating
minor decarbonation in this suite. Carbonate δ13C values retain the signatures typical
of marine carbonates and are relatively uniform across grade, with some shifts to
lower values in low-calcite samples containing abundant reduced C. Carbonaceous
matter shows some increase in δ13C with increasing grade that could reflect
decarbonation, but more likely varying degrees of exchange with carbonate in the
1
same samples. All carbonate analyzed has δ18O significantly lower than is typical for
marine carbonates, partly caused by exchange with silicate phases but, for more
carbonate-rich samples, seemingly requiring some infiltration by externally-derived
fluids. These results indicate that relatively little decarbonation occurred in
carbonate-bearing sediments subducted to depths of up to 90 km, arguing against
extensive decarbonation driven by massive infiltration of the sediments by H2O-rich
fluids released from mafic and ultramafic parts of the underlying slab (as was
modeled in some theoretical studies). Metapelitic rocks intercalated with the
carbonate-rich rocks released some H2O-rich fluid during prograde metamorphism,
providing a more proximal source for fluid to shift the δ18O values of the carbonates.
2
CONTEXT OF THIS WORK IN LARGER C SUBDUCTION PROJECT
This thesis is part of a broader study at Lehigh University on the retention of
carbon in subduction zones and the role of subduction zone metamorphism in global
C cycling, with contributions also from Timm John (Universitat Munster, Munster,
Germany), Samuel Angiboust (GeoForschungsZentrum, Potsdam, Germany), and
Philippe Agard (Universite PM Curie, Paris, France). Behavior of C in subducting
sediments was investigated in this thesis study, and related research on the
underlying mafic lithologies is currently underway, the latter work also based on
exposures in the Italian Alps. Plans are for publication of a short-format paper
including preliminary data for all rock types (and also other suites), with two
separate papers then published containing more in-depth analyses of C residency and
behavior in mafic/ultramafic and sedimentary rocks. This thesis has largely been
sized and formatted for submission to Earth and Planetary Science Letters or
Chemical Geology (both Elsevier journals).
1. Introduction
There has been great recent interest from both the scientific community and the
general public in the short-term (including anthropogenic) controls of atmospheric
CO2 concentrations and global warming. Relatively little attention has been paid to
characterizing the key fluxes affecting longer-term C cycling (at millions of years
timescales), including C subduction and the return of some fraction of this subducted
3
C into the atmosphere in arc volcanic gases (e.g. Sleep and Zahnle, 2001; Tajika and
Matsui, 1992; Zhang and Zindler, 1993). Particularly if significant amounts of
decarbonation occur, forearc to subarc metamorphism of subducted sections of
oceanic lithosphere and sediment could strongly affect the efficiency of this transfer
of C into arc volcanic gases and deeper parts of the mantle. It is generally accepted
that, during subduction, devolatilization and partial melting of the downgoing slab
lead to the release of “fluids” (a term that can refer to relatively dilute hydrous fluid,
silicate melt, or supercritical liquids; see Hermann et al., 2006) from deeply
subducted oceanic lithosphere and overlying sediment. These fluids enter the
overlying mantle wedge, driving partial melting and conveying distinctive “slab
signatures” in arc magmas and arc volcanic volatiles (Elliott, 2003; Hilton et al.,
2002; Morris and Tera, 1989). There is still much debate regarding the amount of
devolatilization and melting that occurs in the forearc and beneath volcanic arcs, and
the roles that this fluid loss could play in the overall input-output mass-balance of C
(see discussion by Dasgupta and Hirschmann, 2010, and update in Dasgupta, 2013).
In models of long-term atmospheric pCO2 (see Berner, 1999, and Berner et al.,
1983), subduction is treated as an integral part of the global C budget. Degassing of
CO2 from the mantle provides the only return of C into the atmosphere, and this CO2
is continually drawn down by silicate weathering and subduction. Caldeira (1991)
proposed that changes between dominantly continental and pelagic carbonate
depositional environments could affect the flux of CO2 to the atmosphere (also see
4
the discussions of sedimentological and tectonic controls by Edmond and Huh, 2003,
and Johnston et al., 2011).
The distribution of C in subducting material is believed to be largely the result of
sedimentation and metasomatic alteration, the latter on the seafloor and perhaps also
in the outer rise region during slab bending. Carbon in slabs is concentrated in
sedimentary cover (as carbonate and as organic C), altered oceanic crust (largely as
carbonate; see Alt and Teagle, 1999), ophicalcites (brecciated serpentinite with
calcite matrix, believed to be more abundant in “Atlantic-type” oceanic crustal
sections), and the associated uppermost mantle (as carbonate). The contribution of
subducting C from carbonated mantle rocks in the subducting oceanic lithosphere is
the least understood of these and Dasgupta and Hirschmann (2010) used the
carbonate concentration in ophicalcites as representing this mantle C flux (see the
updated flux estimate by Dasgupta, 2013). Quantification of the mafic C subduction
inputs is complicated by the great heterogeneity in the altered oceanic crust on the
modern seafloor (see Jarrard, 2003). Despite the relatively small fraction of the
subducting slab section that it represents, sediment cover appears to account for
about one-third of global C subduction (see Bebout, 2007), delivering C into
subduction zones as seafloor carbonate and as organic matter. The composition of
seafloor sediment cover varies widely across the globe, with carbonate contents
ranging from 0-20% and averaging ~0.3% (Plank and Langmuir, 1998). Most
modern carbonate deposition is occurring in the Atlantic and Indian Oceans, whereas
subduction is largely occurring in the Pacific Ocean, leading to the formation of
5
large C sinks in the Atlantic and Indian Oceans (Edmond and Huh, 2003). It follows
that, at this time, only a few subduction margins (Central America, East Sunda,
Vanuatu, Makran) are expected to contain significant amounts of carbonate (Plank
and Langmuir, 1998).
Figure 1. Cartoon of deep Earth C cycle, showing current estimates of C fluxes in blue and
reservoir volumes in red (in moles of CO2). Fluxes shown suggest a return efficiency of
23-100% (and possibly higher, although that is less plausible). After Dasgupta and
Hirschmann (2010) with flux updates from Dasgupta (2013).
Extreme variability in the physical and chemical characteristics of modern
subduction zones has hampered efforts to determine global estimates of subduction C
flux. Jarrard (2003) suggested that 3.47 x 1012 moles of CO2 are subducted every
year in sediments and crust, on a global basis, with 1.2 x 1012 mol CO2/yr attributed
solely to the sediments. Dasgupta (2013) estimated a similar 2.1-6.5 x 1012 mol
CO2/yr for sediment and oceanic crust combined and suggested a further 0.4-0.8 x
6
1012 mol CO2/yr in the upper mantle part of subducting oceanic lithospheric slabs
(see Fig. 1). Bebout (2007) estimated a C input flux of 4.0-8.8 x1012 mol/yr CO2, for
sediment and oceanic crust, overlapping with the estimate of Dasgupta and
Hirschmann (2010) without flux in mantle rocks. Only the estimate by Bebout
(2007) includes organic C in addition to carbonate lithologies. Predicted ranges of
volcanic arc return, on a global basis, are similarly varied but the generally accepted
range is 1.5-3.1 x 1012 mol CO2/yr (Hilton et al., 2002; Marty and Tolstikhin, 1998;
Sano and Williams, 1996; see the somewhat lower flux estimated by Snyder et al.,
2001). Mid-ocean ridge volcanism returns a comparable amount of CO2 into the
oceans and atmosphere (1.0-5.0 x 1012 mol CO2/yr; see Fig. 1). All of these estimates
have significant error bars due to sampling limitations, difficulties in estimating the
amount of offscraping/erosion occurring within subduction zones, and lateral
variability.
