JOURNAL OF PETROLOGY
VOLUME 46
NUMBER 8
PAGES 1725–1746
2005
doi:10.1093/petrology/egi034
A New Interpretation of Centimetre-scale
Variations in the Progress of Infiltrationdriven 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
ADVANCE ACCESS PUBLICATION APRIL 22, 2005
KEY WORDS: Alpine Barrovian metamorphism; diffusion; metamorphic
fluid composition; metamorphic fluid flow; reaction progress
Progress ( j) of the infiltration-driven reaction, 4olivine þ 5CO2 þ
H2O ¼ talc þ 5magnesite, that occurred during Barrovian
regional metamorphism, varies at the cm-scale by a factor of
35 within an 3 m3 volume of rock. Mineral and stable
isotope compositions record that XCO2, d18Ofluid, and d13Cfluid
were uniform within error of measurement in the same rock
volume. The conventional interpretation of small-scale variations
in j in terms of channelized fluid flow cannot explain the
uniformity in fluid composition. Small-scale variations in j
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 j image variations in the (Mg þ 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 j 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.
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 (x) 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 x conventionally are
interpreted in terms of channelized, layer-parallel fluid
flow, with elevated flow in the high-x layers and reduced
flow in low-x layers (Ferry, 1987, 1994). The interpretation is correct only if there is no significant chemical
communication during metamorphism between adjacent
high-x and low-x layers, either by advection, diffusion or
mechanical dispersion (the combination of the latter two
*Corresponding author. Telephone: 410-516-8121. Fax: 410-5167933. E-mail: [email protected]
# The Author 2005. Published by Oxford University Press. All
rights reserved. For Permissions, please e-mail: journals.permissions@
oupjournals.org
INTRODUCTION
JOURNAL OF PETROLOGY
VOLUME 46
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 CO2, H2O and other fluid species by
diffusion–dispersion. If correct, at least some layerby-layer variations in the progress of infiltration-driven
reactions demand another explanation. We propose the
alternative explanation that cm- to m-scale variations in
x 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,
4ðMg,FeÞ2 SiO4 þ 5CO2 þ H2 O
olivine
fluid
¼ ðMg,FeÞ3 Si4 O10 ðOHÞ2 þ 5ðMg,FeÞCO3
talc
magnesite
ð1Þ
driven by infiltration of metaperidotite by chemically
reactive, relatively CO2-rich, CO2–H2O fluid. Measured
variations in the progress of reaction (1), x1, are up to
a factor of 26 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; d 18OMgs and d18OOl for
d18Ofluid; and d13CMgs for d13Cfluid, where subscripts Ol
and Mgs refer to olivine and magnesite, respectively).
The proxies allow accurate determination of the scale of
NUMBER 8
AUGUST 2005
Germany
0
N
m
Austria
10
Switzerland
sample locations
study area
Italy
1
5
4
3
surrounding rock
and cover
schlieren facies
(Ol-Tlc-Mgs-Chl±En)
prismatic enstatite facies
(En-Ol-Tlc-Mgs-Chl)
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 7089/13238 (Swiss national grid). Primary
metaperidotite lithologies distinguished by texture and mineralogy.
homogenization of fluid composition relative to the scale
of variations in x1 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 x1 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).
GEOLOGIC SETTING
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.
1726
FERRY et al.
CENTIMETRE-SCALE VARIATIONS IN METAMORPHIC REACTIONS
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 ¼ 6–8 kbar and T ¼ 600–660 C (Grond
et al., 1995).
The metaperidotite body at Guglia, Val d’Efra, is
exposed over a 400–500 m2 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 (05–38 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 þ 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.
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
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.
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
1727
JOURNAL OF PETROLOGY
VOLUME 46
NUMBER 8
AUGUST 2005
Table 1: Mineral assemblages and modes for selected samples of metaperidotite
Sample:
1
3
7H
7S
12
14
16B
16D
16H
16I2
16 J
16 M
Lithology:*
PEF
SF
SF/VH
VS
SF
SF
SF/MT
SF/MT
SF/MT
SF/MT
SF/MT
SF/MT
Olivine
26.34
15.45
22.58
52.32
24.88
50.59
29.26
32.95
35.06
68.08
17.17
45.23
24.41
32.32
37.63
34.90
40.89
32.55
30.97
38.64
31.59
27.01
35.41
11.85
11.69
20.47
22.04
14.04
28.02
17.99
0
9.03
0
6.26
23.01
0.67
4.67
22.05
12.04
0
4.97
0
4.82
0
5.88
0
6.84
0
5.22
0
5.57
8.88
1.42
6.86
0
0.05
0
0.25
0
0.49
0
0.44
0
0.79
0
0.15
0
0
0.05
0
0.25
0
0.15
0.10
0.24
0.20
0.30
0.20
0.19
0.43
0
0.30
0.20
0
0.24
tr
0.49
0.20
tr
0.20
0.05
tr
0.25
0.05
tr
1.16
0
1.18
tr
1.52
0
0
1.94
0
0.10
0.39
0.05
0.20
tr
0.10
3.59
0.10
2.26
0.10
4.81
Talc
Magnesite
Enstatite
Chlorite
Anthophyllite
Chromite
Pentlandite
Pyrrhotite
Magnetite
Serpentine
0.10
56.59
38.43
0
3.83
0.19
0
2.71
4.47
0.39
0.39
tr
2.95
0.10
5.67
0.10
19.87
Values in vol %; tr, <0.05%.
*PEF, prismatic enstatite facies; SF, schlieren facies, sample from part of metaperidotite body other than m-scale traverse;
SF/VH, schlieren facies, host rock adjacent (<10 cm) to composite vein and its selvage; VS, selvage of composite vein;
SF/MT, schlieren facies, sample from m-scale traverse.
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.
METHODS OF INVESTIGATION
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 m2 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
1728
FERRY et al.
CENTIMETRE-SCALE VARIATIONS IN METAMORPHIC REACTIONS
schlieren facies
prismatic enstatite facies
pelitic schist
sample
locations
11
14
13
17 A 16
M
15 18
1
4
2
7
5
N
3
6
12
outcrop of carbonated
metaperidotite
0
m
20
19
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.
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).
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 diamondtipped 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 (d 18O ¼ 2865%, VSMOW; d13C ¼ 195%,
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
01% (1s). All d18O 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 d18O 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 d 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 (d18O ¼ 58%; Valley et al., 1995), whose
1729
JOURNAL OF PETROLOGY
VOLUME 46
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 d 18OOl is
considered 01% (1s).
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 31, 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. 31, following methods
described by Carpenter & Ferry (1984).
MINERALOGY AND MINERAL
CHEMISTRY
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
(05–38%) in 20% of the samples. The modal amount
of Mgs varies by a factor of 6 (47–280%). 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 þ
Fe) and Ni/(Mg þ Fe) are both <0004 and Mn/(Mg þ
Fe) is <0005. Minerals have remarkably uniform Mg/
(Fe þ Mg) and Ol, in particular, displays no growth zoning. Chlorite contains significant but fairly constant
amounts of Cr, 020–026 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 þ Mg) of silicates and carbonate
in analysed samples from the prismatic enstatite facies
NUMBER 8
AUGUST 2005
(0081–0093) overlaps with that of En-free samples from
the schlieren facies (0069–0093). Likewise, the range in
bulk (Mg þ Fe)/Si of silicates and carbonate in analysed
samples of the prismatic enstatite facies (148–172)
overlaps with that of En-free samples from the schlieren
facies (147–189). 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
(01%) 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 02–19% 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 þ
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).
STABLE-ISOTOPE GEOCHEMISTRY
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.
The O-isotope composition of Mgs depends on its
occurrence. Magnesite in the schlieren facies >10 cm
from a composite vein has fairly uniform d 18OMgs ¼
92–97% (VSMOW); d 18OMgs in composite veins and
their selvages is significantly higher, at 109–119%
(Table 4). Magnesite in the schlieren facies <10 cm from
the vein selvages has intermediate d 18OMgs ¼ 98–112%.
Values of d 18OMgs 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
18
O-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 d13CMgs ¼ –64 to 74% (VPDB). There is
also 13C-enrichment within and adjacent to the composite veins. Measured d 13CMgs for veins and their selvages
is 53 to 64%. The halo of 13C-enrichment around
the composite veins, however, appears to be restricted to
1730
FERRY et al.
