Journal of Petrology Advance Access originally published online on April 22, 2005
Journal of Petrology 2005 46(8):1725-1746; doi:10.1093/petrology/egi034
© The Author 2005. Published by Oxford University Press. All rights reserved. For Permissions, please e-mail: journals.permissions@oupjournals.org
A New Interpretation of Centimetre-scale Variations in the Progress of Infiltration-driven Metamorphic Reactions: Case Study of Carbonated Metaperidotite, Val d'Efra, Central Alps, Switzerland
JOHN M. FERRY1,*,
DOUGLAS RUMBLE, III2,
BOSWELL A. WING3 and
SARAH C. PENNISTON-DORLAND1
1 DEPARTMENT OF EARTH AND PLANETARY SCIENCES, JOHNS HOPKINS UNIVERSITY, BALTIMORE, MD 21218, USA
2 GEOPHYSICAL LABORATORY, CARNEGIE INSTITUTION OF WASHINGTON, 5251 BROAD BRANCH ROAD, NW, WASHINGTON, DC 20015, USA
3 EARTH SYSTEM SCIENCE INTERDISCIPLINARY CENTER, UNIVERSITY OF MARYLAND, COLLEGE PARK, MD 20742, USA
RECEIVED
JULY 28, 2004;
ACCEPTED
MARCH 7, 2005
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ABSTRACT
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Progress (
) of the infiltration-driven reaction, 4olivine +
5CO
2 + H
2O = talc + 5magnesite, that occurred during Barrovian
regional metamorphism, varies at the cm-scale by a factor of
3·5 within an
3 m
3 volume of rock. Mineral and stable
isotope compositions record that
XCO2,
18O
fluid, and
13C
fluid were uniform within error of measurement in the same rock volume.
The conventional interpretation of small-scale variations in
in terms of channelized fluid flow cannot explain the uniformity
in fluid composition. Small-scale variations in
resulted instead
because (a) reactant olivine was a solid solution, (b) initially
there were small-scale variations in the amount and composition
of olivine, and (c) fluid composition was completely homogenized
over the same scale by diffusion–dispersion during infiltration
and subsequent reaction. Assuming isochemical reaction, spatial
variations in
image variations in the (Mg + Fe)/Si of the parent
rock rather than the geometry of metamorphic fluid flow. If
infiltration-driven reactions involve minerals fixed in composition,
on the other hand, spatial variations in
do directly image
fluid flow paths. The geometry of fluid flow can never be determined
from geochemical tracers over a distance smaller than the one
over which fluid composition is completely homogenized by diffusion–dispersion.
KEY WORDS: Alpine Barrovian metamorphism; diffusion; metamorphic fluid composition; metamorphic fluid flow; reaction progress
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INTRODUCTION
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Carbonation and decarbonation reactions during metamorphism
in the crust typically are driven by infiltration of rocks by
chemically reactive fluids (e.g. Ferry & Gerdes, 1998
; Ferry
et al., 2002
). Significant differences in the progress (
) of
the infiltration-driven reactions commonly occur between contrasting
lithologic layers within individual outcrops (Ferry, 1994
; Ferry
& Rumble, 1997
; Ferry
et al., 1998
, 2001
) and, in some cases,
large differences occur between adjacent layers only
1 cm thick
(Ferry, 1987
). The variations in
conventionally are interpreted
in terms of channelized, layer-parallel fluid flow, with elevated
flow in the high-
layers and reduced flow in low-
layers (Ferry,
1987
, 1994
). The interpretation is correct only if there is
no significant chemical communication during metamorphism between
adjacent high-
and low-
layers, either by advection, diffusion
or mechanical dispersion (the combination of the latter two
is referred to as ‘diffusion–dispersion’ in
the rest of the paper). The conventional interpretation once
appeared reasonable because limited cross-layer chemical communication
during metamorphism seemed to be documented by significant layer-by-layer
differences in fluid composition (e.g. Rumble, 1978
; Ferry,
1979
; Kohn & Valley, 1994
), some at the cm-scale (e.g. Rumble
& Spear, 1983
). More recent field (e.g. Bickle
et al., 1997
;
Evans
et al., 2002
; Ague, 2003
), theoretical (e.g. Ague, 2000
,
2002
) and experimental studies (e.g. Wark & Watson 2004
),
however, indicate efficient homogenization of fluid composition
over several metres across lithologic layers during regional
metamorphism caused by exchange of CO
2, H
2O and other fluid
species by diffusion–dispersion. If correct, at least
some layer-by-layer variations in the progress of infiltration-driven
reactions demand another explanation. We propose the alternative
explanation that cm- to m-scale variations in
may result when
adjacent layers initially contain different amounts and/or compositions
of reactant mineral solid solutions, and fluid composition is
homogenized across layering by diffusion–dispersion at
all times during subsequent infiltration and reaction. The importance
of cross-layer diffusion–dispersion in driving metamorphic
devolatilization reactions has been recognized by others as
well (e.g. Hewitt, 1973
; Ague & Rye, 1999
).