In an effort to provide better-constrained subduction input/output comparisons,
additional work has focused on mass-balancing inputs and outputs for individual
subduction zones. By far the most thorough work to date has been done on the
Central American margin. Based on detailed study of several DSDP cores, Li and
Bebout (2005) estimated a total amount of downgoing C of 1.4 x 1011 g/yr, as TOC,
and 1.5 x 1012 g/yr, as oxidized C (in carbonate), over the entire length of the margin.
de Leeuw et al. (2007) analyzed C output flux in Central American volcanoes and,
comparing this flux with the estimate input flux of Li and Bebout (2005), estimated
recycling efficiencies of 12-18% in Costa Rica and 29% in El Salvador. They
7
suggested that all C emissions could be accounted for through decarbonation of
sediment cover, with a 76% contribution from marine carbonate, 14% from organic
matter, and 10% from mantle wedge material. While there has not been such
thorough study of most subduction zones, it is generally agreed that a significant
fraction of the released C comes from sedimentary material (again, perhaps one-third
of the C subduction input on a global basis; Bebout, 2007). As an approximation, the
total oceanic crust+sediment C subduction input inventory is composed of
carbonates and reduced matter in an approximately 4:1 ratio globally (see
calculation, not including subduction in ultramafic rocks, and the related discussion
by Bebout, 2007). However, this ratio varies considerably among modern margins
into which widely varying sediment sections are being subducted.
Current models suggest that subarc slab surface temperatures range from 600900°C, only sufficient to cause dry melting at select few subduction zones and rarely
in the forearc (Hermann et al., 2006; Syracuse et al., 2010). It has been hypothesized
that dehydration of oceanic crust could produce a substantial amount of H2O and
facilitate melting and decarbonation of overlying sedimentary material (Kerrick and
Connolly, 2001b). Significant field evidence exists for substantial fluid flow through
deeply subducted sediments and oceanic crustal rocks, as preserved in veins and
textures in the Dabie Shan region of China, Zambezi Belt in Zambia, and the
Franciscan complex in California (Franz et al., 2001; John and Schenk, 2003; Romer
et al., 2003; Sadofsky and Bebout, 2001). These suites show no field, petrographic,
or geochemical evidence for having been partially melted and most or all of them
8
experienced peak metamorphism firmly within the subsolidus P-T region (see
discussion by Bebout, 2007). Based on these previous results, it is expected that,
within the forearc region, C is likely to be lost via decarbonation and other
devolatilization reactions rather than via wholesale melting.
In this study, I assessed the extents of retention of oxidized and reduced C to
depths approaching those beneath volcanic fronts through detailed field and
petrographic work and analyses of the concentrations and stable isotope
compositions of carbonate and organic/reduced C in HP/UHP metasedimentary rocks
in the Western Alps. My focus was on a particularly well-studied traverse of the
Schistes Lustres in the Cottian Alps, exposures of sediments associated with the
Zermatt-Saas ophiolite, in the Val d’Aosta region, and UHP metapelitic and
metacarbonate rocks exposed at Lago di Cignana (Valtournenche; see Agard et al.,
2001a; Angiboust et al., 2009; Bebout et al., 2013; see Fig. 2). As in Bebout et al.
(2013), which focused on devolatilization history only of the metapelites and on
possible mobility of some key trace elements, this metasedimentary suite was
selected because of the known high degree of preservation of prograde to peak
metamorphic paragenesis due to the isothermal to down-temperature exhumation
history experienced by the rocks. This suite represents an aggregate prograde P-T
path intermediate to the “cool” and “warm” paths calculated for rock subducting into
modern subduction zones (see Syracuse et al., 2010; discussion by Bebout et al.,
2013; Busigny et al., 2003). Petrography and stable isotope analyses were
undertaken to elucidate the processes contributing to C distribution and potential
9
forearc C loss. Studies of this type and magnitude have previously been undertaken
only on carbonate-poor metamorphosed shales and sandstones in the Catalina Schist,
Western Baja Terrane, and Franciscan Complex, together representing subduction to
only 15-45 km (Bebout, 1995; Sadofsky and Bebout, 2001, 2003). There are,
however, other more scattered reports of marble in UHP suites (Becker and Altherr,
1992; Ogasawara et al., 2000; Ohta et al., 2003; Zhang et al., 2003; Zhang and Liou,
1996), leading these authors to suggest that carbonate can be retained to great depths
in subduction zones. These other studies of marbles do not provide stable isotope
datasets of the type generated here allowing consideration of the possible evolution
in isotope composition across grade in a single paleosubduction margin. Work of the
type presented in this thesis can contribute information regarding the efficiencies of
delivery of oxidized and reduced C reservoirs to depths approaching those beneath
volcanic fronts, beyond which some fraction of this C inventory is added to the
mantle wedge (and outgassed in volcanic arcs), or entrained into the deeper mantle to
beyond subarcs. Constraints on the input-output subduction zone mass-balance of C
can inform models in which assumptions are made regarding the arc volcanic return
of C to the atmosphere, contributing to its pCO2 (e.g., Berner, 1999), and are
necessary in evaluating the long-term (and modern) net flux of C among the surface
and deep-Earth reservoirs (see Marty and Tolstikhin, 1998; Zhang and Zindler,
1993).
10
Figure 2. Simplified geologic map of the Western Alps, showing location of Lago di
Cignana (enlarged in C), Zermatt-Saas sampling localities, and Cottian Alps Schistes
Lustres traverse (box in A; enlarged in B), in addition to relevant lithologies. Inset (B)
shows P-T data from Agard et al. (2001a) and Angiboust et al. (2009). Fr: Fraiteve, As:
Assietta, Alb: Albergian, Fin:Finestre, Z-S: Zermatt-Saas, Cig: Lago di Cignana. Figure
after Bebout et al. (2013). See detailed sampling localities in Appendices 3 and 4.
2. Geologic Setting
The variably metamorphosed Schistes Lustres oceanic sediments were deposited
in the Jurassic and Cretaceous, then were subducted in the convergence that
11
ultimately led to continental collision and the formation of the Alps (e.g., Coward
and Dietrich, 1989; de Graciansky et al., 2011; Lardeaux et al., 2006). Beginning at
165 Ma, continental rifting led to the formation of the Valais and Liguro-Piemontais
Oceans (de Graciansky et al., 2011). The Tethyan ocean floor began subducting
under the Apulian plate in the early Eocene, eventually culminating in continental
subduction (45-40 Ma) and formation of an accretionary wedge. Following the
Oligocene continental collision and the complete destruction of the Apulian plate,
the area has since been in a state of minor extension.
The Schistes Lustres section has been shortened during accretionary wedge
formation, resulting in the thick exposures of Upper Mesozoic pelagic
metasedimentary rocks seen today (e.g., Deville et al., 1992; Lemoine and Tricart,
1986). The Cottian Alps traverse (extensively sampled in this study, see Figs. 2 A,B)
has been shown to exhibit a relatively continuous increase in peak metamorphic
pressures and temperatures towards the east, ranging from 1.4 GPa and 330°C at the
Fraiteve exposures in the west to 2.0 GPa and 480°C in the Finestre region (Agard et
al., 2001a; see the inset P-T diagram in Fig. 2B). These exposures consist of
metacarbonates and metapelites interlayered at all scales (but mostly mm to m
scales; see Fig. 3), and samples collected for this study represent the full range of
metasedimentary compositions exposed (i.e., a continuum of mixtures of metapelite
and metacarbonate). Protoliths of these exposures have been compared with the
sediment section subducting into the modern-day subduction zone at the Banda
forearc, sampled at ODP Sites 262 and 765 (Ernst, 2005; Trümpy, 2003; see
12
description of this sediment section in Plank, 2013). In addition to the thorough
geothermobarometry studies (Agard et al., 2001a; Beyssac et al., 2002), some stable
isotope work has previously been done on the Cottian Alps traverse (Bebout et al.,
2013; Busigny et al., 2003; Henry et al., 1996). However, Henry et al. (1996) worked
on only the lowest-grade units in this HP-metamorphosed suite and largely
concentrated on units to the west of our traverse that did not experience high-P/T
metamorphism. A recently published study of the devolatilization history in the
metapelitic rocks in this suite (Bebout et al., 2013; also see Busigny et al., 2003)
focuses on loss of volatiles and mobilization of trace elements during subduction
along the P-T path collectively represented by the various units. For the same Cottian
Alps traverse, Beyssac et al. (2002) investigated the crystallinity of the
metamorphosed organic matter, using Raman spectroscopic and HRTEM methods.