CENTIMETRE-SCALE VARIATIONS IN METAMORPHIC REACTIONS
Table 2: Compositions of minerals (cations per formula unit) in selected samples of metaperidotite
Olivine
Sample:
1
3
7H
12
14
16B
16D
16H
16I2
16J
16 M
Lithology:
PEF
SF
SF/VH
SF
SF
SF/MT
SF/MT
SF/MT
SF/MT
SF/MT
SF/MT
1.785
0.212
1.785
0.214
1.787
0.210
1.773
0.221
1.793
0.198
1.779
0.217
1.787
0.209
1.774
0.220
1.791
0.207
1.780
0.219
1.770
0.234
0.001
0.005
0.003
0.005
0.002
0.005
0.002
0.005
0.004
0.007
0.003
0.006
0.003
0.006
0.002
0.005
0.002
0.006
0.002
0.005
0.002
0.005
0.000
0.998
0.000
0.997
0.000
0.997
0.000
1.000
0.000
0.999
0.000
0.998
0.000
0.997
0.000
0.999
0.000
0.997
0.000
0.997
0.000
0.995
100.28
0.891
99.73
0.890
100.24
0.891
100.08
0.886
99.98
0.896
100.31
0.887
100.25
0.891
99.91
0.886
100.09
0.893
99.89
0.887
100.40
0.880
Mg
Fe
Mn
Ni
Ca
Si
Oxide sum
Mg/M2þ
Talc
Sample:
1
3
7H
7S
12
14
16B
16D
16H
16I2
16J
16M
Lithology:
PEF
SF
SF/VH
VS
SF
SF
SF/MT
SF/MT
SF/MT
SF/MT
SF/MT
SF/MT
2.903
0.066
2.898
0.068
2.882
0.068
2.872
0.073
2.887
0.067
2.892
0.061
2.899
0.065
2.916
0.065
2.900
0.067
2.918
0.064
2.929
0.069
2.891
0.074
0.000
0.008
0.001
0.008
0.000
0.009
0.000
0.009
0.001
0.007
0.001
0.011
0.001
0.009
0.001
0.009
0.001
0.009
0.001
0.009
0.001
0.009
0.000
0.009
0.007
4.006
0.008
4.006
0.006
4.016
0.004
4.020
0.005
4.013
0.004
4.014
0.004
4.010
0.005
4.001
0.005
4.008
0.005
4.001
0.006
3.991
0.009
4.006
94.96
0.975
95.09
0.974
94.63
0.974
95.20
0.972
94.68
0.975
94.90
0.975
95.31
0.975
95.18
0.975
95.60
0.974
95.39
0.975
94.69
0.974
94.75
0.972
Sample:
1
3
7H
7S
12
14
16B
16D
16H
1612
16J
16M
Lithology:
PEF
SF
SF/VH
VS
SF
SF
SF/MT
SF/MT
SF/MT
SF/MT
SF/MT
SF/MT
Mg
Fe
Mn
Ni
Al
Si
Oxide sum
Mg/M2þ
Magnesite
Mg
Fe
Mn
Ca
0.938
0.058
0.936
0.056
0.929
0.065
0.922
0.073
0.936
0.060
0.938
0.054
0.931
0.060
0.937
0.056
0.934
0.060
0.942
0.051
0.939
0.055
0.932
0.062
0.001
0.003
0.005
0.003
0.003
0.003
0.003
0.002
0.001
0.003
0.004
0.004
0.005
0.004
0.003
0.004
0.003
0.003
0.003
0.004
0.002
0.004
0.002
0.004
48.89
49.08
49.02
49.26
48.93
48.93
48.98
49.23
49.11
48.67
48.81
49.02
Sample:
1
3
7H
7S
12
14
16B
16D
16H
1612
16J
16M
Lithology:
PEF
SF
SF/VH
VS
SF
SF
SF/MT
SF/MT
SF/MT
SF/MT
SF/MT
SF/MT
Oxide sum
Chlorite
Mg
Fe
Mn
Ni
Al
Cr
Si
Oxide sum
Mg/M2þ
4.801
0.246
4.871
0.263
4.848
0.215
4.631
0.403
4.782
0.237
4.723
0.232
4.785
0.228
4.724
0.233
4.802
0.256
4.780
0.219
4.783
0.244
4.770
0.268
0.001
0.009
0.002
0.011
0.000
0.010
0.000
0.010
0.000
0.007
0.001
0.016
0.001
0.010
0.001
0.011
0.002
0.011
0.001
0.010
0.000
0.010
0.002
0.011
1.543
0.162
1.435
0.252
1.551
0.248
1.515
0.232
1.557
0.244
1.511
0.250
1.510
0.227
1.503
0.258
1.496
0.198
1.528
0.235
1.493
0.248
1.479
0.243
3.192
87.24
3.162
87.22
3.114
87.60
3.168
87.34
3.135
86.81
3.193
87.03
3.185
87.13
3.194
87.42
3.194
87.92
3.173
87.62
3.175
87.02
3.184
87.08
0.949
0.946
0.956
0.918
0.951
0.950
0.952
0.951
0.947
0.954
0.949
0.944
1731
JOURNAL OF PETROLOGY
VOLUME 46
NUMBER 8
AUGUST 2005
Table 2: Continued
Enstatite and anthophyllite
Sample:
1
12
16 M
Lithology:
PEF
SF
SF/MT
VS
Mineral:
Enstatite
Enstatite
Enstatite
Anthophyllite
0.000
1.800
0.001
1.780
0.000
1.791
0.006
5.660
0.199
0.003
0.206
0.003
0.202
0.003
1.286
0.015
0.001
0.001
0.001
0.001
0.001
0.001
0.007
0.007
0.002
1.996
0.002
2.001
0.002
1.999
0.015
7.997
99.75
0.899
100.63
0.894
100.58
0.897
97.64
0.812
Ca
Mg
Fe
Mn
Ni
Cr
Al
Si
Oxide sum
Mg/M2þ
7S
Analyses are averages of five ‘spot’ analyses of three to five grains in thin section (except for olivine analyses that are
averages of 1027 analyses). Mineral formulas for olivine are cations per 4 oxygen atoms; for talc, cations per 11 oxygen
atoms (less H2O); for magnesite, cations per oxygen atom (less CO2); for chlorite, cations per 14 oxygen atoms (less H2O);
for enstatite, cations per 6 oxygen atoms; for anthophyllite, cations per 23 oxygen atoms (less H2O). Oxide sum refers to
the sum of oxide wt %, excluding CO2 and H2O, with all Fe as FeO. M2þ ¼ Ca þ Mg þ Fe þ Mn þ Ni. Notation for sample
lithology as in footnote to Table 1.
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 d 18OOl ¼ 44–47%
(Table 4). The single sample of analysed Ol collected
<10 cm from a composite vein (16L1) has slightly higher
d18OOl ¼ 48%—a value, however, the same as the
others within error of measurement.
PRESSURE, TEMPERATURE AND
FLUID COMPOSITION
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 61 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 ¼ 08 (as explained
below). Results are P ¼ 75 14 (2s) kbar (sample 6)
and 74 16 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 (2s)
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 ¼ 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.
The composition of CO2–H2O 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 ¼
020 001 (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.
1732
FERRY et al.
CENTIMETRE-SCALE VARIATIONS IN METAMORPHIC REACTIONS
Table 3: Compositions of minerals (cations per formula unit)
in analysed pelitic schists
CARBONATION OF
METAPERIDOTITE
Carbonation reaction
Micas
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 þ Tlc þ 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) þ Ol þ Chl; brucite
(Brc) þ Atg þ Chl or Atg þ Tlc þ Chl, depending
on whole-rock (Mg þ Fe)/Si; and chrysotile/lizardite
(Ctl/Lz) þ Brc þ Chl or Ctl/Lz þ Tlc þ Chl, depending
on (Mg þ Fe)/Si. The Ctl/Lz þ Brc þ Chl and Ctl/Lz þ
Tlc þ 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 þ Tlc þ 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 þ 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 þ Brc þ Chl, Atg þ Tl þ Chl, Ctl/Lz þ
Brc þ Chl or Ctl/Lz þ Tlc þ Chl, either at conditions
of lower-grade Barrovian metamorphism or prior to
Sample:
6
19
6
19
Mineral:
Muscovite
Muscovite
Biotite
Biotite
0.808
0.142
0.079
0.830
0.120
0.074
0.869
0.888
0.048
1.215
0.038
1.183
0.083
0.000
0.029
0.102
0.000
0.029
1.052
1.071
0.003
0.123
0.004
0.127
1.830
0.881
1.816
0.864
0.435
1.250
0.438
1.264
3.119
95.94
0.488
3.136
2.751
2.736
95.52
0.421
96.05
0.536
95.71
0.525
K
Na
Fe
Mg
Mn
Ti
AlVI
AlIV
Si
Oxide sum
Fe/(Fe þ Mg)
Garnet and staurolite
Sample:
6
19
6
Mineral:
Garnet
Garnet
Staurolite
Fe
Mg
Mn
Ca
Ti
Al
Si
Oxide sum
Fe/(Fe þ Mg)
2.303
0.439
2.161
0.471
0.048
0.258
0.045
0.364
n.m.
2.013
n.m.
1.988
2.964
100.15
2.985
100.57
0.840
0.821
2.959
0.574
0.027
n.m.
