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ORIGINAL PAPER
Polyphase strain caps
Tamer Abu-Alam &Kurt Stüwe
Received: 6 October 2010 /Accepted: 22 November 2010 / Published online: 22 December 2010
#Springer-Verlag 2010
Abstract Strain caps are one of a series of microstructures that
typically form during deformation of a softer matrix around
hard objects. However—in contrast to other microstructures
around porphyroblasts, for example pressure shadows—strain
caps are rarely described in the literature. Here we describe
strain caps with particular focus on strain caps associated with
growth of a new phase, not elsewhere present in the
paragenesis. Examples from foliated, amphibolite facies,
metapelitic schists from Alaska, Sinai and Bhutan are
discussed. All examples show chlorite growth exclusively in
strain caps formed around porphyroblasts. Porphyroblasts
around which the strain caps grow are muscovite, staurolite
and garnet, respectively. In all of these examples strain caps
formed synkinematically, but the chlorite grew statically at a
later stage. Three mechanisms can explain the formation of
new phases in the strain cap region: (a) the strain cap region
may have experienced different P-T conditions from the
matrix; (b) the strain cap region has a different effective bulk
composition from the surrounding matrix; (c) fluid flow that
is preferentially focused parallel to the foliation planes causing
only local adjustment to retrograde metamorphism in the
strain cap region. We show that the third hypothesis is the
most preferable mechanism. Indeed, the absence of chlorite
outside the strain cap region allows a quantification of the
amount of fluid that infiltrated the rock. It is shown that for
Bhutan sample about 8.5 mole% more water must have been
added to the rock during fluid infiltration to cause the strain
cap formation.
Introduction
Strain caps are one of a series of microstructures that
typically form during deformation of a softer matrix around
hard objects (e.g. porphyroblasts and-clasts). As such, they
can be used to infer aspects of the stress and strain fields
around the porphyroblasts. Other microstructural elements
around pophyroblasts include their strain- and pressure
shadows, as well as the line that effectively separates the
flow around the porphyroblast from the “eddy-flow”in the
strain shadow: the separatrix (Fig. 1a). Some of these
microstructural elements bear characteristic information on
the vorticity of the stress field and have—therefore—
received a lot of attention in the literature. For example,
the orientation of a pressure shadow with respect to the
flow lines is a direct consequence of the ratio of pure- to
simple shear and can thus be used to estimate the vorticity
(Fig. 1b, c and d). A GEOREF search lists more than 36
entries for the term “pressure shadow”or “strain shadow”
in the title of a paper (e.g. Takagi and Ito 1988; Etchecopar
and Malavieille 1987; Tenczer et al. 2001). In contrast,
strain caps are barely described outside their definition (e.g.
Passchier and Trouw 1996,2005), possibly because strain
caps are considered to merely reflect the denser flow lines
around the porphyroblast without much additional signifi-
cance for the orientation of the stress field. Typically, strain
caps are evidenced by a region that is depleted in quartz
and enriched in micas occurring on opposite sides of the
rigid body, in the quarters orthogonal to the strain shadow.
In one of the few studies that have discussed strain caps,
Trouw et al. (2008) showed this orthogonal relationship
between strain cap and strain shadow may be responsible
for the orthogonal arrangement of spiral inclusion trails in
garnet. They argue that synkinematically grown garnet
typically grows preferentially in direction of the strain cap
Editorial handling: J. Raith
T. Abu-Alam (*):K. Stüwe
Institut für Erdwissenschaften, Universität Graz,
Universitätsplatz 2,
8010 Graz, Austria
e-mail: tamer.abu-alam@uni-graz.at
Miner Petrol (2011) 101:1–19
DOI 10.1007/s00710-010-0141-7
as this region is concentrated in the nutrient cations. As the
porphyroblast continues to elongate in direction of the
strain cap by growth, it eventually rotates into the direction
of the fabric attractor, before continuing to grow on its sides
towards the new strain cap.
In this paper we describe strain caps with the opposite
relationship: Not preferential porphyroblast growth, but
preferential consumption of the porphyroblasts in the region
of the strain cap. This process leads to reaction and
formation of new phases (typically chlorite) that grow
exclusively in the strain cap region. Such a microstructure
can obviously not be explained through deformation alone,
but requires an intergraded interpretation in terms of the
metamorphic reaction in relationship to the deformation.