Modal, mineral chemical and stable isotope data for the carbonated metaperidotite body in Val d'Efra, Central Alps, Switzerland (Evans & Trommsdorff, 1974), were used to test whether the conventional or new interpretation better explains cm-scale variations in the progress of an infiltration-driven reaction during one instance of Barrovian regional metamorphism. The metaperidotite is nearly ideal for the investigation because of several reasons. First, most samples experienced a single, simple mineral–fluid reaction at or near the peak of Barrovian metamorphism,
 | (1) |
driven
by infiltration of metaperidotite by chemically reactive, relatively
CO
2-rich, CO
2–H
2O fluid. Measured variations in the progress
of
reaction (1),
1, are up to a factor of 2·6 over a
distance of <1 m. Second, rocks contain numerous proxies
for metamorphic fluid composition (mole fraction of the forsterite
component in olivine,
Xfo,Ol, for
XCO2;
18O
Mgs and
18O
Ol for
18O
fluid; and
13C
Mgs for
13C
fluid, where subscripts Ol and Mgs
refer to olivine and magnesite, respectively). The proxies allow
accurate determination of the scale of homogenization of fluid
composition relative to the scale of variations in
1 without
explicit consideration of
T. Third, the minerals are close to
binary Fe–Mg solid solutions and
reaction (1) involves
only one reactant mineral. The relationship between
1 and the
amount and composition of mineral reactants therefore can be
completely and quantitatively represented on a single two-dimensional
diagram. Fourth, the study benefits from three decades of excellent
mineralogical and petrologic work, both on the metaperidotite
body (Evans & Trommsdorff, 1974
) and on associated rocks
in the region (summaries by Pfiffner & Trommsdorff, 1998
;
Pfiffner, 1999
; Nimis & Trommsdorff, 2001
).
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GEOLOGIC SETTING
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The metaperidotite body at Guglia, Val d'Efra (
Fig. 1), is one
of numerous boudins composed of metamorphosed ultramafic and
mafic rocks and rodingite, ranging from several metres to several
hundred metres in size, in the Cima Lunga unit of the Penninic
nappe system (Pfiffner & Trommsdorff, 1998
; Nimis &
Trommsdorff, 2001
). The best known are at Alpe Arami and Cima
di Gagnone. In Val d'Efra, the boudins are set in a matrix of
felsic gneisses, pelitic schists and metacarbonate rocks (Evans
& Trommsdorff, 1974
). Boudins and their host rocks are considered
to have been part of an ocean basin near a continental margin
and exhumed oceanic mantle lithosphere that were subducted,
metamorphosed and uplifted during the Alpine orogeny. The ultramafic
boudins represent oceanic mantle lithosphere. Some of the metamorphosed
mafic and ultramafic rocks retain a mineralogical record of
Eocene (35–43 Ma) ultra-high pressure (UHP) metamorphism.
Mineral equilibria in prograde metamorphosed garnet lherzolite
at Cima di Gagnone,
1 km SW of the metaperidotite body in Val
d'Efra, for example, record
P 30 kbar and
T 740°C (Nimis
& Trommsdorff, 2001
). Where fluids gained access to ultramafic
rocks, as in Val d'Efra, however, all mineralogical evidence
for UHP metamorphism was obliterated by later Alpine Barrovian
regional metamorphism at
P = 6–8 kbar and
T = 600–660°C
(Grond
et al., 1995
).
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Fig. 1. Geologic map of the metaperidotite body at Guglia, Val d'Efra, Central Alps, Switzerland. Body is located on the Osogna 1:25 000 topographic map at 708·9/132·38 (Swiss national grid). Primary metaperidotite lithologies distinguished by texture and mineralogy.
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The metaperidotite body at Guglia, Val d'Efra, is exposed over
a 400–500 m
2 area (
Fig. 1) and vertically over 10–15
m on a vertical exposure that bounds its western margin. The
contact between metaperidotite and surrounding rock is buried
by vegetation and alluvium. Metaperidotite is primarily composed
of two mappable lithologies (
Fig. 1). The schlieren facies is
schist composed of olivine (Ol), talc (Tlc), magnesite (Mgs)
and chlorite (Chl) with and without enstatite (En). [These and
other abbreviations for minerals follow Kretz (1983)
]. Schlieren
are defined by wispy lighter-colored regions, richer in Tlc,
set in a darker matrix richer in Ol [
Fig. 2a of this study and
plate 1A of Evans & Trommsdorff (1974)
]. All samples of
schlieren and matrix collected for this study contain Ol, Tlc,
Mgs and Chl. Some schlieren contain small amounts of En in addition
(0·5–3·8 modal %); the matrix to the schlieren
contains no En. The matrix grades into schlieren over
1 cm;
boundaries between schlieren and matrix cut foliation at a low
angle. Evans & Trommsdorff (1974)
concluded that the schlieren
developed by replacement of the matrix. The schlieren, however,
differ from adjacent matrix in bulk composition [lower (Fe +
Mg)/Si] rather than simply in greater progress of
reaction (1) (e.g. compare modes of matrix sample 16B with schlieren samples
16H and 16 M,
Table 1). We interpret the schlieren as features
that (a) developed prior to Barrovian metamorphism, either by
a primary magmatic process during formation of the igneous parent
rock, by deformation in the mantle, or by Si-metasomatism of
the parent rock during serpentinization, and (b) were then ductilely
deformed during regional metamorphism.