0.135
17.694
7.811
97.21
0.837
Plagioclase and ilmenite
Sample:
Mineral:
Ca
Na
K
Al
Si
Oxide sum
Xan
6
Plagioclase
19
Sample:
Plagioclase
Mineral:
0.196
0.791
0.006
0.276
Fe
0.709
0.006
Mg
1.185
2.813
1.274
2.727
100.17
0.193
99.75
0.277
Mn
Ti
Oxide sum
Fe/(Fe þ Mg)
6
Ilmenite
0.990
0.005
0.012
0.997
100.32
0.995
19
Ilmenite
0.975
0.005
0.024
0.998
99.87
0.995
Analyses are averages of 511 ‘spot’ analyses of two to five
grains in thin section. Mineral formulas for micas are cations
per 11 oxygen atoms (less H2O); for garnet, cations per 12
oxygen atoms; for staurolite, cations per 46 oxygen atoms
(less H2O); for plagioclase, cations per 8 oxygen atoms; for
ilmenite, cations per 3 oxygen atoms. n.m., not measured.
Other notation as in footnotes to Tables 1 and 2.
1733
JOURNAL OF PETROLOGY
VOLUME 46
NUMBER 8
AUGUST 2005
Table 4: Measured reaction progress, olivine composition, and stable isotope compositions for samples of metaperidotite
d 18OMgs (%)x
Sample:
Lithology
Location (cm)*
x1 (mol/l)y
Xfo,Olz
2
SF
3
SF
7 host
SF/VH
1.21 (0.13)
0.93 (0.11)
0.91 (0.11)
0.892 (0.003)
0.890 (0.005)
0.891 (0.004)
7 selvage
VS
11
SF
12
SF
3.78 (0.21)
1.21 (0.13)
1.97 (0.16)
0.888 (0.004)
0.886 (0.006)
13
SF
14
SF
15
SF
0.56 (0.09)
0.35 (0.07)
0.64 (0.09)
0.898 (0.004)
0.896 (0.004)
0.899 (0.003)
16A
SF/MT
0
16B
SF/MT
5
16C
SF/MT
15
1.38 (0.14)
1.73 (0.15)
0.99 (0.12)
0.888 (0.005)
0.888 (0.003)
0.885 (0.007)
16C1
SF/MT
15, 20h, 0v
16D
SF/MT
29
16E
SF/MT
37
1.34 (0.13)
1.88 (0.15)
1.55 (0.15)
0.886 (0.003)
0.891 (0.003)
0.890 (0.002)
16F
SF/MT
44
16G
SF/MT
52
16H
SF/MT
60
1.53 (0.14)
1.72 (0.15)
1.13 (0.12)
0.890 (0.004)
0.889 (0.003)
0.886 (0.003)
16I
SF/MT
68
16I1
SF/MT
68, 378h, 0v
16I2
SF/MT
68, 378h, 76v
1.26 (0.13)
1.29 (0.13)
2.52 (0.18)
0.887 (0.003)
0.885 (0.007)
0.893 (0.005)
16 J
SF/MT
79
16K
SF/MT
96
16L
SF/MT
107
1.50 (0.14)
1.31 (0.13)
0.88 (0.11)
0.887 (0.006)
0.884 (0.004)
0.877 (0.023)
16L1
SF/MT
107, 142h, 0v
16 M
SF/MT
113
17 host
SF/VH
1.47 (0.14)
0.72 (0.10)
0.34 (0.07)
0.887 (0.004)
0.881 (0.023)
0.893 (0.002)
17 selvage
VS
3.91 (0.21)
18 selvage
VS
n.m.
d 18OOl (%)x
d 13CMgs (%)x
9.57
9.33
n.m.
4.44
7.22
6.70
9.8210.28
n.m.
10.9211.23
9.22
4.57
7.36 to 7.28
6.36 to 5.66
6.37
9.62
9.28
4.54
n.m.
7.26
6.49
l.s.
n.m.
l.s.
9.38
n.m.
7.00
7.01
n.m.
4.65
6.38
6.86
n.m.
6.64
6.75
9.41
l.s.
9.41
9.40
9.22
9.40
9.23
9.48
9.44
n.m.
4.70
n.m.
4.73
4.45
4.46
7.10
6.87
6.90
6.75
n.m.
6.66
6.70
l.s.
9.67
n.m.
4.64
6.76
6.82
9.22
9.19
4.55
n.m.
6.68
6.94
l.s.
9.79
n.m.
4.78
6.91
7.10
6.44
6.71
5.83 to 5.26
6.05 to 6.02
11.20
n.m.
n.m.
11.05
10.9511.31
11.7411.86
Notation for sample lithology as in footnote to Table 1. n.m., not measured.
*Horizontal distance in cm along m-scale traverse measured perpendicular to foliation relative to position of sample 16A.
Samples 16C1, 16I1, 16I2 and 16L1 are offset from line of traverse parallel to foliation with distance of offset in horizontal (h)
and vertical (v) dimensions noted in cm.
yProgress of reaction (1) relative to 1 litre OlTlcChl precursor. Numbers in parentheses are 2s uncertainties estimated
from point counting statistics (Chayes, 1956).
zMole fraction Mg2SiO4 component in olivine. Average of 1027 measurements (15 measurements typical). Numbers in
parentheses are 2s uncertainties.
xOxygen isotope compositions relative to VSMOW; carbon isotope compositions relative to VPDB. 2s uncertainty is 0.2%
for all analyses. d18OMgs not reported for large samples (l.s.), >25 mg (see text). Range in values for samples 7, 17 and
18 represents two to four analyses separated by 0.54 cm on polished slab (see Figs 4 and 5).
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 þ
Atg) ¼ 019–100. 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
1734
34ðMg,FeÞ2 SiO4 þ 20CO2 þ 31H2 O
olivine
fluid
ð2Þ
¼ ðMg;FeÞ48 Si34 O85 ðOHÞ62 þ 20ðMg;FeÞCO3 :
antigorite
magnesite
9.0
host schlieren facies
adjacent to selvage
-6.0
>10 cm
δ13C Mgs (‰ VPDB)
>10 cm
10.0
<10 cm
vein selvage
11.0
m-scale
traverse
sample 7
all other SF (>10 cm)
sample 7
δ18O Mgs (‰ VSMOW)
-5.0
m-scale
traverse
<10 cm
12.0
other SF (>10 cm)
CENTIMETRE-SCALE VARIATIONS IN METAMORPHIC REACTIONS
host schlieren facies
adjacent to selvage
vein selvage
FERRY et al.
-7.0
others
others
-8.0
0
10
20
30
40
50
0
10
20
30
40
50
distance perpendicular to
vein, from vein center (mm)
distance perpendicular to
vein, from vein center (mm)
Fig. 4. Left-hand panel illustrates O-isotope composition of Mgs
(stippled rectangles) from the selvage around the composite vein in
sample 7 and from adjacent host rock of the schlieren facies. Rectangles
have vertical dimensions that correspond to the 2s uncertainty in
measured d 18OMgs and horizontal dimensions that correspond to the
distance over which rock was sampled for analysis. Right-hand panel
summarizes d 18OMgs for all other samples of the schlieren facies (SF).
Vertical lines represent the range in measured values expanded by
02% (2s). The same elevated d 18OMgs in the schlieren facies
adjacent to the vein selvage in sample 7 also occurs in samples of the
schlieren facies from the m-scale traverse collected <10 cm from a
composite vein. Taken together, all analyses define a halo of
18
O-enrichment within the vein selvages and extending 10 cm into
adjacent schlieren facies.
Fig. 5. Left-hand panel illustrates C-isotope composition of Mgs
(stippled rectangles) from the selvage around the composite vein in
sample 7 and from adjacent host rock of the schlieren facies. Dimensions of rectangles as in Fig. 4. Right-hand panel summarizes d 13CMgs
for all other samples of the schlieren facies (SF). Vertical lines represent
the range in measured values expanded by 02% (2s). As in the
schlieren facies adjacent to the vein selvage in sample 7, there is no
elevated d 13CMgs in samples of the schlieren facies from the m-scale
traverse collected <10 cm from a composite vein. Together, all analyses define a halo of 13C-enrichment in Mgs within the selvage that, in
contrast to the 18O-enrichment, does not extend into adjacent host
schlieren facies.
Regardless of the Atg/(Ol þ Atg) of possible precursors,
except for implausibly fortuitous combinations of wholerock (Mg þ Fe)/Si and Fe/(Fe þ 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 þ 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 þ Fe)/Si
and Fe/(Fe þ 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 003–020 at
7 kbar. Carbonation could have occurred at a single T or
over any range of T between 575 and 645 C.