Farther north, in the Val d’Aosta region (see Fig. 2A), the Zermatt-Saas ophiolite
is largely composed of mafic and ultramafic exposures with relatively few associated
metamorphosed cover sediments ranging in composition from metapelites to nearly
pure marbles. The ophiolitic rocks exposed in this area were exhumed in larger slices
and do not show the progression of grades observed for the sediment-dominated
Schistes Lustres exposures such as that seen in the Cottian Alps (Agard et al., 2009).
Sampling in the Val d’Aosta was performed at three locations (Servette Valley,
Clavalité Valley, and near Cervinia) as well as at the associated UHP Lago di
Cignana exposure (in the nearby Valtournenche; Reinecke, 1998). The rocks in the
Val d’Aosta are known to collectively have experienced peak metamorphism at 52013
550°C and 2.3-2.4 GPa (Angiboust et al., 2009). Estimation of peak pressures for the
metasedimentary rocks at the Lago di Cignana exposure have proved more difficult
due to the extensive exhumation-related overprinting of these rocks (see Bebout et
al., 2013, for discussion), but current estimates suggest 550°C and 2.5-3.0 GPa.
However, Frezzotti et al. (2011) reported occurrences of microdiamond in garnet
porphyroblasts in siliceous, Mn-rich schists, indicating the possibility of higher peak
pressures in a nearby small tectonic slice, as was suggested by Reinecke (1991) and
Groppo et al. (2009). Previous work by Cartwright and Barnicoat (1995, 1999) in the
Zermatt-Saas and Täschalp regions focused on stable isotope analyses of the
ophiolitic rocks, but their smaller number of stable isotope data for the sediment
cover largely overlaps with those presented here.
3. Stable Isotope Methods
Calcite C and O isotope values were obtained following the procedure developed
by McCrea (1950) involving formation of CO2 by reaction with phosphoric acid.
Some samples were reacted overnight at 27°C, cryogenically purified on a glass
vacuum extraction line, and then analyzed in dual-inlet mode on a Finnigan MAT
252 mass spectrometer at Lehigh University. Other samples were reacted with 0.3
mL of 100% phosphoric acid for a minimum of thirty minutes (three hours for Lago
di Cignana samples containing minor amounts of dolomite) at 72°C and introduced
into the mass spectrometer by a Gas Bench II coupled with a Combi PAL
autosampler. Regular analysis of a house standard and international standard NBS-19
14
Figure 3. Representative field photos and photomicrographs (long dimension = .164 mm,
carb: carbonate layer, pel: pelitic layer, gt: garnet, clz: clinozoisite). (A) Mostly pelitic
outcrop at Finestre locality with thin carbonate layers. Lens cap for scale in A and B.
(B) Carbonate-rich outcrop at Lago di Cignana locality, showing cm-scale interlayering.
(C) Lower grade pelitic sample (SL99-23A, Albergian). Contains fine-grained
carbonate, quartz, organic matter, chlorite, and oxides. (D) Lower grade carbonate-rich
sample (SL99-19D, Albergian). Contains white mica, quartz, organic matter, and trace
biotite. (E) Siliceous marble from Zermatt-Saas (CV12-3J), consistent with high-grade
metamorphism. Contains coarser-grained garnet, carbonate, white mica, and clinozoisite.
(F) Metapelite from Lago di Cignana (02-LDCS-6), showing major recrystallization,
quartz zoning, and high-grade mineral assemblage. Contains garnet, quartz, white mica,
rutile, and oxides.
15
allowed monitoring and correction of the data, resulting in a standard deviation of
0.2‰. For reduced C analyses, samples were first treated with 1N HCl, to remove all
calcite, then oxidized by combustion in sealed quartz tubes (6 mm o.d.) with CuCuOx reagent before extraction and analysis in dual-inlet mode (see description of
methods by Li and Bebout, 2005). Graphite standard USGS-24 was used to ensure
accuracy, with a long-term standard deviation of <0.1‰. Samples for which reduced
C was measured include the relatively small number of metapelitic rocks studied by
Bebout et al. (2013) for their major and trace element and N isotope compositions.
13C and 18O values are reported relative to the VPDB and VSMOW standards,
respectively.
For most of the analyses, samples were hand-powdered by mortar and pestle to
produce several grams of whole-rock powder. Microdrilled samples (using a 2 mm
drill bit) across metacarbonate layers (2-6 cm) adjacent to metapelite were analyzed
to test for C and O isotope exchange across lithological boundaries. Traverses of 3 or
4 points (depending on sample size) were performed on 11 hand specimens.
4. Results
4.1 Petrography
Fig. 4 and Appendix 1 contain petrographic observations made on the suite of
metasedimentary rocks for which isotopic data are presented in this paper (also see
the summaries of mineralogy of the more metapelitic rocks in Agard et al., 2001b;
Bebout et al., 2013). Figure 3 shows some representative mineralogy and textures at
16
the thin section scale of the metapelitic and more carbonate-rich rocks across grades.
All localities contain a wide variety of metasedimentary rocks ranging from quartzrich metapelitic schist, with negligible carbonate, to marbles with >70% carbonate.
Lowest-grade samples from the Fraiteve locality commonly contain secondary
phases biotite and chlorite, as well as lawsonite and the rare Mg-Fe mineral
carpholite. Calc-silicate phases that could represent decarbonation reactions (e.g.,
lawsonite) are lacking or very minor in modal abundance. Metamorphosed organic
matter is abundant in the strongly foliated metapelites. Assietta samples have
mineralogy and fine-grained textures similar to those at the Fraiteve locality, but
with increased modal abundance of chloritoid and carpholite. Organic matter in these
samples is concentrated largely in the more micaceous domains, and is far less
abundant in the quartz-rich domains. Samples from the Albergian region (slightly off
the strike of the main traverse, the latter which was collected along a small mountain
road; see Agard et al., 2001a; see Appendix 3), are extremely carbonate-rich or are
fine-grained metapelites containing abundant organic matter, but without carpholite
(see Fig. 2B). Finestre samples tend to be coarser-grained than those at the lower
grades and organic matter in these rocks occurs in clusters rather than in wispy,
deformed domains. Most samples from this locality contain calc-silicate phases with
total modal percentage <20% and generally less than <10%; clinozoisite, garnet, and
titanite are the most abundant calc-silicate minerals. Tourmaline abundance in
metapelitic Finestre samples is notably higher than in lower-grade lithologic
equivalents (see Appendix 1; Bebout et al., 2013). The Zermatt-Saas
17
metasedimentary lithologies appear fairly similar to those found at the Finestre
locality (likely due to the similar peak P-T they experienced), and contain significant
amounts of clinozoisite, rutile, garnet, titanite, and tourmaline. The UHP Cignana
metasedimentary rocks contain clinozoisite, garnet, rutile, and titanite with very clear
distinctions between metacarbonates and quartz-rich metapelites (see Fig. 3) and far
greater abundance of the calc-silicate minerals in the more carbonate-rich samples.