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Fig. 2. Field exposures of metaperidotite. Knife handle in both panels is 9 cm long. (a) Schlieren facies on exposure oblique to foliation. Wispy light-coloured schlieren contain more Tlc and less Ol than surrounding dark matrix. (b) Prismatic enstatite facies with randomly oriented cm-sized En prisms set in a matrix of coarse Tlc, Mgs and Ol.
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Rocks of the prismatic enstatite facies are composed of randomly
oriented, prismatic En crystals, up to several centimetres long,
set in a finer-grained foliated matrix of Ol, Tlc, Mgs and Chl
[
Fig. 2b of this study and Plate 3 of Evans & Trommsdorff
(1974)
]. They differ from rocks of the schlieren facies, both
in their larger grain size and mineralogy (significantly more
En in samples collected for this study, 11–42 modal %).
In three dimensions (3D), the prismatic enstatite facies appears
to form a thin shell,
1–2 m thick, around the margin of
the metaperidotite body. Because of its much greater volume,
this study focused on the schlieren facies.
The metaperidotite is cut by three sets of veins. The commonest, the ‘composite veins’ of Evans & Trommsdorff (1974), are vertical, with NE strike, 1 mm wide, composed of Tlc, Mgs and anthophyllite (Ath), and bounded by a selvage of Ol-free, Tlc–Mgs–Chl rock [Plates 1B and 5 of Evans & Trommsdorff (1974)]. Composite veins are well exposed in the schlieren facies, with typical spacings of 20–40 cm (minimum 0; maximum 130 cm); selvages have fairly uniform thickness, with a half-width of 1–2 cm. As recognized by Evans & Trommsdorff (1974), the composite veins also record infiltration of metaperidotite by reactive CO2-rich, CO2–H2O fluid. Composite veins and their selvages cut across both foliation and schlieren. Selvages of composite veins and adjacent host rocks of the schlieren facies have significantly different Tlc and Mgs contents (e.g. compare samples 7H and 7S, Table 1) and significantly different O- and C-isotope compositions, separated by steep gradients in both modes and isotopic composition. Fluids that produced the composite veins therefore were different from those that drove reaction (1) in the schlieren facies, and they infiltrated the metaperidotite body along fractures after the mineralogy of the schlieren facies developed. The composite veins are not directly relevant to the study's focus on mineral reactions in the schlieren facies. Nevertheless, the composite veins cannot be ignored because their later formation disturbed the stable isotope composition of adjacent samples of the host schlieren facies. In addition to the composite veins, there is a 7-m-long, folded actinolite (Act)–Chl vein and a set of millimetre-wide Ath veins without selvages. Because they are minor constituents, the Act–Chl and Ath veins are not considered further.
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METHODS OF INVESTIGATION
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Internal contacts between lithologies within the metaperidotite
body were mapped in 3D to decimetre accuracy, using a laser
rangefinder and a digital fluxgate compass (
Fig. 1). The locations
of samples within areas of <10 m
2 were recorded with compass
and metal measuring tape.
Twenty-four samples of the schlieren facies, three of the prismatic enstatite facies and three of the composite veins and their selvages were collected for modal, mineral and stable isotope analysis (Fig. 3). All samples of the prismatic enstatite facies and one of the schlieren facies were obtained in place. Because of the difficulty in sampling glacially polished surfaces of the schlieren facies and to avoid defacing the beautiful exposures of the metaperidotite body, the remaining samples of the schlieren facies and the composite veins were obtained from rectangular blocks, 1–5 m in long dimension, that have fallen from the vertical exposure that bounds the western margin of the boudin. The rectangular shapes of the blocks result from their breaking along parallel composite veins that often define two faces of the blocks. Thirteen samples of schlieren facies were obtained from a single block over a 110-cm-long traverse oriented perpendicular to foliation (location 16, Fig. 3). These were supplemented by four other samples from the same block, offset from the line of traverse parallel to foliation. Together, the 17 samples are referred to in the text, figures and tables as from the ‘m-scale traverse’ and are designated samples 16A–16M (a numerical suffix indicates the four samples collected offset from the line of traverse). One of the 17 samples contains a composite vein (16M); no other composite veins occur within or between the other samples. An additional six samples of the schlieren facies (designated 2 and 11–15) were collected from other blocks, and they are representative of the range in colour (and hence in proportions of Ol, Tlc and Mgs) of rocks from the schlieren facies exposed in situ. The three samples of composite veins (designated 7, 17 and 18) include the complete selvage as well as host schlieren facies rock outside the selvage on both sides of the vein. A sample of pelitic schist was collected from each of two outcrops located 10–35 m from the metaperidotite body for mineral thermometry and barometry (locations 6 and 19, Fig. 3).