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
Source of carbon
Values of d 13CMgs ¼ –53 to 74% provide the only
constraints on the origin of C involved in the carbonation
1735
6ðMg,FeÞ2 SiO4 þ ðMg,FeÞ3 Si4 O10 ðOHÞ2
olivine
talc
þ 9H2 O ¼ 5ðMg,FeÞ3 Si2 O5 ðOHÞ4 :
fluid
serpentine
ð3Þ
JOURNAL OF PETROLOGY
VOLUME 46
NUMBER 8
AUGUST 2005
700
En
Fo +
Tlc
En
Tlc + Mg s
Fo
+M
gs
a fo = a en = 0.80
a tlc = 0.93
a mgs = 0.94
a atg = 0.20
620
Tlc
T (˚C)
660
Fo
gs
+M
En
Atg
Fo + Tlc
Atg
580
Tlc + Mgs
0.1
m-scale traverse
on line
offset
3.0
I2
378h
76v
1.0
D
B
2.0
20h
0v
F
G
378h
0v
C1 E
A
C
L1
142h
0v
J
I1
H
I
K
L
M
0.0
40
80
120
0
distance perpendicular to foliation (cm)
Fo
Atg + Mgs
540
0.0
4.0
>10 cm
<10 cm
selvages
P = 7000 bars
ξ1 (mol/L Ol-Tlc-Chl precursor)
others
0.2
0.3
0.4
0.5
X CO 2
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 020 at 7 kbar. Reaction (1) among Ol, Tlc, Mgs and fluid
occurs between the two isobaric invariants points: T 575–645 C
and XCO2 003–020.
The increase in rock volume caused by reactions (1) and
(3) is the sum of x(DVs) for each reaction, where DVs is the
solid molar volume of reaction. All corrections for the
formation of Srp were small—02–51% of x1. Five samples from the schlieren facies contain small amounts of En
that required an additional correction. The pair Tlc þ
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
08ðMg,FeÞ3 SiO4 O10 ðOHÞ2 þ 06ðMg,FeÞ2 SiO4
talc
olivine
þ 02ðMg,FeÞCO3 ¼ 19ðMg,FeÞ2 Si2 O6
magnesite
enstatite
þ 08H2 O þ 02CO2 :
f luid
ð4Þ
The measured amounts of En in the five samples were
corrected for by running reaction (4) backwards to x4 ¼ 0
Fig. 7. All measured values of the progress of reaction (1), x1, in
samples from the schlieren facies and selvages to the composite veins.
Error bars represent 2s based on the statistics of point counting
(error bar not displayed when smaller than size of symbol). Left-hand
panel illustrates data for all samples from the m-scale traverse, sample
location 16 (circles). Filled circles correspond to samples collected along
the line of the traverse. Open circles represent samples displaced from
the line of traverse parallel to foliation over horizontal (h) and vertical
(v) distances given in centimetres. Sample designations have location
‘16’ prefix omitted; a number suffix identifies a sample displaced from
the line of traverse. The total variation in x1 is by a factor of 35;
significant differences in x1 occur between samples a few centimetres
apart. Right-hand panel summarizes all other x1 measurements.
Values for vein selvages (open squares) correspond to rocks in which
reaction (1) has gone to completion. The absence of a significant
difference in x1 between samples from the schlieren facies <10 cm
(filled diamonds) and >10 cm (open diamonds) from composite veins
demonstrates that formation of the veins had no effect on x1 outside
the vein selvage.
and adjusting the measured amounts of Tlc, Ol and Mgs
accordingly. All corrections for the formation of En were
very small, 01–09% of x1. Calculated values of x1 are
listed in Table 4 and illustrated in Fig. 7.
Along the line of the m-scale traverse at location 16,
x1 ¼ 072–188 mol/l—variation of a factor of 26 over
1 m (Fig. 7). Reaction (1) has occurred but not gone to
completion in every sample along the traverse. Significant differences in x1 occur over distances of several centimetres. Considering samples collected from positions
offset from the line of traverse as well, x1 varies by a
factor of 35 within a volume of rock 3 m3. Values of
x1 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 x1 than in surrounding rock, but
are regions where Tlc/(Ol þ Tlc) was elevated in the
Ol–Tlc–Chl precursor prior to carbonation because of
lower whole-rock (Mg þ Fe)/Si. The range in measured
x1 for the m-scale traverse is similar to the range for
samples of the schlieren facies collected from other parts
1736
CENTIMETRE-SCALE VARIATIONS IN METAMORPHIC REACTIONS
0.91
others
m-scale traverse
on line
offset
0.89
M
L
0.87
weighted mean = 0.888±0.001
MSWD = 1.42 (N = 17)
>10 cm
<10 cm
of the metaperidotite body (open and filled diamonds,
Fig. 7). Metamorphic processes that controlled x1 along
the m-scale traverse therefore are representative of those
that affected the body as a whole. Values of x1 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 x1 for the selvages (open squares, Fig. 7)
indicate that the maximum value of x1 4 mol/l.
Measured values of x1 in samples from the schlieren facies
therefore correspond to 9–63% reaction.
measured X fo,Ol
FERRY et al.
0.85
0
SPATIAL DISTRIBUTION OF FLUID
COMPOSITION
80
120
distance perpendicular to foliation (cm)
Rocks of the schlieren facies contain four proxies for
metamorphic fluid composition: Xfo,Ol, d 18OMgs, d18OOl
and d 13CMgs. 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),
K1 ¼ ½ðafo Þ4 ðaCO2 Þ5 ðaH2 O Þ=½ðatlc Þðamgs Þ5 40
ð5Þ
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,
ðFe=MgÞOl =ðFe=MgÞTlc ¼ KOl=Tlc
ð6Þ
ðFe=MgÞOl =ðFe=MgÞMgs ¼ KOl=Mgs :
ð7Þ
In CO2–H2O fluids, XH2O is 1XCO2. 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 142 (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, 0888 0001
(95% confidence interval for the standard error).
Fig. 8. Average mole fraction forsterite component of Ol (Xfo,Ol) in
samples from the schlieren facies based on 10–27 analyses per sample.
Error bars represent 2s uncertainty. Symbols are as in Fig. 7. Larger
uncertainties for samples L and M result from more extensive
serpentinization. Left-hand panel illustrates that all measured data
from the m-scale traverse (sample location 16) are statistically
consistent with a single value whose best estimate is the weighted
mean ¼ 0888 (dashed line); grey band represents the 95%
confidence interval based on the standard error (0001). Right-hand
panel illustrates that small but statistically significant differences in
Xfo,Ol exist between samples from the m-scale traverse and from
other parts of the metaperidotite body. The complete overlap in
Xfo,Ol between samples from the schlieren facies >10 cm and <10 cm
from composite veins demonstrates that formation of the veins had no
effect on Xfo,Ol outside the vein selvage. Taken together, all data
indicate the scale of XCO2 homogenization was between 1 and 30 m.
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 < 001 did develop over the scale of
the entire metaperidotite body ( 30 m or less).
The d 18O of Mgs and Ol are proxies for d 18O of fluid.
With two exceptions, measured values of d18OMgs along
the m-scale traverse are consistent with a single value
(MSWD ¼ 190) whose best estimate is 937 006%
(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,
1737
JOURNAL OF PETROLOGY
VOLUME 46
NUMBER 8
AUGUST 2005
others
5.5
weighted mean = +9.37±0.06‰
MSWD = 1.90 (N = 12)
L1
9.0
omitted from fit
8.0
0
40
120
80
distance perpendicular to foliation (cm)
4.0
weighted mean = +4.62±0.09‰
MSWD = 1.50 (N = 8)
3.5
0
experienced O-enrichment associated with formation
of the vein (cf. Fig. 4 and right-hand panel of Fig. 9).
Measured values of d 18OOl along the m-scale traverse
are consistent with a single value (MSWD ¼ 150) whose
best estimate is 462 009% (Fig. 10). The O-isotope
compositions of Mgs and Ol record uniform d 18Ofluid
along the m-scale traverse during Barrovian metamorphism within error of measurement. Although d18OMgs
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, d 18Ofluid
therefore was uniform over distances comparable to the
m-scale traverse but not over distances on the scale of the
entire body.
The d13C of Mgs is a proxy for d13C of fluid. With
the exception of three samples at the ends (A, L1, M),
measured values of d13CMgs from the m-scale traverse are
consistent with a single value (MSWD ¼ 166), whose
best estimate is 681 006% (Fig. 11). Although
values of d13CMgs along the m-scale traverse overlap
with those of samples of the schlieren facies from other
80
120
Fig. 10. Measured d 18OOl for samples from the schlieren facies. Error
bars and symbols same as in Figs 7 and 8. Left-hand panel illustrates
that all measured data from the m-scale traverse (sample location 16)
are statistically consistent with a single value whose best estimate is
þ462 009% VSMOW. Right-hand panel illustrates d 18OOl for
samples from other parts of the metaperidotite body. Taken together,
all data confirm the scale of d 18Ofluid homogenization was 1 m.
others
weighted mean = -6.81±0.06‰
MSWD = 1.66 (N = 14)
-5.0
δ13C Mgs (‰ VPDB)
Fig. 9. Measured d OMgs for samples from the schlieren facies and
selvages to the composite veins. Error bars and symbols are as in Figs 7
and 8. Left-hand panel illustrates that data from the m-scale traverse
(sample location 16), excluding samples L1 and M, are statistically
consistent with a single value whose best estimate is þ937 006%
VSMOW. Right-hand panel illustrates that small but statistically
significant differences in d 18OMgs exist between samples from the
m-scale traverse and from other parts of the metaperidotite body.