Grain sizes are coarser in the metapelites and the metacarbonates at this and the Val
d’Aosta localities, relative to grain sizes in the Cottian Alps, reflecting their
recrystallization at higher peak metamorphic P-T. In summary, the most clear change
in mineralogy, in these samples, is at the boundary between the Albergian and
Finestre localities, corresponding to peak P-T of 1.9-2.0 GPa and 430-480°C, with
the Finestre metacarbonates containing a significantly greater modal abundance of
calc-silicate phases (on average, increasing from <<5% to <10%; Appendix 1).
4.2 Stable Isotope Analyses
Oxygen isotope ratios of carbonate are relatively uniform across grades (average
δ18OVSMOW = +20.1‰, standard deviation = 2.2‰), but show a slight decrease at the
higher grades (average of +20.8‰ at Fraiteve to +18.2‰ at Cignana; see Fig. 5B).
δ13CVPDB of the carbonate is also fairly uniform, with values for most samples falling
within the typical marine carbonate range of -1.5 to +1‰ (Hoefs, 2009; see Fig. 5A).
Some samples containing smaller amounts of carbonate show downward shifts in
δ13C, with values as low as approximately -6‰. As noted earlier, carbonate
18
Figure 4. Percent of samples from each grade containing various mineral phases. Although
most samples contain quartz and carbonate, compositions range from quartz-dominated
metapelites to marbles. Note increase in calc-silicate phases between Albergian and
Finestre. For P-T estimates see Fig. 2B, and for modal abundance information see
Appendix 1.
19
weight percent ranges from 0-100%, with the Schistes Lustres samples averaging
30% carbonate and the higher grade samples averaging 15% carbonate (see
Appendix 2).
The traverses across metacarbonate layers adjacent to metapelites, obtained by
microdrilling, showed little variation in either δ13Ccarbonate or δ18Ocarbonate (in nearly all
cases less than 0.5‰; Appendix 5). As no consistent trend indicative of a spatial
control on isotope exchange was found across layers, the whole-rock analyses can
safely be interpreted as representative values without averaging out interesting
variations within rock layers.
Reduced (organic) C δ13C shows significant variation across grade (see Fig. 6A),
with values at higher grades showing a strong relationship with the ratio of carbonate
C to organic C in the same samples. The values for Fraiteve samples are uniformly
low (mean δ13C = -22.5‰; 1 = 1.4‰), similar to values for unmetamorphosed
marine organic matter (Hoefs, 2009), whereas at higher grades the reduced C shows
increase in δ13C. Values for the Albergian rocks average -18.2‰ and samples from
Finestre show a bimodal distribution with one group averaging -25‰ (values similar
to those of the Fraiteve samples) and the other continuing the trend of increase
initiated at the lower grades and averaging -12‰. Values for Zermatt-Saas samples
average -11.8‰ (see Fig. 6A), similar to the group of higher- δ13C values for the
Finestre samples, whereas the δ13C values for Cignana samples show more even
scatter from -26.0‰ to -8.6‰ (Fig. 6A). Weight percent of reduced C ranges from
20
Figure 5. Stable isotope compositions of carbonate in the W. Alps metasedimentary suite.
Data are ordered by increasing grade (consistent with other figures). For the Schistes
Lustres localities, samples are ordered west to east (which corresponds to grade as a
general rule, see Fig. 2B). For the Zermatt-Saas and Cignana localities samples are not
organized within grades. (A) δ13CVPDB of carbonate in samples, ordered by increasing
grade and shaded/sized according to ratio of carbonate:reduced C weight percents. In
general, samples with smaller proportional amounts of carbonate exhibit larger decreases
from standard marine values (-1.5‰ to +1‰ VPDB). (B) δ18OVSMOW of carbonate in
samples, ordered by increasing grade. Values are uniformly lower than values for marine
carbonate (for the latter, +28 to +30‰) with slightly lighter values at the Zermatt-Saas
and Lago di Cignana localities consistent with greater exchange with silicates.
Shaded/sized according to ratio of carbonate:silicate weight percents. Highest-grade
samples may show a slight correlation with ratio. Otherwise, shifts are not seen to
correlate with carbonate weight percent, silicate weight percent, or δ13C.
21
Figure 6. (A) δ13CVPDB of organic matter, ordered by grade as in Figs. 2B, 5 and shaded
according to the ratio of carbonate weight percents to reduced C. Carbonate-dominated
samples exhibit large increase with grade, whereas samples with lower ratios remain
near typical values for marine organic matter values (-25 to -20‰). (B) Reduced C
weight percents also ordered by grade, showing significant scatter and sharp decrease at
Lago di Cignana.
0.02% to 0.84%, showing an overall average of 0.21% (Fig. 6B, Appendix 2) and,
with the possible exception of Lago di Cignana, no obvious variation across grade
(although such an assessment is highly complicated by the extreme protolith
heterogeneity). These data are in line with previous analyses of wt. % reduced C in
22
these metasedimentary rocks reported by Beyssac et al. (2002) and Henry et al.
(1996).
5. Possible C Loss in the W. Alps Metasedimentary Section
In this section, we discuss the evidence for possible loss of C from this
metasedimentary section at the higher grades using petrographic observations (i.e.,
presence or absence of Ca-rich silicate phases that could reflect decarbonation) and
discussing the possibility that shifts in the C isotope compositions of the highergrades are the effects of isotopic fractionation during loss of C to metamorphic fluids
(as CO2 or CH4). Most of the previous consideration of subduction zone C loss has
focused on decarbonation (loss of volatile CO2 from carbonate minerals) as the most
likely process at currently accepted pressures and temperatures, but other pathways
should be considered as well. In particular, dissolution of carbonate into
metamorphic fluids has been suggested as a more efficient means of removing C
from subducting sedimentary rocks (see Frezzotti et al., 2009). In this section, we
evaluate each of these mechanisms of C loss and the degree to which the W. Alps
metasedimentary suite shows evidence for loss by these mechanisms.
5.1 Possible Effects of Devolatilization on Isotope Compositions
Shifts in isotopic composition could be explained by devolatilization, perhaps
with carbonate-rich rocks largely releasing CO2 and metapelitic rocks rich in reduced
C (and containing little carbonate) releasing CH4. Using accepted fractionation
factors (Chacko et al., 1991), it is evident that, even with 50% decarbonation at
23
550°C, the carbonate δ13C would be shifted by less than 3‰ (see Fig. 7A). This is a
fairly extreme case, and the relative lack of Ca-rich phases obviously resulting from
decarbonation argues against decarbonation as a significant process in these rocks
(see Appendix 1). Decarbonation would also be expected to result in lowered
carbonate wt. % at the higher grades but comparisons of carbonate content across
grade are complicated by the extreme protolith heterogeneity. Work on the largely
metapelitic and metapsammitic Catalina Schist in California documented a
significant decrease in calcite weight percents at the higher grades with loss of small
amounts of diagenetic cement at 1GPa and 300°C (Sadofsky and Bebout, 2003), a
scenario that is certainly not replicated as clearly here. Petrographic work in this
study indicated only minor concentrations (generally 0-10% but ranging up to 20%)
of calc-silicate phases (clinozoisite, garnet, titanite) that could reflect decarbonation,
providing further support for significant C retention in this suite to at least 90 km.
The first (lowest-grade) occurrence of these phases at the Finestre locality indicates
that this minor decarbonation was initiated at about 2.0 GPa and 430 to 480°C.
Although the possibility of decarbonation at all grades cannot be ruled out based on
isotopic evidence, the retention of significant C and lack of calc-silicate phases
suggests that only minor decarbonation could have occurred.