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Fig. 3. Location map for all samples described in this study. Samples 1, 2, 4 and 5 of metaperidotite and samples 6 and 19 of pelitic schist were collected from outcrop. Other samples were collected from large blocks fallen from the metaperidotite body.
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Mineral assemblages were determined in thin section with optical
petrography and backscattered electron (BSE) imaging using the
JEOL JXA-8600 electron microprobe at Johns Hopkins University.
Compositions of minerals in all samples of the schlieren and
prismatic enstatite facies, two samples of selvages to the composite
veins and the two samples of pelitic schist were determined
by electron microprobe using wavelength-dispersive spectrometry
with natural and synthetic mineral standards and a ZAF correction
scheme (Armstrong, 1988
). X-ray maps were made of garnets in
the pelitic schists, and regions near the rim with the lowest
Mn contents were then analysed for mineral thermometry and barometry.
Modes of all metaperidotite and two vein-selvage samples were
measured by counting
2000 points in thin section using BSE imaging.
Any uncertainty in the identification of a particular point
was resolved by obtaining an energy-dispersive X-ray spectrum.
Magnesite in all samples of the schlieren facies and the three samples of vein selvages was analysed for O- and C-isotope composition, following procedures described by Rumble et al. (1991). Approximately 6–30 mg finely powdered rock was obtained with a 2 mm diamond-tipped drill from a polished rock slab. Magnesite was dissolved overnight in phosphoric acid (McCrea, 1950) in evacuated reaction vessels at 100°C. Evolved CO2 was analysed with the Finnigan MAT 252 mass spectrometer at the Geophysical Laboratory. The acid fractionation factor was taken from Sharma et al. (2002). Results were normalized to the composition of calcite standard NBS-19 (18O = 28·65
, VSMOW; 13C = 1·95, VPDB, Coplen, 1988, 1996). Analyses of NBS-19 and a working calcite standard indicate that analytical precision for both oxygen and carbon isotopes is approximately ±0·1 (1
). All 18O analyses of samples weighing >25 mg are suspect because a significant decrease in T occurred during the relatively long time it took to remove the reaction vessel from the Al-metal heating block and mix the larger powdered samples with phosphoric acid. Values of 18O for samples >25 mg therefore are not reported. Because variations in T during reaction do not affect measurements of C-isotope composition, all measured values of 13C are reported.
The O-isotope composition of Ol in 11 samples of the schlieren facies was measured following procedures of Yui et al. (1995). Olivine separates were obtained by gently crushing samples and hand picking grains under a binocular microscope, followed by ultrasonic cleaning in distilled H2O. Oxygen was extracted from 2 mg of mineral separate in an atmosphere of BrF5 using a CO2 laser fluorination system similar to that of Sharp (1990). The O2 gas was collected, purified and directly analysed with the Finnigan MAT 252 mass spectrometer at the Geophysical Laboratory. Duplicates of all but one sample were measured. Results were normalized to garnet standard UWG-2 (18O = 5·8; Valley et al., 1995), whose composition was measured at the beginning and end of each analytical session. Based on multiple analyses of UWG-2 and of Ol pairs, the precision for 18OOl is considered ±0·1 (1).
Modal abundances of minerals were converted to molar abundances using mineral compositions and molar volumes of mineral components from Holland & Powell (1998). All calculations of mineral equilibria used Holland & Powell's (1998) thermodynamic database and THERMOCALC (version 3.1, 2001). Except for the anorthite component of plagioclase, activities of components in mineral solid solutions were computed from measured mineral compositions and Holland & Powell's AX program. The activity coefficient of the anorthite component in plagioclase was calculated from the experimental data of Goldsmith (1982) at 650°C and 9 kbar using thermodynamic data from Holland & Powell (1998) and THERMOCALC, v. 3·1, following methods described by Carpenter & Ferry (1984).
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MINERALOGY AND MINERAL CHEMISTRY
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Modes and mineral compositions in selected samples of metaperidotite
from the schlieren and prismatic enstatite facies and from the
selvages of the composite veins are listed in
Tables 1 and
2.
Compositions of minerals in the two samples of pelitic schist
used for mineral thermometry and barometry are given in
Table 3.