Values of d 18OMgs in the schlieren facies <10 cm from composite
veins are intermediate between those of the vein selvages and those of
the schlieren facies >10 cm from veins (see also Fig. 4). The d 18O of the
schlieren facies evidently is disturbed for 10 cm from the veins by
18
O-enrichment associated with formation of veins. For this reason,
d 18OMgs of samples L1 and M from the m-scale traverse, that occur
<10 cm from a composite vein, were omitted from the estimate of
the weighted mean for the rest of the traverse. Taken together,
all data indicate the scale of d 18Ofluid homogenization was between
1 and 30 m.
40
distance perpendicular to foliation (cm)
18
18
others
4.5
m-scale traverse
on line
offset
-6.0
A
M
-7.0
L1
omitted from fit
selvages
10.0
5.0
>10 cm
M
>10 cm
<10 cm
m-scale traverse
on line
offset
δ18O Ol (‰ VSMOW)
11.0
m-scale traverse
on line
offset
>10 cm
<10 cm
selvages
δ18O Mgs (‰ VSMOW)
12.0
-8.0
0
40
120
80
distance perpendicular to foliation (cm)
Fig. 11. Measured d 13CMgs for samples from the schlieren facies and
selvages to the composite veins. Error bars and symbols are as in Figs 7
and 8. Left-hand panel illustrates that, with the exception of samples at
the ends (A, L1, M), measured data from the m-scale traverse (sample
location 16) are statistically consistent with a single value whose best
estimate is 681 006% VPDB. Right-hand panel illustrates 1%
13
C-enrichment in the vein selvages compared with rocks of the schlieren facies (see also Fig. 5). The overlap in d 13CMgs between samples
<10 cm and >10 cm from veins demonstrates that the 13C-enrichment
does not extend out of vein selvages into the adjacent host schlieren
facies. Small but statistically significant differences in d 13CMgs exist
between samples from the m-scale traverse and in the schlieren facies
from other parts of the metaperidotite body. Taken together, all data
indicate that the scale of d 13Cfluid homogenization was between
1 and 30 m.
1738
CENTIMETRE-SCALE VARIATIONS IN METAMORPHIC REACTIONS
parts of the metaperidotite body, small but statistically
significant differences occur. Like XCO2 and d 18Ofluid,
d13Cfluid was uniform over distances comparable to the
m-scale traverse but not over distances on the scale of
the entire body.
Taken together, data in Figs 7–11 firmly establish that
fluid composition during Barrovian metamorphism was
uniform within error of measurement in the same 3 m3
volume of rock in which progress of infiltration-driven
reaction (1) varies by a factor of 35.
X CO 2
0.197
0.200
0.205
INTERPRETATIONS OF cm-SCALE
VARIATIONS IN REACTION
PROGRESS
Conventional interpretation
0.211
0.220
0.230
X ofo = 0.9215
noOl = 21.2 mol/L
noTlc = 0
0.92
calculated X fo,Ol
FERRY et al.
0.88
X fo = 0.888
o
X fo = 0.9215
noOl = 10.6 mol/L
noTlc = 3.4 mol/L
0.84
0.80
for all:
noChl = 0.3 mol/L
X ofo = 0.9000
noOl = 21.2 mol/L
noTlc = 0
0.76
When a carbonation or decarbonation reaction is driven
by infiltration of rock by chemically reactive fluid, reaction progress is proportional to the time-integrated
fluid flux (Baumgartner & Ferry, 1991; Ferry & Gerdes,
1998). Metre- to cm-scale variations in the progress of
infiltration-driven reactions, as in Fig. 7, therefore are
conventionally interpreted in terms of channelized fluid
flow with elevated flow in the high-x areas and reduced
flow in the low-x areas. The interpretation implies that
rocks were physically and chemically isolated from each
other at the cm–m scale over which the variations in x1
are observed. Assuming isochemical metamorphism (in
the petrologic sense), Xfo,Ol decreases and XCO2 correspondingly increases with increasing x1 because of the
fractionation of Fe and Mg among Ol, Tlc and Mgs
(Fig. 12). If samples from the m-scale traverse were isolated systems, the different values of x1 should systematically correlate with differences in Ol composition;
specifically, high-x samples should contain Ol with relatively low Xfo,Ol and low-x samples should contain Ol with
relatively high Xfo,Ol. The predicted range in Xfo,Ol for
the observed range in x1, 07–25 mol/l, is at least 002
(Fig. 12)—a range that would be unequivocally detected
by microprobe analysis (Fig. 8). The absence of any
statistically significant correlation between x1 and Xfo,Ol
therefore suggests some other process accounts for the
variations in x1 in Fig. 7. The process, in particular,
must be consistent with spatially uniform Xfo,Ol and
fluid composition.
New interpretation
Qualitative description and quantitative analysis
Spatial variations in the progress of an infiltration-driven
reaction inevitably develop in a suite of rocks that experiences the same reaction if (a) at least one of the mineral
reactants is a solid solution, (b) different rocks initially
contain different amounts and/or compositions of the
0
1
2
3
4
5
ξ1 (mol/L Ol-Tlc-Chl pr otolith)
Fig. 12. Representative examples of the decrease in Xfo,Ol caused by
progress of reaction (1) in metaperidotite. Curves calculated from mass
balance of Fe and Mg for rocks with initial amounts and compositions
of minerals as indicated, using model mineral compositions and Fe–Mg
exchange constants in Table 5. Associated values of XCO2 refer only to
o
¼ 09215, and were computed
rock with noOl ¼ 212 mol/l and Xfo;Ol
for P ¼ 7 kbar and T ¼ 645 C from K1, assuming CO2–H2O fluid,
calculated mineral compositions, and a–X relations described in
text. Right-hand ends of curves correspond to reaction (1) gone to
completion.
reactant mineral(s), and (c) fluid composition is the same
at all times and in all samples during reaction (if, for
example, it is homogenized by diffusion–dispersion).
Samples of metaperidotite that record different values of
x1 illustrate the process. Consider two rocks with Ol of
the same composition prior to reaction (e.g. Xfo,Ol ¼
09215), but the first rock contains twice as much Ol
(e.g. 212 mol/l) as the second (e.g. 106 mol/l). The
remainder of the rocks is Chl Tlc but not Mgs. Fluid
composition changes as reaction (1) proceeds, but at any
one time it is the same in both rocks; Ol composition
must be the same as well. At all times, x1 in the first rock
therefore must be approximately twice x1 in the second in
order to maintain the equality in Ol composition (Fig. 12).
In this case, the difference in x1 between the two rocks
develops because the rate of reaction (1), qx1/qt, in the
Ol-rich rock is approximately twice that in the Ol-poor
rock on a volume basis. Alternatively, consider two rocks
with the same amount of Ol prior to reaction (e.g.
212 mol/l) but one contains more Mg-rich Ol (e.g.
Xfo,Ol ¼ 09215) compared with the other (e.g. Xfo,Ol ¼
09000). If, for example, reaction (1) initiates in the
first rock at 7 kbar and 645 C, fluid composition is
buffered by reactants and products to XCO2 ¼ 0197
(Fig. 12). Fluid with XCO2 ¼ 0197, however, is in
equilibrium with Ol with Xfo,Ol ¼ 09, and reaction (1)
1739
VOLUME 46
NUMBER 8
AUGUST 2005
o
o
m
ol
/L
nOl -X fo values
m-scale traverse
on line
offset
o l=
o
l =
O
o
ol
2
=1
mo
n Ol
2.0
o
n Ol
=8
/L
m
16
nO
20
3.0
pr
ec
ur
so
r
4.0
n
does not proceed in the second rock. With continued
infiltration and progress of reaction (1), Ol composition
in the first rock eventually reaches Xfo,Ol ¼ 09 (Fig. 12).
With further infiltration, reaction (1) then proceeds
in both rocks. If infiltration and reaction (1) cease when
Xfo,Ol < 09 in both samples (e.g. Xfo,Ol ¼ 0888), x1 in the
first rock will be larger than x1 in the second (Fig. 12). In
the second case, differences in x develop not so much
from differences in reaction rate, but from differences in
the duration of reaction. Specifically, all else being equal,
reaction (1) proceeds longer in rocks that initially contain
relatively Mg-rich Ol and for a shorter time in rocks that
initially contain relatively Mg-poor Ol.
A systematic quantitative analysis of the process, specifically for progress of reaction (1) driven by infiltration
of metaperidotite by CO2–H2O fluid, appears in Fig. 13.