Carbonaceous matter could also undergo devolatilization, depending on oxygen
fugacity producing either CO2 or CH4. Release of CO2 can be ruled out based on
theoretical calculations by Bottinga (1969) that predict a decrease in the δ13C of the
reduced matter, a result inconsistent with the isotope shifts seen here (Fig. 6A). In
24
contrast, conversion of reduced C into CH4, in H2O-rich fluids generated by
devolatilization (see Bebout et al., 2013), would preferentially leave the residual
reduced C enriched in 13C and increase δ13C in rocks experiencing such loss. Bebout
(1995) attributed shifts in the δ13C of reduced C in metasandstones of the Catalina
Schist to speciation of reduced C as isotopically fractionated CH4 in H2O-rich fluids
produced by dehydration reactions. Given the fluid-rock fractionation expected
during this process (103ln = -4‰ according to Bottinga (1969), corresponding to a
temperature of 550˚C; see Fig. 7B), 90% loss of reduced C to CH4 would be required
to explain the shifts of up to approximately 15‰ observed in this traverse (Fig. 6A).
This would require some higher-grade rocks to have contained unreasonably large
amounts of organic matter before loss to metamorphic fluids, unlikely given their
oceanic origin and the reduced C concentrations in their lower-grade equivalents in
this same suite. As for the estimation of carbonate loss due to decarbonation (see
discussion above), it is difficult to estimate the degree of reduced C loss, based on
the weight percent data, due to the heterogeneity of samples and their likely
protoliths (Fig. 6B). While it is not possible to rule out protolith variation, samples
from Lago di Cignana (average 0.09 wt. %) certainly contain smaller amounts of
reduced C than samples from other localities (average 0.24 wt. %), conceivably due
to some degree of loss of reduced C to fluids during devolatilization. Comparison
with whole-rock data could potentially serve to elucidate this distinction in future.
25
Figure 7. Demonstration of C isotope shifts resulting from Rayleigh and batch-loss for (A)
Carbonate (assuming CO2 as the fluid C species), and (B) Graphite (assuming CH4 as the
fluid species). Fractionation factors from Chacko et al. (1991) and Bottinga (1969).
5.2 Open vs. Closed System Decarbonation
It is possible to envision decarbonation occurring in the carbonate-bearing rocks
in this suite in several different fluid-rock scenarios. In the closed system model laid
out by Kerrick and Connolly (2001a), any fluid formed during decarbonation is
removed from the system without further interaction but the rock has no fluid added
to it from external sources. Decarbonation is largely driven by temperature (and
pressure) increase and fluid composition can be controlled internally. The amount of
decarbonation that occurs also depends on the presence of suitable reactants, prior to
reaction, and the relative modal proportions of the reactants. By this model, rocks
very rich in carbonate and poor in silicate phases with which to react would be
26
expected to experience only minimal decarbonation and related isotopic shift in
residual carbonate. Such a model has been invoked in a number of studies of
devolatilization in which it appears that shifts in concentration and isotopic
composition can be explained by a Rayleigh or batch-loss model (e.g., Bebout and
Fogel, 1992, for loss of N as N2 into metamorphic fluids generated during subduction
zone metamorphism of the Catalina Schist). Calculated decarbonation histories for
subducting carbonate-rich sediments in the closed system (Rayleigh-like) model by
Kerrick and Connolly (2001a and updated in Connolly, 2005) predict that much of
the carbonate could be stable to subarc depths and into the deeper mantle.
In contrast with the Rayleigh-like models, models of open system behavior
presented by Gorman et al. (2006) predict significant decarbonation in the forearc. In
these models, H2O-rich fluid from underlying rocks in the slab infiltrate subducting
carbonate-rich rocks, at any given temperature along prograde P-T paths driving
decarbonation reactions and resulting in far greater amounts of reaction. Thus, a far
smaller fraction of the carbonate initially subducted is retained through the forearc
and to depths beneath arcs and beyond. Further variations on this theme could
involve the addition of dissolved material in these infiltrating fluids capable of
reacting with the carbonate in decarbonation reactions (e.g., dissolved SiO2). In some
cases, this could be a smaller-scale process in which the infiltrating H2O-rich fluids
are generated in directly adjacent metapelitic rocks. Infiltration from any external
source can also potentially result in shifts in isotopic composition without any
decarbonation, depending on fluid-rock ratios, if the infiltrating fluids are initially
27
out of isotopic equilibrium with the rocks. Alternatively, large amounts of
isotopically non-reactive fluid could infiltrate rocks such as these without leaving
any isotopic imprint, for example, if they have previously equilibrated isotopically
with similar rocks (see Sadofsky and Bebout, 2001).
The very low modal abundance of calc-silicate phases in even the highest-grade
rocks in this traverse (see Fig. 4, Appendix 1) indicates that these rocks experienced
little decarbonation and related C loss. Significant carbonate is present at all grades
that were studied (up to 70% even at Lago di Cignana), consistent with the more
minor decarbonation expected from calculations by Kerrick and Connolly (2001a).
However, the shifts in δ18O, relative to values for presumed protoliths (see Fig. 5B),
require further explanation if the rocks were for the most part closed systems as they
devolatilized (see Section 6.2, 7).
5.3 Dissolution of Carbonate as a Mechanism for Mobilizing CO2
If deeply subducted carbonate-bearing rocks experience only minimal
decarbonation (as predicted by the models of Kerrick and Connolly, 2001a,b), it
could be necessary to find an alternative mechanism by which CO2 can be released
from subducting sediments to explain the volcanic arc CO2 outputs. Dissolution of
carbonate, by infiltrating fluids, has been invoked as a more efficient means of
liberating CO2 from such rocks (Caciagli and Manning, 2003; Dolejš and Manning,
2010; Frezzotti et al., 2011). Study of fluid inclusions in garnets at Lago di Cignana,
by Frezzotti et al. (2011) revealed the presence of carbonate crystals and dissolved
28
bicarbonate in diamond-bearing fluid inclusions, perhaps indicating dissolution
behaviors at the highest grades.
6. Closed-System Isotopic Exchange Between Mineral Phases
6.1 Carbon
Isotopic equilibration of C in coexisting carbonaceous matter and carbonate has
received some attention as a potential geothermometer (overview in Valley, 2001),
particularly at higher grades at which exchange of carbonate with the recalcitrant
graphite can be more readily achieved (Kitchen and Valley, 1995). The large
difference in initial δ13C of the carbonate (δ13C = -1.5 to +1‰) and marine organic
matter (δ13C = -25 to -20‰) allows assessment of the degree of isotopic equilibration
of the two C reservoirs as a function of increasing temperature. Considerations of
this type must incorporate the large effects on any isotopic shifts of the relative
abundances of the oxidized and reduced C reservoirs (i.e., the fraction of C residing
in carbonate and carbonaceous matter; see Bergfeld et al., 1996). Graphite-rich rocks
containing only very small amounts of the more chemically reactive calcite are more
likely to show apparent isotopic equilibrium between the two phases, as graphite is
known to be sluggish to reequilibrate isotopically.
Previous isotope studies of the lowest-grade units of the Catalina Schists
showed evidence of this calcite-graphite C isotope exchange, with 103ln of +15‰
(Bebout, 1995) achieved in low-calcite metasedimentary rocks rich in carbonaceous
matter, whereas data for the two reservoirs in metasedimentary rocks in the
29
Franciscan Complex indicate a lack of isotopic equilibrium, likely due to the lower
peak metamorphic temperatures experienced by these rocks relative to the low-grade
Catalina Schist equivalents (Sadofsky and Bebout, 2003). Figure 8 demonstrates the
relationships among metamorphic grade and the δ13C and relative proportions of the
two C reservoirs in the W. Alps traverse. For the lowest-grade rocks (Fraiteve), there
is no significant change in the ∆13Ccarb-red (20 to 25‰) as a function of relative
proportions across a very wide range of carbonate:reduced C, though the isotope
values are consistent with expected equilibrium values for these metasediments. For
higher grades, ∆13Ccarb-red shows significant decrease and, for Finestre, Zermatt-Saas,
and Lago di Cignana, the values are as low as 5 to 10‰, this latter range similar to
that suggested in previous studies as representing isotopic equilibrium. Interestingly,
as shown in Figs. 5A and 6A, the C isotope shifts that occur in the two reservoirs are
in general shifted from their initial marine organic or marine limestone values toward
upper mantle δ13C values near -6‰ (see discussion below).