All samples from the schlieren facies contain Ol, Tlc, Mgs and
Chl, with and without En, along with accessory chromite (Chr),
pyrrhotite (Po) and pentlandite (Pn). Retrograde serpentine
(Srp) and magnetite (Mag) are ubiquitous (
Table 1). Enstatite
occurs in small amounts (0·5–3·8%) in
20%
of the samples. The modal amount of Mgs varies by a factor of
6 (4·7–28·0%). Olivine, Tlc, Mgs and Chl
are close to binary Fe–Mg solid solutions (
Table 2). The
principal divalent cations other than Fe and Mg are Ca, Mn and
Ni, and they occur in relatively small concentrations. In all
analysed minerals, Ca/(Mg + Fe) and Ni/(Mg + Fe) are both <0·004
and Mn/(Mg + Fe) is <0·005. Minerals have remarkably
uniform Mg/(Fe + Mg) and Ol, in particular, displays no growth
zoning. Chlorite contains significant but fairly constant amounts
of Cr, 0·20–0·26 atoms per formula unit.
Analysed samples from the prismatic enstatite facies have the same mineral assemblage as those from the schlieren facies, except that En is always present in substantial amounts (11–42%). There is a complete overlap in measured mineral compositions between the prismatic enstatite and schlieren facies (Table 2). Reconstructed from measured modes and mineral compositions, the range in bulk Fe/(Fe + Mg) of silicates and carbonate in analysed samples from the prismatic enstatite facies (0·081–0·093) overlaps with that of En-free samples from the schlieren facies (0·069–0·093). Likewise, the range in bulk (Mg + Fe)/Si of silicates and carbonate in analysed samples of the prismatic enstatite facies (1·48–1·72) overlaps with that of En-free samples from the schlieren facies (1·47–1·89). The greater amounts of En in rocks of the prismatic enstatite facies compared with those of the schlieren facies cannot be explained in any simple way by differences either in mineral chemistry or in bulk-rock composition.
Selvages to the composite veins are composed of Mgs, Tlc and Chl with accessory Chr, Po and Pn. Selvages are devoid of Ol, except minute quantities (0·1%) that occur as isolated inclusions in Mgs. The selvages are also devoid of retrograde Srp and Mag (e.g. sample 7S, Table 1). A sharp interface, 1 mm wide, separates vein selvages with no Ol (except as inclusions in Mgs) from adjacent rock of the schlieren facies with normal Ol contents (cf. samples 7H and 7S, Table 1). The veins themselves contain the same assemblage as in the vein selvages with the addition of 0·2–1·9% Ath. Compositions of Mgs, Tlc and Chl in the vein selvages are similar to those in the schlieren and prismatic enstatite facies but have systematically slightly higher Fe/(Fe + Mg), the result of reaction (1) having gone to completion in the selvages.
Analysed pelitic schists contain garnet, muscovite, biotite, kyanite, staurolite, plagioclase and quartz, with accessory ilmenite, rutile, monazite and Po, all with unexceptional compositions (Table 3).
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STABLE-ISOTOPE GEOCHEMISTRY
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Measured O- and C-isotope compositions of Mgs and Ol from the
schlieren facies and of Mgs from the selvages to the composite
veins are listed in
Table 4.
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Table 4: Measured reaction progress, olivine composition, and stable isotope compositions for samples of metaperidotite
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The O-isotope composition of Mgs depends on its occurrence.
Magnesite in the schlieren facies >10 cm from a composite
vein has fairly uniform
18O
Mgs = 9·2–9·7
(VSMOW);
18O
Mgs in composite veins and their selvages is significantly
higher, at 10·9–11·9
(
Table 4). Magnesite
in the schlieren facies <10 cm from the vein selvages has
intermediate
18O
Mgs = 9·8–11·2
. Values of
18O
Mgs measured along a traverse from the composite vein in
sample 7 through the vein selvage into adjacent schlieren facies
(
Fig. 4) indicate that a narrow
18O-enrichment halo,
10 cm wide,
exists in the schlieren facies adjacent to the composite vein
selvages.
Magnesite in the schlieren facies likewise has fairly uniform
13C
Mgs = –6·4 to –7·4
(VPDB). There
is also
13C-enrichment within and adjacent to the composite
veins. Measured
13C
Mgs for veins and their selvages is –5·3
to –6·4
. The halo of
13C-enrichment around the
composite veins, however, appears to be restricted to the vein
selvage and does not extend into adjacent host rocks of the
schlieren facies (
Fig. 5).
Analysed Ol in the schlieren facies >10 cm from composite
veins has remarkably uniform
18O
Ol = 4·4–4·7
(
Table 4). The single sample of analysed Ol collected <10
cm from a composite vein (16L1) has slightly higher
18O
Ol =
4·8
—a value, however, the same as the others within
error of measurement.
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PRESSURE, TEMPERATURE AND FLUID COMPOSITION
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Published estimates of
P recorded by mineral assemblages developed
during Barrovian regional metamorphism in the area are: 6–7
kbar (Heinrich, 1982
), 6–8 kbar (Grond
et al., 1995
) and
6·1 kbar (Todd & Engi, 1997
,
fig. 7). Additional
P estimates were calculated from the ‘average
PT’
routine of THERMOCALC, using mineral compositions in the two
analysed samples of pelitic schist (
Table 3) and the mineral
components an, ab, mu, pa, phl, ann, east, alm, py, gr, ilm,
ru, fst, ky and q [abbreviations from Holland & Powell (1998)
],
with
XH2O = 0·8 (as explained below). Results are
P =
7·5 ± 1·4 (±2
) kbar (sample 6) and
7·4 ± 1·6 kbar (sample 19). The preferred
value of
P, based on all four sets of estimates, was taken as
7 ± 1 kbar.