Rocks are considered composed of Ol, Tlc, Chl
(0289 mol/l) and 1 modal % other inert minerals (Chr,
Po, Pn) prior to reaction. Initial amounts of Ol (noOl ) are
varied between 8 and 20 mol/l; Tlc content prior to
reaction is the difference in volume between 1 l and the
volumes of the other minerals. Mineral solid solutions are
modelled with the formulas in Table 5. Initial composio
) are varied between 088 and 094
tions of Ol (Xfo;Ol
Xfo,Ol; compositions of the other minerals both prior to
and at all times during reaction are specified by the
composition of Ol and the Fe–Mg exchange constants
in Table 5. The ranges in initial amounts and compositions of Ol are those appropriate to metamorphism of the
carbonated metaperidotite body in Val d’Efra. Figure 13
illustrates the value of x1 needed to achieve a final Ol
composition of Xfo,Ol ¼ 0888 (Fig. 8) as a function of the
amount and composition of Ol prior to reaction (1).
Plotting coordinates are chosen to linearize the relationo
and x1. As expected, x1 at
ship between noOl , Xfo;Ol
constant initial Ol composition is greater in rocks
with greater initial amounts of Ol. At constant initial
abundance of Ol (inclined solid contours), x1 is greater
in rocks with Mg-richer Ol prior to reaction. No reaction
occurs in any rock with Xofo,Ol 0888. Significant
differences in x1 can result from differences in Xofo,O1 as
small as 001. Open and filled circles correspond to the
initial amounts and compositions of Ol in all samples
from the m-scale traverse, obtained by running measured
values of x1, x3 and (where relevant) x4 backwards to zero
and then computing the corresponding initial mineral
abundances and (from mass balance of Fe and Mg)
their initial compositions. The measured range in x1
for samples from the m-scale traverse (07–25 mol/l)
can be completely explained by modest cm-scale
variations in the amount (noOl ¼ 124–199 mol/l)
o
¼ 0901–0920) of Ol prior to
and composition (Xfo;Ol
reaction (1) and complete homogenization of fluid
composition across the traverse, most plausibly by
diffusion–dispersion.
ξ1 for final Xfo,Ol = 0.888 (mol/L)
JOURNAL OF PETROLOGY
l/L
l/L
mo
1.0
0.0
0.88
0.90
0.92
0.94
o
Xfo,Ol in Ol-Tlc-Chl precursor prior to reaction (1)
Fig. 13. Quantitative relationship between progress of reaction (1), x1,
needed to produce a final Ol composition of Xfo,Ol ¼ 0888 (Fig. 8) and
o
) of Ol prior to reaction.
the amount (noOl ) and composition (Xfo;Ol
o
¼
Curves specifically are for the range noOl ¼ 8–20 mol/l and Xfo;Ol
088–094, relevant to carbonation of the metaperidotite body, and are
calculated using model mineral formulas and Fe–Mg exchange
constants in Table 5. Open and filled circles correspond to calculated
o
for all samples from the m-scale traverse
values of noOl and Xfo;Ol
(sample location 16). The measured range in x1 ¼ 07–25 mol/l
along the traverse can be explained by a range in noOl ¼ 12–20 mol/
o
¼ 090–092, and complete homogenization of fluid
l, a range in Xfo;Ol
composition along the traverse during reaction (1).
Table 5: Mineral formulas and Fe–Mg exchange
constants used to model mineral reactions
Mineral
Model formula
Olivine
(Mg,Fe)2SiO4
Talc
(Mg,Fe)3Si4O10(OH)2
Magnesite
(Mg,Fe)CO3
Enstatite
(Mg,Fe)2Si2O6
Chlorite
(Mg,Fe)515Al146Cr024Si315O10(OH)8
K Ol=Tlc ¼ ½ðFe=MgÞOl =½ðFe=MgÞTlc ¼ 5276ð0214Þ
K Ol=Mgs ¼ ½ðFe=MgÞOl =½ðFe=MgÞMgs ¼ 1950ð0143Þ
K Ol=En ¼ ½ðFe=MgÞOl =½ðFe=MgÞEn ¼ 1108ð0049Þ
K Ol=Chl ¼ ½ðFe=MgÞOl =½ðFe=MgÞChl ¼ 2392ð0129Þ
Mineral formulas based on microprobe data, ignoring trace
Ca, Mn, Ni and Ti in all phases and Al and Cr in all but Chl.
K values are the average of measured values for all samples;
1s uncertainty in parentheses.
Evidence against the conventional interpretation
There remains a remote possibility that rocks indeed
were chemically isolated at the cm scale during metamorphism, that the differences in measured values of x1
1740
CENTIMETRE-SCALE VARIATIONS IN METAMORPHIC REACTIONS
(Fig. 7) represent different values of time-integrated
o
, and that any resulting
flux independent of noOl and Xfo;Ol
variations in Xfo,Ol after reaction are simply below the
detection limits of electron microprobe analysis. The
possibility was evaluated by a set of representative calculations that predict the composition of Ol that would
result if rocks were chemically isolated along the
m-scale traverse and if time-integrated flux, and hence
o
. The
x1, were completely unrelated to noOl and Xfo;Ol
measured value of x1 in each sample from the m-scale
traverse was first assigned at random to one of the other
16 samples. Starting with the abundances and compositions of minerals in each sample prior to reaction (1),
reaction (1) was then run forward by the new amount
and the final Ol composition that developed was calculated. Results appear in Fig. 14. The model for chemically isolated samples predicts a range of final Xfo,Ol
¼ 0857–0901—far larger than that observed (Fig. 8).
Given a representative analytical precision of Xfo,Ol in the
samples of metaperidotite (00032, 1s; the average of
1s uncertainties for samples from location 16 in Table 4),
the set of calculated Ol compositions in Fig. 14 is inconsistent with a single value at a high degree of statistical
significance (MSWD ¼ 155). Thus, except for an unreasonably contrived set of initial conditions, any model of
metamorphism of the metaperidotite based on chemical
isolation of samples at a scale <1 m fails to account for
both the large variations in measured x1 and the uniformity in measured Xfo,Ol (Figs 7 and 8). The new interpretation, based on spatially uniform XCO2 at the m scale
during metamorphism, is further supported by evidence
for spatial uniformity in d18Ofluid and d13Cfluid at the
same scale as well (Figs 9–11).
Origin of differences in the amount and composition of
Ol in Ol–Tlc–Chl schist
A full understanding of reaction progress in the metaperidotite requires an explanation of the process
that controlled the initial amounts and compositions of
Ol in Ol–Tlc–Chl schist prior to reaction (1). Variations
in the amount and composition of Ol prior to reaction
(1) were caused by variations in the bulk composition
of the metaperidotite prior to Barrovian metamorphism.
Many metaperidotites in the Central Alps, and the one
in Val d’Efra in particular, do not have bulk compositions
of normal igneous rocks (Trommsdorff & Evans, 1974;
Pfiffner, 1999). They have too low Ca contents to be
typical oceanic peridotites (Dick et al., 1984) and too
high Si contents to be dunites. The metaperidotite in
Val d’Efra could have been either a peridotite that lost
nearly all Ca or a dunite that gained Si by metasomatism,
e.g. associated with serpentinization (Pfiffner, 1999). Both
possibilities are modelled to explain the variations in the
0.90
calculated X fo,Ol
FERRY et al.
0.88
m-scale traverse
on line
offset
0.86
MSWD = 15.5 (N = 17)
0.84
0
40
80
120
distance perpendicular to foliation (cm)
Fig. 14. Calculated Xfo,Ol that results when measured progress of
reaction (1), x1, in each sample from the m-scale traverse (sample
location 16) is randomly assigned to one of the other 16 samples, and
reaction (1) then is run forward by that amount starting from the
o
of each sample. Calculations used model
calculated noOl and Xfo;Ol
mineral formulas and Fe–Mg exchange constants in Table 5.
Measured mean Xfo,Ol (dashed line) and 95% confidence interval
(grey band) for samples from the traverse shown for reference (from
Fig. 8). Error bars on calculated Xfo,Ol (circles) correspond to a 2s
uncertainty of 00064, the average 2s of measured values in Table 4.
Calculated values of Xfo,Ol, 0857–0901, are significantly more
variable than measured compositions. The calculated values of Xfo,Ol
are not consistent with a single value to a high level of statistical
significance. Results demonstrate that the measured variations in x1
and uniform values of Xfo,Ol along the m-scale traverse were not
produced by channelized fluid flow through samples that were
chemically isolated from each other while reaction (1) proceeded.
amount and composition of Ol in the Ol–Tlc–Chl schist
prior to carbonation.