Degrees to which one or the other reservoir show these shifts depend on the
relative proportions of the two phases and the degrees to which they isotopically
reequilibrated, in turn related to exchange kinetics. In some cases, the two phases
were intimately intergrown in the samples in which they occur, better allowing
equilibration. However, in other cases, isotopic equilibration at the scale sampled in
the whole-rock powders could have been less likely, requiring exchange of reduced
and oxidized C in adjacent µm- to mm-scale interlayers/domains. Previous work by
Beyssac et al. (2002) found that carbonaceous matter underwent continuous
30
structural changes and graphitization throughout the Cottian Alps traverse,
potentially facilitating exchange. They also note the formation of harder ring
structures around carbonaceous matter, which could shield the inner material and
explain why only some samples appear to have shifted significantly towards
equilibrium. Similarity of calculated bulk-rock δ13C, using the present δ13C values
and proportions of the two reservoirs, with calculated pre-subduction bulk-sediment
δ13C (using 0‰ for carbonate and -22.5‰ for organic matter and the present
proportions of carbonate and reduced C) is consistent with a closed-system model of
equilibration (Appendix 6).
Figure 8. Carbon isotope reequilibration as a function of metamorphic grade in the W. Alps
metasedimentary suite. Measured Δ13Ccarb-red values are plotted in order of increasing
grade (see Figs. 2B and 5) and colored/sized according to carbonate:reduced weight
percents. Expected equilibrium values from previous work are shown as dotted lines,
sourced from Beaudoin and Therrien (2009). Note that few samples exhibit complete
equilibration. Higher carbonate:reduced ratios generally correspond to smaller Δ13C
values. (Some data could not be plotted due to very low weight percents, resulting in
difficulty determining isotope ratios for one phase.)
31
6.2 Oxygen
Oxygen in these rocks is largely contained in the carbonate, quartz, and
phyllosilicate fractions. Although this study analyzed O isotope ratios only for the
carbonate fraction, several other studies (e.g., Cartwright and Barnicoat, 2003;
Cartwright and Buick, 2000; Henry et al., 1996) analyzed the silicate and
phyllosilicate fractions at exposures of the Schistes Lustres in the Cottian Alps
traverse, the Täschalp region, and Corsica. Figure 9 takes the approach of Henry et
al. (1996) in demonstrating the expected O isotope evolution of the three fractions as
a function of their proportions and the expected isotope fractionation among the
related mineral phases (quartz, phyllosilicate minerals, and calcite). The assumptions
were that quartz and phyllosilicates were present in a 1:2 ratio and that premetamorphism values of δ18O values for quartz, phyllosilicates, and calcite were
+13‰, +18‰, and +29‰, respectively (consistent with discussion in Henry et al.
(1996). Depending on the relative proportions of the silicate and carbonate fractions,
each can exert control over the 18O of the other. At the low carbonate %, the
carbonate 18O values appear consistent with closed-system control by the dominant
silicate fraction, whereas at higher carbonate wt. %, the uniformly low 18O cannot
be explained by closed-system exchange. The most carbonate-rich samples appear to
have been shifted in 18O by more than 10‰ relative to values expected purely by
closed-system exchange.
6.3 Isotopic Exchange Among Adjacent Layers
32
The majority of the microdrilled traverses across metacarbonate layers showed
little variation in either δ13Ccarbonate or δ18Ocarbonate (less than 1‰), as a function of
distance from adjacent metapelite layers, and only one of the microsampled
carbonate layers showed evidence for retention of the much higher marine carbonate
18O values (Appendix 5). This observation is consistent with the lack of obvious
mineralogical zonations at lithological contacts representing metasomatic exchange.
The pervasive nature of the 18O shift is suggestive that fluid infiltration and fluidrock equilibration was achieved at scales greater than those of the individual
interlayers, with the amounts of fluid producing the shifts representing aggregate
sampling of far larger volumes of Schistes Lustres or some external source.
Figure 9. Carbonate O isotope data ordered by carbonate weight %, as compared with
modeled whole-rock equilibration after Henry et al. (1996). The low 18O values for
most of the carbonate-rich samples appear to require control by an alternative reservoir,
likely via infiltration by fluids, whereas at low grades, 18O of the carbonate is consistent
with that expected for closed-system reequilibration with silicate phases. Fractionation
factors from Chacko et al. (1991) and Zheng (1999).
33
7. Open-System Behavior to Explain the O Isotope Systematics
The lowered 18O values for carbonate-rich rocks, relative to those expected
purely for closed-system reequilibration (see Fig. 9), appear consistent with
exchange with fluid previously equilibrated with the silicate fraction, perhaps in
dehydrating metapelitic rocks within the Schistes Lustres section. Bebout et al.
(2013) suggested that the metapelitic rocks in this traverse lost up to 20 % of their
initial H2O during prograde devolatilization reactions, with the largest loss at P-T
corresponding to the Finestre and Cignana units (i.e., over an approximate
temperature range of 450-575 ˚C). Thus, we suggest that it is possible that the O
isotope evolution of the more carbonate-rich rocks can be explained by exchange of
these rocks with H2O-rich fluids generated “internally” within the Schistes Lustres
thrust slices. This hydrous fluid could have been homogenized in an extensive fluid
network or simply have had similar O isotope composition because of derivation
from similar rocks at similar temperatures. Several other studies of the Western Alps
sedimentary suites (including studies at Fraiteve, the lowest grade sampled here)
have postulated an internal fluid source based on the consistency of isotope ratios
between the veins and the wall rock, though some of these models call for O isotope
equilibration over the entire Schistes Lustres metasedimentary unit, requiring
significant fluid infiltration (Cartwright and Barnicoat, 2003; Cartwright and Buick,
2000; Henry et al., 1996). Although it is also possible that infiltration of these
subducting sediments by fluids generated by devolatilization in underlying mafic
34
mineralogies (δ18O = +6 to +20‰; Alt et al., 1992) could have caused this shift, no
evidence has been found that would require this more complicated mechanism.
7.1 Fluid Characteristics
Fluid inclusion studies of the Cottian Alps traverse by Agard et al. (2000)
identified two generations of veins: a series of early blueschist-facies carpholitebearing veins and a set of retrograde greenschist-facies chlorite-bearing veins. All
fluid inclusions were aqueous and contained negligible quantities of CO2, CH4, and
N2 (Agard et al., 2000). The lack of C within the fluid is consistent with the lack of
large scale homogenization of C isotope compositions. Trace element studies of the
Schistes Lustres Cottian Alps traverse and Lago di Cignana report little to no loss of
fluid mobile elements such as Li, Cs, Ba, and N (Bebout et al., 2013; Busigny et al.,
2003; Busigny, 2004), despite the estimate of H2O loss of at least 20% during
prograde devolatilization (Bebout et al., 2013).
7.2 Timing of Fluid Infiltration
Thermodynamic modeling of metasedimentary rocks in the Cottian Alps traverse
by Bebout et al. (2013) allows insight into the timing of potential fluid infiltration.