Published estimates of T recorded by mineral assemblages developed during Barrovian regional metamorphism in the area are: 600–650°C (Heinrich, 1982), 600–660°C (Grond et al., 1995) and 645°C (Todd & Engi, 1997). Published estimates are consistent with those computed from the ‘average PT’ routine—628 ± 24°C (±2) for sample 6 and 629 ± 28°C for sample 19. An additional T of metamorphism is recorded independently by the equilibrium between coexisting Ol, Mgs, Tlc and En in the metaperidotite and CO2–H2O fluid. Using representative reduced activities for the Mg-components in the minerals and THERMOCALC, calculated T = 643°C at 7 kbar (Fig. 6). The range in measured mineral compositions and the uncertainty in P of ±1 kbar introduce uncertainties of ±3 and ±7°C, respectively. The T of equilibrium among Ol, Mgs, Tlc, En and CO2–H2O fluid, computed from the data of Berman (1988, updated 1991), using ideal ionic mixing models to calculate reduced activities of Mg-components in minerals, is nearly the same—645°C at 7 kbar. Mineral assemblages in the metaperidotite evidently equilibrated at the same conditions, as did mineral assemblages in other lithologies in the region during Barrovian regional metamorphism. The preferred T of equilibration, based on all five sets of estimates, was taken as 645 ± 10°C.
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Fig. 6. Isobaric T–XCO2 diagram, illustrating selected equilibria among Ol, En, Tlc, Mgs, Atg and CO2–H2O fluid relevant to the metaperidotite. Curves computed for reduced activities of the Mg-components as indicated, estimated from the average compositions of minerals in the schlieren and prismatic enstatite facies (this study) and from Fe–Mg partitioning between coexisting Ol–Atg pairs elsewhere (Ferry, 1995). Coexisting Ol, En, Tlc, Mgs and CO2–H2O fluid record T 645°C and XCO2 0·20 at 7 kbar. Reaction (1) among Ol, Tlc, Mgs and fluid occurs between the two isobaric invariants points: T 575–645°C and XCO2 0·03–0·20.
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The composition of CO
2–H
2O fluid in equilibrium with metaperidotite
of the schlieren and prismatic enstatite facies at the preferred
P–
T conditions of mineral equilibration during Barrovian
metamorphism was
XCO2 = 0·20 ± 0·01 (
Fig. 6),
explaining the value of
XH2O used in the ‘average
PT’ calculations. The uncertainty in
XCO2 is based on
the range in measured mineral compositions.
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CARBONATION OF METAPERIDOTITE
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Carbonation reaction
In principle, carbonation of the metaperidotite body could have
occurred either during amphibolite facies Barrovian regional
metamorphism by
reaction (1), sometime earlier at lower grades
of Barrovian metamorphism, or even prior to Barrovian metamorphism.
If the precursor mineral assemblage subject to carbonation was
not Ol + Tlc + Chl, progressive metamorphism of metaperidotite
in the Alps indicates other plausible possibilities (Trommsdorff
& Evans, 1974
). At progressively lower grades of metamorphism,
rocks with compositions equivalent to those in the metaperidotite
body in Val d'Efra are composed of antigorite (Atg) + Ol + Chl;
brucite (Brc) + Atg + Chl or Atg + Tlc + Chl, depending on whole-rock
(Mg + Fe)/Si; and chrysotile/lizardite (Ctl/Lz) + Brc + Chl
or Ctl/Lz + Tlc + Chl, depending on (Mg + Fe)/Si. The Ctl/Lz
+ Brc + Chl and Ctl/Lz + Tlc + Chl assemblages also correspond
to the mineralogy of any serpentinite precursor that could have
been carbonated prior to Barrovian metamorphism.
There are several arguments that carbonation of the metaperidotite body in Val d'Efra occurred by reaction (1) during Barrovian metamorphism, and that the observed mineral assemblages do not simply represent metamorphism of ultramafic rock carbonated at an earlier time. First, Mgs is the dominant carbonate mineral in Alpine metaperidotites from the amphibolite facies but not in metaperidotites from lower grades (Trommsdorff & Evans, 1974). Metaperidotite elsewhere in the Central Alps at a grade equivalent to that in Val d'Efra is typically composed of Ol + Tlc + Chl. The regional distributions of minerals imply that Mgs formed at conditions of the amphibolite facies by reaction (1).
Second, comparison of the mineral assemblage in the selvages of the composite veins with that in adjacent host rock of the schlieren facies unequivocally demonstrates that Mgs + Tlc in the selvages formed from Ol by reaction (1). The veins and their selvages are undeformed and therefore could not have formed prior to amphibolite facies Barrovian regional metamorphism (Pfiffner, 1999). The selvages of the composite veins are proof that at least some parts of the metaperidotite body were carbonated by reaction (1) during Barrovian metamorphism.