In the case of Ca-depleted peridotite, the parent rock is
considered composed of 98 modal % Ol þ En þ diopside
(Di), 1% Al–Cr Spl that reacts with Ol and En to form
0289 mol/l Chl, and 1% minerals that remain inert
during metasomatism and regional metamorphism
(Chr, Po, Pn). Olivine, En and Di in peridotite have
uniform compositions. The proportions of Ol and pyroxene (Px), however, are allowed to vary in the range Px/
(Ol þ Px) ¼ 0–08 by volume. The relative proportion of
En to Di is fixed and corresponds to that in average
oceanic peridotite, Di/En ¼ 0175 by volume (Dick
et al., 1984). The variations in Px/(Ol þ Px) are intended
to represent dm-scale alternations of Px- and Ol-rich
layers observed in oceanic peridotite (Loney et al., 1971;
Dick & Sinton, 1979; Boudier & Coleman, 1981). In the
first stage of the reaction history, peridotite loses all Ca
(at constant Mg, Fe and Si) and is hydrated to Ol–Tlc–
Chl schist using Ca(Mg,Fe)Si2O6 as for the composition
of Di and formulas in Table 5 for the other minerals.
The exact sequence of reactions, that in nature would
1741
VOLUME 46
have involved Ctl/Lz, Atg, Brc, and possibly UHP minerals as intermediate reaction products (Trommsdorff &
Evans, 1974; Pfiffner & Trommsdorff, 1998), is inconsequential. Amounts of Ol and Tlc in the model Ol–Tlc–
Chl schist depend on Px/(Ol þ Px) of the parent rock; Ol
composition was computed from mass balance of Fe and
Mg and the Fe–Mg exchange constants in Table 5. The
partitioning of Fe and Mg between Ol and Di was taken
as that for Ol and En (Table 5), as appropriate for the
elevated T at which the peridotite parent originally
formed (Loucks, 1996). During the second stage, reaction (1) proceeds in Ol–Tlc–Chl schist until Ol composition reaches 0888 Xfo,Ol (Fig. 8). Different values of x1
develop at the end of the second stage of reaction,
o
prior to reaction (1), that,
depending on noOl and Xfo;Ol
in turn, depend on Px/(Ol þ Px) of the parent peridotite.
The range in measured values of x1 along the m-scale
traverse (Fig. 7) are quantitatively reproduced for Xfo,Ol ¼
0923 (a representative mantle value) and Px/(Ol þ Px) in
the range 005–050 by volume in the original peridotite
(Fig. 15). There could have been variations in Xfo,Ol as
well as in Px/(Ol þ Px) in the peridotite, but these are not
required by the data. For Xfo,Ol ¼ 0923 and Px/(Ol þ
Px) ¼ 005–050 in the peridotite parent, predicted
o
¼ 0902–
ranges in noOl ¼ 121–205 mol/l and Xfo;Ol
0920 of Ol–Tlc–Chl schist nearly match those inferred
for samples along the m-scale traverse from measured
modes and mineral compositions (124–199, 0901–
0920, Fig. 13).
In the case of a silicified dunite, the parent rock is
considered composed of 98 modal % Ol, 1% Al–Cr Spl
that reacts with Ol and SiO2 to form 0289 mol/l Chl,
and 1% minerals that remain inert during metasomatism
and regional metamorphism (Chr, Po, Pn). Olivine has
uniform composition. The amount of SiO2 added to
dunite is allowed to vary in the range 0–07 mol SiO2/
mol Ol (in excess of what is required to react with Ol and
Spl to form Chl). In the first stage of the reaction history,
dunite is hydrated and silicified to Ol–Tlc–Chl schist
using mineral formulas in Table 5. The exact sequence
of reactions, that in nature would have involved
other minerals as intermediate reaction products, is
inconsequential. Amounts of Ol and Tlc in the
model Ol–Tlc–Chl schist depend only on the amount of
SiO2 added; Ol composition was computed from
mass balance of Fe and Mg and the Fe–Mg exchange
constants in Table 5. During the second stage, reaction
(1) proceeds in Ol–Tlc–Chl schist until Ol composition
reaches 0888 Xfo,Ol (Fig. 8). Different values of x1
develop at the end of the second stage of reaction,
o
prior to reaction (1), that,
depending on noOl and Xfo;Ol
in turn, depend on the amount of SiO2 added to dunite.
The range in measured values of x1 along the m-scale
traverse (Fig. 7) is quantitatively reproduced for Xfo,Ol ¼
0924 in the dunite (a representative mantle value) and
NUMBER 8
AUGUST 2005
3.0
ξ 1 for final X fo,Ol = 0.888 (mol/L)
JOURNAL OF PETROLOGY
range in ξ1 for samples
from m-scale traverse
peridotite parent:
Ol+En+Di = 98%
X fo = 0.923
X en = Xdi = 0.930
2.0
1.0
0.0
0.0
0.2
0.4
0.6
0.8
volumetric Px/(Ol+Px) in peridotite
Fig. 15. Progress of reaction (1), x1, predicted by simple two-stage
model for mineral reaction in the metaperidotite body. Assumed parent rock is peridotite composed of 98 modal % Ol þ En þ Di, 1% Al–
Cr Spl and 1% other inert minerals; Px/(Ol þ Px) may vary between 0
and 08; Di/En ¼ 0175 by volume. In the first stage of reaction,
peridotite is stripped of Ca and hydrated to produce Ol–Tlc–Chl schist
o
that depend on Px/(Ol þ Px) and the
with variable noOl and Xfo;Ol
specified Ol composition in the peridotite. In the second stage, reaction
(1) is driven in the Ol–Tlc–Chl schist by infiltration of rock by reactive
CO2–H2O fluid until Ol with Xfo,Ol ¼ 0888 is produced. The range in
measured x1 from the m-scale traverse (sample location 16), 07–
25 mol/l (grey area) is reproduced if Ol in the peridotite parent has
Xfo,Ol ¼ 0923 and Px/(Ol þ Px) is variable in the range 005–050.
Spatial variations in x1 along the m-scale traverse therefore might
simply represent spatial variations in Px/(Ol þ Px) in a peridotite
parent rock.
the amount of SiO2 added ¼ 004–043 mol/mol Ol
(Fig. 16). There could have been variations in Xfo,Ol, as
well as a range in amount of SiO2 added, but these are
not required by the data. For Xfo,Ol ¼ 0924 in the dunite
parent and addition of 004–043 mol SiO2/mol Ol,
o
¼
predicted ranges in noOl ¼ 125–202 mol/l and Xfo;Ol
0902–0920 of Ol–Tlc–Chl schist almost exactly match
those inferred for samples along the m-scale traverse from
measured modes and mineral compositions (124–199,
0901–0920, Fig. 13).
Regardless of whether the carbonated ultramafic rock
evolved from Ca-depleted peridotite, silicified dunite, Cadepleted and silicified peridotite, or something else, the
measured variations in x1 along the m-scale traverse
therefore record nothing about the geometry or distribution of fluid flow during Barrovian regional metamorphism. The variations in x1 simply image spatial variations
in the bulk composition of the altered ultramafic rock,
(Mg þ Fe)/Si in particular, prior to metamorphism.
A variation of the new interpretation
Significant variations in x1 can also be produced by
infiltration of Ol–Tlc–Chl schist by CO2–H2O fluid,
even if all samples contain the identical amount and
1742
CENTIMETRE-SCALE VARIATIONS IN METAMORPHIC REACTIONS
ξ 1 for final X fo,Ol = 0.888 (mol/L)
3.0
range in ξ1 for samples
from m-scale traverse
2.0
dunite parent:
Ol = 98%
X fo = 0.924
1.0
0.0
0.0
0.2
0.4
0.6
moles SiO 2 added per mole olivine in dunite
Fig. 16. Progress of reaction (1), x1, predicted by a second simple twostage model for mineral reaction. Assumed parent rock is a dunite
composed of 98 modal % Ol, 1% Al–Cr Spl and 1% other inert
minerals. In the first stage of reaction, dunite is hydrated and
o
that
silicified to produce Ol–Tlc–Chl schist with noOl and Xfo;Ol
depend on the amount of SiO2 added and the specified Ol
composition in the dunite. Amounts of SiO2 added were considered
in the range 0–07 mol SiO2/mol Ol (in excess of what is needed to
convert Ol þ Spl to Chl). In the second stage, reaction (1) is driven in
the Ol–Tlc–Chl schist by infiltration of rock by reactive CO2–H2O
fluid until Ol with Xfo,Ol ¼ 0888 is produced. The range in measured
x1 from the m-scale traverse (sample location 16), 07–25 mol/l (grey
area) is reproduced if Ol in the unreacted dunite has Xfo,Ol ¼ 0924
and the amount of SiO2 added is 004–043 mol/mol Ol. Spatial
variations in x1 along the m-scale traverse therefore could
alternatively simply represent spatial variations in silicification of a
dunite parent rock.
composition of Ol prior to reaction, provided there is
partial rather than complete homogenization of fluid
composition among the samples. Consider Ol–Tlc–Chl
schist with initial mineral abundances and Ol composition as indicated in Fig. 17. Chemically reactive fluid
infiltrates some reference sample, reaction (1) then proceeds, and XCO2 is buffered by reactants and products.