Between 450 and 550°C, pelitic sediments are expected to lose about 20% water
during the breakdown of chlorite and carpholite. This process (likely responsible for
forming the blueschist-grade veining identified by Agard et al., 2000) could have
produced enough H2O-rich fluid to cause unit-scale equilibration of carbonate δ18O
values. Calculations of expected isotopic shifts were performed as described in
Rumble (1982) for a system undergoing infiltration by one batch of hydrous fluid in
35
isotopic equilibrium with present day quartz values (δ18O = +22‰ for calc-schists
and δ18O = +15‰ for metapelites; from Henry et al., 1996, and Cartwright and
Barnicoat, 1999). According to these results, a fluid in equilibrium with quartz in the
calc-schists would require a fluid:rock volumetric ratio of greater than four. In
contrast, a fluid in equilibrium with metapelites could accomplish this batch shift
with only a 3:4 fluid:rock volumetric ratio. However, the significant variation in
δ18OCarb of these sediments suggests that either the fluid did not reach complete
equilibrium or multiple fluid batches were involved. This is in agreement with work
at Lago di Cignana by Reinecke (1991), which suggested only minor fluid
infiltration could have been present in order to preserve the HP metamorphic
assemblages. Shifts to slightly lower δ18O at the Lago di Cignana locality could be
explained by either localized fluid flow or the noticeably smaller amount of
carbonate in this slice.
Mineralogical evidence from several studies may further help to constrain the
timing of fluid infiltration. Studies by Cartwright and Buick (2000) and Cartwright
and Barnicoat (1999) have found that metasedimentary rocks associated with the
Zermatt-Saas ophiolite have δ18O values of phyllosilicates and quartz that are in
equilibrium at independently determined peak metamorphic temperatures. In
contrast, quartz-calcite values predicted a range of temperatures, with Täschalp
samples predicting temperatures from 400-700°C and Corsican samples exhibiting
complete isotopic reversal as well as temperatures ranging from 200 to 800°C.
However, quartz and calcite tend to exhibit poor equilibrium behavior due to the
36
diffusion properties of quartz, different cooling times, further exchange during
cooling, and retrogression alteration (Sharp and Kirschner, 1994). While it is more
likely that quartz, calcite, and phyllosilicates were at equilibrium at peak
temperatures and that calcite experienced retrograde alteration the possibility of
disequilibrium during crystallization cannot be ruled out. Without further study of
these metasedimentary rocks, it is not possible to conclusively determine when the
oxygen isotope compositions were reset.
7.3 Consistency with Thermodynamic Calculations of Devolatilization in Metapelites
and Metacarbonates
Comparison of preserved mineralogy with published pseudosections allows for
analysis of expected devolatilization behaviors. Pelite metamorphism of the Schistes
Lustres has been well modeled by Bebout et al. (2013). Preserved mineral
assemblages correspond well with predicted phases, and are dominated by chlorite,
carpholite, chloritoid and white mica (in addition to saturated quartz). Carpholite is
predicted to break down around 400°C, coincident with the disappearance of
carpholite in thin sections, and garnet is introduced at temperatures greater than
500°C (though found preserved in lower-grade samples). However, this work does
not predict the modal abundance of calc-silicate phases in high-grade samples at
Finestre and the Zermatt-Saas, as the focus of the calculations was on the Ca-poor
metapelites and the systems modeled did not include CaO.
The mineral assemblages of the carbonate-rich samples investigated in this study
were compared with those predicted in previously published pseudosections (Kerrick
37
and Connolly, 2001a) of marine sediments from the Marianas Trench based on the
similarity of this Marinas Trench composition with the composition of one
carbonate-rich sample from the Albergian Schistes Lustres locality for which wholerock data are available (see the data for sample MIU-955 in Bebout et al., 2013).
According to these calculations, samples from this study would be expected to
contain significant glaucophane and clinopyroxene, not found in any of the samples
studied here, perhaps due to the presence of significant tourmaline (see the
discussion of this issue by Bebout et al., 2013, for the Ca-poor metapelites). The
presence of dolomite is also predicted. It is likely that most samples contain minor
dolomite, perhaps limited by the uptake of Fe and Mg during carpholite formation,
but dolomite is only common at the Lago di Cignana locality. The stability of
lawsonite is predicted by Perple-X modeling and confirmed in two low-grade
localities sampled. Aragonite, while predicted by the thermodynamic calculations,
was probably retrogressed to calcite during decompression. It is interesting to note
that, for typical prograde subduction zone P-T paths, these calculations predict little
or no formation of calc-silicates, which in the Schistes Lustres/Cignana suite are rare
at the lower grades and only minor in abundance at the higher-grade localities (see
Appendix 1).
As noted above (in Section 5.2), the observed relatively minor decarbonation in
this study is wholly inconsistent with the more extensive decarbonation history
predicted by the models of Gorman et al. (2006). For some of these models, which
invoke infiltration of subducting sediment sections by H2O-rich fluid from the
38
underlying oceanic lithosphere (altered oceanic crust and sub-crustal hydrated mantle
rocks), >50% of the carbonate fraction is lost in forearcs and is thus unavailable for
additions to volcanic arcs or the deeper mantle. For the Schistes Lustres/Cignana
suite studied here, a fluid source within the subducting shale-carbonate section
appears to have been adequate to produce the 18O shifts observed in the carbonaterich rocks.
7.4 Comparison with C Isotope Behavior
Under this model, C isotopes behave according to a closed-system devolatilization
model and O isotopes require open-system mechanisms. While the limited evidence
for decarbonation implies only minor fluid flow, explanation of the O isotope
systematics requires fluid to facilitate the significant decrease in O isotope values.
This complication could be caused by a dearth of silicate minerals required for
decarbonation reactions, by fluid infiltration at a temperature too low to facilitate
release of C, or through an alternative explanation for the decrease in O isotope
ratios.
8. Implications for Carbon Cycling Models
First and foremost, this study (together with work on the Franciscan-like suites)
indicates that most of the carbonate and reduced C in subducting shale-carbonate
sediment sections can largely be preserved to depths approaching those beneath
volcanic fronts, depending on the subduction-zone thermal conditions (i.e., the
prograde P-T taken by the rocks). At depths beyond those observed in this study,
39
retention of C in such rocks will be dictated by extents of mobilization during partial
melting, with carbonated fluids traveling through the mantle wedge and contributing
to arc degassing. As investigated in the experimental studies by Tsuno and Dasgupta
(2011, 2012), carbonate in subducting sediments is expected to largely remain stable
through subarc regions and only melt significantly within the deep mantle.
The results of this study can provide an estimate of the recycling efficiency for C
in sediments at this particular subduction zone. Based on the minor proportion of
calc-silicate minerals in the high-grade marbles (usually <10%), we estimate that 8090% of C reached depths of slabs beneath volcanic fronts in the paleo-Tethyan
margin. More broadly considered, and taking into account C inputs in both
sedimentary and oceanic lithosphere, if <20% of initially subducted C is lost in
forearcs, and 15-30% of the C initially subducted is returned to the surface in arcs,
over 50% of the subducting C would be available to enter the deep mantle. Without
an in-depth theoretical investigation of the reaction history, and a more thorough
understanding of subarc processes (e.g., constraining the significance of C release
into fluids or melts), this must remain a general estimate. Also, the uncertainties in
the trench inputs, extents of forearc accretion/underplating/erosion, and arc gas
outputs will complicate estimates of recycling efficiency, with work on individual
margins shows the greatest potential to address these issues.