Third, Mgs in the metaperidotite body commonly contains inclusions of Ol but not of Tlc or other minerals. The Ol inclusions are often in optical continuity with each other (Evans & Trommsdorff, 1974) and sometimes with Ol in the matrix. Carbonation of Ol-free equivalents composed of Atg + Brc + Chl, Atg + Tl + Chl, Ctl/Lz + Brc + Chl or Ctl/Lz + Tlc + Chl, either at conditions of lower-grade Barrovian metamorphism or prior to metamorphism, appears to be ruled out. In principle, Mgs with Ol inclusions could develop from either an Ol–Atg–Chl or an Ol–Tlc–Chl precursor. The Ol–Atg–Chl precursor is less likely for two reasons. When modes of analysed samples of metaperidotite are recast as an isochemical equivalent combination of Ol, Atg and Chl, the equivalent Ol–Atg–Chl rocks have modal Atg/(Ol + Atg) = 0·19–1·00. More than half the equivalent Ol–Atg–Chl rocks contain too little Ol to form amounts of Mgs now observed in the metaperidotite body in Val d'Efra by the reaction
 | (2) |
Regardless
of the Atg/(Ol + Atg) of possible precursors, except for implausibly
fortuitous combinations of whole-rock (Mg + Fe)/Si and Fe/(Fe
+ Mg), the uniform measured compositions of Ol in metaperidotite
(
Table 4) cannot be explained by carbonation
reaction (2) followed
by reaction of Atg to Ol + Tlc. On the other hand, as presented
later, carbonation of metaperidotite by
reaction (1) can lead
in a simple and straightforward way to both the observed amounts
of Mgs and the uniform Ol compositions, no matter what the (Mg
+ Fe)/Si and Fe/(Fe + Mg) of the Ol–Tlc–Chl precursors.
T–XCO2 conditions of reaction
Replacement of Ol with Tlc and Mgs by reaction (1) occurs at T between the Atg–Ol–Tlc–Mgs and En–Ol–Tlc–Mgs isobaric invariant points, 575–645°C at 7 kbar (Fig. 6). The corresponding range in XCO2 is 0·03–0·20 at 7 kbar. Carbonation could have occurred at a single T or over any range of T between 575 and 645°C.
Source of carbon
Values of 13CMgs = –5·3 to –7·4 provide the only constraints on the origin of C involved in the carbonation reaction. The C probably was derived from a combination of marine carbonate and reduced organic material, although a source in the mantle cannot be ruled out (Kyser, 1986).
|
SPATIAL DISTRIBUTION OF REACTION PROGRESS
|
|---|
Progress of
reaction (1) was computed for samples from the schlieren
facies and from the selvages to the composite veins as (moles
Mgs)/5, referenced to 1 l of Ol–Tlc–Chl schist prior
to reaction. Measured modes therefore were corrected for the
increase of rock volume caused by
reaction (1) and by the retrograde
reaction that produced Srp. Because Srp replaces both Tlc and
Ol in Ol-bearing rocks, but is absent from the Ol-free selvages
to the composite veins, the Srp-producing reaction probably
was
 | (3) |
The increase in rock volume caused by reactions (1) and (3) is the sum of (
Vs) for each reaction, where Vs is the solid molar volume of reaction. All corrections for the formation of Srp were small—0·2–5·1% of 1. Five samples from the schlieren facies contain small amounts of En that required an additional correction. The pair Tlc + Mgs is stable at or at a lower T than, and En is stable at or at a higher T than the Ol–Tlc–Mg–En isobaric invariant point in Fig. 6. Following Evans & Trommsdorff (1974), En is considered to have formed after reaction (1) by an increase in T and reaction at the P–T–XCO2 conditions of the isobaric invariant point. Under these conditions, the reaction is
 | (4) |
The measured amounts of En in the five samples were corrected for by running reaction (4) backwards to 4 = 0 and adjusting the measured amounts of Tlc, Ol and Mgs accordingly. All corrections for the formation of En were very small, 0·1–0·9% of 1. Calculated values of 1 are listed in Table 4 and illustrated in Fig. 7.
Along the line of the m-scale traverse at location 16,
1 = 0·72–1·88
mol/l—variation of a factor of 2·6 over
1 m (
Fig. 7).