If there is imperfect CO2–H2O exchange by diffusion–
dispersion between the reference sample and some other
remote sample (e.g. 1–10 m away) through which there is
no fluid flow, a gradient in XCO2 develops and Xfo,Ol will
differ between the rocks. In a steady state involving
imperfect CO2–H2O exchange between the samples,
the difference in XCO2 and Xfo,Ol will be constant, and
reaction (1) will proceed in the remote sample at a
reduced rate. If reaction proceeds in the reference sample
until Xfo,Ol ¼ 0888, Fig. 17 illustrates the calculated
value of x1 in the remote sample as a function of the
steady-state difference in Xfo,Ol. Variations in the degree
of communication result in a range of x1 in the remote
sample between zero (for no communication at all) and
276 mol/l (for perfect communication). Variations in
x1 similar in magnitude to those measured along the
m-scale traverse therefore can be produced by imperfect
ξ 1 (remote sample, mol/L Ol-Tlc-Chl protolith)
FERRY et al.
3
o
X fo = 0.9215
o
nOl
= 21.2 mol/L
o
nChl
= 0.3 mol/L
2
o
n Tlc
=0
o
Vother = 10 cm 3/L
1
Mgs-forming reaction (1)
0
0.00
0.01
0.02
0.03
0.04
X fo,Ol (remote sample) - X fo,Ol (reference sample)
Fig. 17. Progress of reaction (1), x1, predicted by simple model for
imperfect chemical communication between two samples of metaperidotite. Both samples have the same modal mineralogy and mineral
compositions prior to reaction (1), as specified in inset. Calculations
used model mineral formulas and Fe–Mg exchange constants in
Table 5. As reaction (1) proceeds in the reference sample, a steadystate difference in Xfo,Ol is maintained between the reference and a
remote sample. Curve illustrates the value of x1 attained in the remote
sample when Xfo,Ol in the reference sample ¼ 0888 as a function of
the steady-state difference in Xfo,Ol. Variations in x1 as large as those
observed along the m-scale traverse can be produced by imperfect
chemical communication during reaction between rocks that initially
are identical.
chemical communication between samples that are in
every other respect identical. The good evidence for
homogenization of fluid composition along the m-scale
traverse (Figs 8–11) rules out imperfect chemical communication as the explanation for variations in x1
along the traverse. Imperfect communication, along
o
, however, may explain
with differences in noOl and Xfo;Ol
differences in x1 between samples from the m-scale
traverse and more remote samples that record a
difference in Xfo,Ol (¼ XCO2), d 18OMgs (¼ d 18Ofluid) and
d 13CMgs (¼ d13Cfluid).
IMPLICATIONS
Whenever an infiltration-driven reaction involves one
or more mineral reactants that are solid solutions and
fluid composition is homogenized by diffusion–dispersion
over a distance greater than the thickness of lithological
layering, layer-by-layer differences in the progress of the
reaction inevitably develop if the layers differ in the initial
amount and/or composition of the reactant mineral(s).
Specific consideration of the carbonated metaperidotite
body in Val d’Efra demonstrates that the process
can produce variations in reaction progress of a factor
of 26 over several decimetres. Other studies that
1743
JOURNAL OF PETROLOGY
VOLUME 46
interpreted cm-scale variations in reaction progress in
terms of channelized fluid flow (e.g. Ferry, 1987) failed
to adequately consider the significance of solid solutions
and can no longer be considered correct. The results of
this study thus resolve the apparent contradiction
between cm-scale variations in progress of infiltrationdriven reactions and isotopic evidence for homogenization of fluid compositions over a distance of >1 m in the
same outcrop of regionally metamorphosed rock (Ferry,
1987; Bickle et al., 1997). The rate of CO2–H2O interdiffusion during regional metamorphism is very rapid (Wark
& Watson, 2004), and the length scale of homogenization
of fluid composition by diffusion–dispersion appears to be
typically 1 m (Bickle et al., 1997; Ague & Rye, 1999;
Ague, 2000, 2002, 2003; Evans et al., 2002). Therefore,
cm-scale variations in the progress of infiltration-driven
reactions involving solid solutions are better interpreted
in terms of cm-scale variations in the initial amount and
compositions of mineral reactants. Variations in reaction
progress at the m scale (e.g. Ferry, 1994) likewise may
be controlled more by initial variations in modes and
mineral chemistry than by the channelization of reactive
fluid flow during regional metamorphism.
If mineral reactants and products involved in an
infiltration-driven reaction are either pure substances
(e.g. calcite, quartz and wollastonite) or are fixed in composition by mineral equilibria (e.g. calcite and dolomite
during the dolomite–periclase–calcite reaction), amounts
and compositions of reactant mineral(s) exert no control
on reaction progress. In this case, the mapped distribution of the progress of the reaction at the outcrop or
larger scale indeed corresponds to the spatial distribution
of time-integrated fluid flux (e.g. Ferry & Rumble, 1997;
Ferry et al., 1998, 2001, 2002; Lackey & Valley, 2004).
Specifically, high-x regions image the location, size and
geometry of channels for elevated flow, and low-x regions
image intervening regions with reduced flow.
A corollary of this study is that the geometry of fluid
flow can never be determined at a length scale smaller
than the characteristic distance over which fluid composition is completely homogenized by diffusion–dispersion.
Because of the homogenization of geochemical tracers in
the fluid by diffusion–dispersion, it is impossible, for
example, to determine over the characteristic distance
whether the physical mechanism of flow was along a
single thin crack, was pervasive and uniform at the
grain-size scale, or was something in between. The conclusion holds, regardless of whether an infiltration-driven
metamorphic reaction involves solid solutions or pure
substances and regardless of the chemical tracer used to
investigate flow (e.g. reaction progress, stable or radiogenic isotope compositions, trace-element concentrations). Studies that determine the characteristic length
scale of diffusion–dispersion (e.g. Bickle & Baker, 1990;
Bickle et al., 1997; Ferry et al., 2001; Evans et al., 2002;
NUMBER 8
AUGUST 2005
Ague, 2003) are essential contributions because they
define the smallest scale at which the geometry of the
flow system can be determined using geochemical tracers.
This smallest scale may differ depending on the specific
tracer considered (Bickle et al., 1997; Ague, 2003).
New methods for direct inversion of spatial patterns of
mineralogical, isotopic and geochemical alteration in
terms of the regional-scale, 3-D pattern of reactive fluid
flow (Wing & Ferry, 2002, 2005) require representative
outcrop-scale estimates of fluid composition and the
progress of infiltration-driven mineral–fluid reactions.
Results of this study suggest that, with some obvious
counterexamples (e.g. Rumble, 1978), fluid composition
may be as surprisingly uniform at the outcrop scale
during regional metamorphism as predicted by Ague
(2000, 2002). For example, with the exception of extensively serpentinized samples 16L and 16M, the range of
measured Ol compositions for all analysed samples from
the metaperidotite body is Xfo,Ol ¼ 0885–0899 (Table 4).
The range in Xfo,Ol, in turn, corresponds to a range in
XCO2 < 001. If fluid composition indeed is typically
so uniform during regional metamorphism at the outcrop
scale, average fluid composition can be adequately
determined from measurements of one or a few samples
per outcrop. Reaction progress, on the other hand, is
much more variable, and average values correspondingly
are more difficult to measure. The obvious solution of
collecting and analysing several dozen samples per outcrop would be impractical for any regional investigation.
This study suggests a more efficient alternative. If a representative outcrop were intensely investigated at low
grade of metamorphism where infiltration-driven reactions did not occur, variations in the amount and composition of reactant minerals could be measured and
average values along with their uncertainties accurately
determined. At higher grades where reaction occurred in
lithologically equivalent rocks, one or a few samples per
outcrop could be used to calibrate the quantitative relationship between reaction progress and the amount and
composition of reactants prior to reaction. An example
specifically for the carbonated metaperidotite body in
Val d’Efra is illustrated in Fig. 13. An average value of
reaction progress and its range for an outcrop as a whole
then could be inferred from the previously determined
average values and ranges for the initial amounts and
compositions of mineral reactants. Thus, quantitative
studies of the 3-D geometry of reactive fluid flow on the
regional scale may become more tractable.
ACKNOWLEDGEMENTS
We thank Volkmar Trommsdorff and Bernard Evans
for advice on the logistics of fieldwork in the area;
Franco Barera for arranging our stay at Capanna Efra;
1744
FERRY et al.
CENTIMETRE-SCALE VARIATIONS IN METAMORPHIC REACTIONS
the Museo Cantonale di Storia Naturale, Lugano,
Switzerland, for permission to conduct the fieldwork;
and Jay Ague, Bernard Evans, and Craig Manning for
their thoughtful reviews. Sarah Carmichael assisted with
preparation of field photographs. Research supported by
grant EAR-0229267 from the Division of Earth Sciences,
National Science Foundation.
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