Most studies of arc gases at individual subduction zones require significant
amounts of CO2 from subducting sediments to balance observed quantities and stable
isotope ratios. However, this work has shown that isotopic exchange between
40
carbonates and reduced C produces large shifts in the δ13C of the subducting C
reservoirs towards mantle values, potentially complicating attempts to distinguish
contributions from the subducting sediment fractions and mantle-derived C (δ13C = 6‰). Current methods of categorizing arc emissions into discrete portions from each
sedimentary lithology (oxidized and reduced C) depend on these isotope shifts (see
examples in Halldórsson et al., 2013; Sano and Marty, 1995). Whereas CO2/3He can
still be used to determine contributions from the mantle vs. sediment cover, isotopic
exchange may make it difficult to assign outgassing CO2 to melting or decarbonation
of a particular sediment source in locations where carbonate and reduced matter are
free to exchange (e.g., Antilles, Marianas).
Additionally, quantities and isotopic compositions of subducting mafic and
ultramafic components remain poorly quantified, and further work is needed to
account for these significant C reservoirs. As the metasedimentary rocks studied here
were not close to complete decarbonation, more substantial decarbonation of
underlying crust and mantle than previously anticipated could be required to explain
present-day measured volcanic C fluxes. In particular, altered oceanic crust has been
suggested to contain significant carbonate, but its role as a contributor to arc gases
could be difficult to evaluate due to its isotopic signature similar to that of mantle C
(Shilobreeva et al., 2011) or marine limestone (for the abundant vein carbonate
produced on the seafloor).
41
9. Conclusions
Intensive study of the metasedimentary rocks exposed in the Schistes Lustres
Cottian Alps traverse and in the Zermatt-Saas ophiolite provides insight into the
metamorphic processes and C retention of subducting sediments. The significant
retention of C and isotopic evolution of these downgoing sediments have important
implications for models of global C cycling and modulation of atmospheric CO2 on
geologically significant timescales. This work has led to the following conclusions.
—This study documents significant C retention to depths of at least 90 km,
approaching those beneath volcanic fronts, arguing against the complete
decarbonation predicted by open system devolatilization models (see Gorman et al.,
2006). However, partial devolatilization (likely not more than 15%) and dissolution
processes cannot be completely ruled out.
—Beyond the depths studied here, a fraction of this C is likely to be transferred
in melts into arcs (following transport through the mantle wedge) while the
remainder continues into mantle storage as a solid.
—Exchange between carbonate and reduced C reservoirs significantly alters the
isotopic signature of metasedimentary rocks during subduction, potentially affecting
mixing models of arc volcanic emissions. If the previously suggested 4:1 ratio of
carbonate to reduced C is subducted, this exchange of C between reservoirs should
not in-bulk change the 13C of the deeply subducted C unless one or the other
reservoir is more effectively removed from subducting sediments by partial melting
beneath volcanic arcs.
42
—These metasedimentary units experienced significant infiltration by a
homogeneous hydrous fluid during subduction, leading to equilibration of O isotopes
at the km-scale. This fluid likely was derived from dehydration of hydrous minerals
in the subducting package (e.g., from dehydrating metapelitic rocks in the same
suite; see Bebout et al., 2013), but could also have been sourced in the underlying
oceanic crust and underlying hydrated mantle.
—The majority of C in these subducting metasediments is retained through the
forearc region, implying significant subduction of global C and limiting the release
of C into the atmosphere, constraining models of long-term C fluxes.
Future work is necessary to further understand the behavior of C in subduction
zones under the wide variety of P-T conditions expected in modern and ancient
subduction zones. Analysis of exhumed subduction-related metamorphic rocks can
complement the larger body of research on subduction zone inputs (sediment section
outboard of trenches) and outputs (in arc volcanic gases), allowing for a better
understanding of C recycling, including the efficiency of return of subducted C into
the atmosphere, and the size and isotopic composition of the C reservoir in the
mantle. Detailed study of the fate of C in subducting altered oceanic crust and
ultramafic lithologies, based on study of exhumed HP/UHP metamorphic suites, will
complement this work and help to form a more complete picture of Earth’s modern
(and ancient) C cycling.
43
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APPENDICES
Appendix 1. Petrography tables, where x: <10 modal percent, xx: 10-50 modal percent, and
xxx: >50 modal percent.
52
53
54
55
Appendix 2. Stable isotope data.
56
57
58
59
Appendix 3. Sampling localities and sample numbers for Cottian Alps Samples. Samples
from Val d’Aosta region were sampled from single outcrop or streams. See Appendix 4 for
coordinates. Place names ©2013 Esri, DeLorme, NAVTEQ, TomTom.
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Appendix 4. Sample coordinates.
Sample
SL98-1
SL98-2
SL98-3
SL98-4
SL99-10
SL99-11
SL99-12
SL99-13
SL99-14
SL99-15
SL99-16
SL99-17
SL99-18
SL99-19
SL99-20
SL99-21
SL99-22
SL99-23
SL99-24
SL99-25
SL99-26
SL99-27
SL99-28
SL99-29
SL99-30
SL99-31
SL99-32
SL99-33
SL12-34
SL12-35
Long.
6.87741
6.87339
6.95732
7.06503
7.05846
6.97338
6.97693
7.02694
7.04264
6.99211
6.99256
6.98366
6.96346
6.86734
6.86393
6.88165
6.86891
6.83160
6.79848
6.80531
6.88138
6.87230
6.86842
6.87612
6.88433
6.88857
6.89144
6.90458
7.07744
7.04949
Lat.
44.97631
44.99232
45.07073
45.05133
44.60743
44.57383
44.57606
44.57613
44.61903
44.66234
44.67503
44.68084
44.69296
44.97494
44.97266
44.97595
44.94972
44.93826
44.96094
44.97578
44.97711
44.98357
44.98843
44.99264
44.99821
45.00170
45.00508
45.02924
45.05297
45.05054
Sample
SL99-36
SL99-37
SL99-38
SL99-39
SL99-40
SL99-41
SL12-42
SL12-43
SL12-44
SL12-45
SL12-46
SL12-47
SL12-48
SL12-49
SL12-50
SL12-51
SL12-52
SL12-54
SL12-56
SL12-58
SL12-59
SL12-60
SL12-61
SL12-62
SL12-63
SL12-64
Clavalité
Servette
Cervinia
Lago di Cignana
61
Long.
6.94005
6.95654
6.97220
6.97943
6.99120
7.02573
7.07421
7.08278
7.07933
7.07547
7.07161
7.05513
7.05050
7.04950
7.04994
7.05465
7.05197
7.05000
7.02543
6.88212
7.11923
7.11595
7.11356
7.11047
7.10615
7.09804
7.49778
7.45335
7.68517
7.59447
Lat.
45.06135
45.06885
45.06012
45.05918
45.06030
45.06365
45.02753
45.03238
45.03272
45.03326
45.05244
45.05003
45.04911
45.06698
45.06862
45.07312
45.07639
45.06181
45.05724
44.97260
45.04668
45.04531
45.04465
45.04592
45.04378
45.04546
45.70408
45.70849
45.92372
45.87691
Appendix 5. Data from microdrilled traverses, plotted as isotope ratio/carbonate weight
percent compared with distance across layer. Sample photos are also shown for context.
62
63
64
65
66
67
68
69
70
71
72
Appendix 6. Comparison of observed bulk δ13C values with values predicted for a closedsystem model, where isotope ratios are controlled by exchange within the system rather than
by either loss or gain of C during metamorphism (the latter loss or gain presumably
involving metamorphic fluids). Calculated with initial δ13C carbonate = 0‰ and initial δ13C
reduced C = -22.5‰. Line shown is for predicted = observed bulk-rock values. Many
samples could not be shown due to difficulty determing isotope ratios for samples with very
low carbonate weight percents.
73
VITA
Jennie Cook-Kollars graduated with honors from Amherst College in 2011
with a B.A. in Chemistry and French. She was a member of the Schupf Scholars
program at Amherst College and a recipient of a Presidential Scholar Fellowship at
Lehigh University.
74