Reaction (1) has occurred but not gone to completion in
every sample along the traverse. Significant differences in
1 occur over distances of several centimetres. Considering samples
collected from positions offset from the line of traverse as
well,
1 varies by a factor of 3·5 within a volume of
rock
3 m
3. Values of
1 are not necessarily higher in Tlc-rich
schlieren than in the Ol-rich matrix (cf. samples 16B, 16H and
16 M,
Tables 1 and
4). The schlieren therefore did not develop
simply from greater
1 than in surrounding rock, but are regions
where Tlc/(Ol + Tlc) was elevated in the Ol–Tlc–Chl
precursor prior to carbonation because of lower whole-rock (Mg
+ Fe)/Si. The range in measured
1 for the m-scale traverse is
similar to the range for samples of the schlieren facies collected
from other parts of the metaperidotite body (open and filled
diamonds,
Fig. 7). Metamorphic processes that controlled
1 along
the m-scale traverse therefore are representative of those that
affected the body as a whole. Values of
1 adjacent to the selvages
of the composite veins (filled diamonds,
Fig. 7) are not greater
than those measured for samples far from the veins (open diamonds).
In terms of reaction progress, the effects of vein formation
do not extend beyond the vein selvages.
Reaction (1), however,
has gone to completion in the vein selvages themselves; measured
values of
1 for the selvages (open squares,
Fig. 7) indicate
that the maximum value of
1 4 mol/l. Measured values of
1 in
samples from the schlieren facies therefore correspond to
9–63%
reaction.
|
SPATIAL DISTRIBUTION OF FLUID COMPOSITION
|
|---|
Rocks of the schlieren facies contain four proxies for metamorphic
fluid composition:
Xfo,Ol,
18O
Mgs,
18O
Ol and
13C
Mgs. The proxies
are considered, rather than the corresponding fluid compositional
variables themselves, because they are directly measured quantities
not subject to uncertainties introduced by estimates of
P and
T and by activity–composition relations. If
P and
T were
uniform across the metaperidotite body at all times during Barrovian
regional metamorphism, spatial variations in fluid composition
can simply be tracked by variations in the proxies.
Activities of components in Ol, Tlc, Mgs and fluid are related through the equilibrium constant for reaction (1),
 | (5) |
where subscripts of the mineral activity terms
refer to the Mg-components. For Fe–Mg Ol, Tlc, and Mgs
solid solutions, the compositions of Tlc and Mgs are related
to the composition of Ol through Fe–Mg exchange constants,
 | (6) |
 | (7) |
In CO
2–H
2O
fluids,
XH2O is 1 –
XCO2. Given
a–
X relations for
the mineral and fluid solutions, a value of
Xfo,Ol therefore
uniquely defines and is a proxy for the
XCO2 of coexisting fluid.
Values of
Xfo,Ol are uniform among samples along the line of
and offset from the m-scale traverse (
Fig. 8). A calculated
mean square weighted deviation (MSWD) of 1·42 (Mahon,
1996
) demonstrates that all values of
Xfo,Ol along the traverse
are consistent within error of measurement with a single value
whose best estimate is the weighted mean, 0·888 ±
0·001 (±95% confidence interval for the standard
error). Correspondingly,
Xfo,Ol records a single value of
XCO2 within error of measurement. Measured values of
Xfo,Ol along
the m-scale traverse are similar to those measured in samples
of the schlieren facies from other parts of the metaperidotite
body (right-hand panel of
Fig. 8). In addition, there are no
large differences in
Xfo,Ol and, hence,
XCO2 between samples
of the schlieren facies <10 cm from composite veins and those
farther away. Some of the small differences in
Xfo,Ol between
samples along the m-scale traverse and samples from other parts
of the body, however, are statistically significant. Whereas
measurable differences in
XCO2 did not occur during metamorphism
over distances of
1 m or less, small differences in
XCO2 <
0·01 did develop over the scale of the entire metaperidotite
body (
30 m or less).
The
18O of Mgs and Ol are proxies for
18O of fluid. With two
exceptions, measured values of
18O
Mgs along the m-scale traverse
are consistent with a single value (MSWD = 1·90) whose
best estimate is 9·37 ± 0·06
(
Fig. 9).
The exceptions, samples 16L1 and 16 M, occur <10 cm from
a composite vein and, like other samples of the schlieren facies
collected near composite veins, experienced
18O-enrichment associated
with formation of the vein (cf.
Fig. 4 and right-hand panel
of
Fig. 9). Measured values of
18O
Ol along the m-scale traverse
are consistent with a single value (MSWD = 1·50) whose
best estimate is 4·62 ± 0·09
(
Fig. 10).
The O-isotope compositions of Mgs and Ol record uniform
18O
fluid along the m-scale traverse during Barrovian metamorphism within
error of measurement. Although
18O
Mgs along the m-scale traverse
is similar to that measured for samples of the schlieren facies
from other parts of the metaperidotite body (
Fig. 9), some of
the small differences are statistically significant. Like
XCO2,
18O
fluid therefore was uniform over distances comparable to
the m-scale traverse but not over distances on the scale of
the entire body.
The
13C of Mgs is a proxy for
13C of fluid. With the exception
of three samples at the ends (A, L1, M), measured values of
13C
Mgs from the m-scale traverse are consistent with a single
value (MSWD = 1·66), whose best estimate is –6·81
± 0·06
TOP
ABSTRACT
INTRODUCTION