Here we use detailed microstructural description, petro-
graphic work and thermodynamic modeling to explain
formation mechanism of new phases grown in the strain
cap region. We shall show that these polyphase strain caps
can only form due to post metamorphic peak fluid
infiltration through matrix foliation and we quantify the
amount of water that is required for their formation.
Geological setting
The strain caps of our interest show chlorite growth in the
strain cap region surrounding various types of porphyro-
blasts. Interestingly, we have discovered this feature in a
series of amphibolite facies metamorphic rocks from
different locations. Here, we present our interpretation for
samples from three different regions: Bhutan, Sinai and
Alaska.
The Bhutan samples which will be studied here are
pelitic schist samples from the Higher Himalaya Crystalline
and are part of a suite of samples that were collected by
Stüwe and Foster (2001). The Higher Himalayan Crystal-
line is a high grade complex located at the Himalayan front
between two major crustal scale shear zones, the Main
Simple shear
PS
SS
RB
Strain cap pressure shadow (PS)
strain
shadow
(SS)
rigid
body
(RB)
Simple shear
(a) (b)
Pure shear
SS=PS
RB
(c)
General shear
PS
SS
RB
(d)
Separatrix
S-planes
Fig. 1 Microstructural elements around rigid objects-(a) shows the
names of different structural elements (modified after Passchier and
Trouw 2005). (b), (c) and (d) show the geometric changes of these
elements in response to simple shear, pure shear and general shear,
respectively. Dashed arrows show the eigenvectors of the flow field.
Continuous arrows show the orientations of the instantaneous
stretching axes
2 T. Abu-Alam, K. Stüwe
Central Thrust (MCT) and the South Tibetan Detachment
Zone (STDZ). The MCT is one of the major features
associated with the collision between the Indoaustralian and
the Asian continental plates. It is associated with inverted
metamorphic isograds both above—in Higher Himalaya
Crystalline rocks (amphibolites facies) and below—in
Lesser Himalaya rocks (greenschist facies) (e.g. Gansser
1964; Le Fort 1975; Hodges et al. 1988; Hubbard 1989;
Vannay and Hodges 1996; Grujic et al. 1996). The Lesser
Himalaya rocks are separated from Sub-Himalayan domain
(mollase type deposits) in the south by the Main Boundary
Thrust (MBT). The main movement on the MCT occurred
at about 15–22.5 Ma (e.g. Hubbard and Harrison 1989;
Hodges et al. 1996) and it was synchronous with the
prograde metamorphism (Grujic et al. 1996). The Higher
Himalaya Crystalline rocks show a wide variation of
pressure-temperature conditions (500–800°C and 4.8–
8.5 kbar; e.g. Dasgupta et al. 2004). However, the transect
that was studied by Stüwe and Foster (2001) shows peak
conditions in the range of 600–650°C at 6.5 kbar. Granitic
magmas originated by the melting of the Higher Himalayan
Crystalline pelitic gneisses by infiltration of fluid at the
time of thrusting along the MCT. These magmas intruded
latter in the higher crust rocks as leucogranite bodies (Le
Fort 1981; Vidal et al. 1982; France-Lanord and Le Fort
1988; England et al. 1992).
Sinai samples are from Taba metamorphic complex
which is directly located to the northwest of the Gulf of
Aqaba and consists of metapelitic schists, migmatites,
metagabbro-diorites, orthogneisses, and metamorphosed
mafic dykes (Kröner et al. 1990; Heimann et al. 1995;
Cosca et al. 1999; Abu El-Enen et al. 1999,2004; Eliwa et
al. 2008). Abu El-Enen et al. (2004) identified four
metamorphic zones in the metapelites; garnet, staurolite,
staurolite-sillimanite and staurolite-cordierite zones. The
samples which will be studied here are from the staurolite
zone. The entire metamorphic complex underwent poly-
phase deformation with three ductile phases D
1
-D
3
(e.g.
Shimron 1980) accompanied by a single metamorphic
event with peak condition (590–640°C and 5–6 kbar; e.g.
Abu El-Enen et al. 2004) followed by a brittle D
4
phase.
Cosca et al. (1999) concluded that regional metamorphism
at 620±10 Ma of the complex is related to the collision of
greater east and west Gondwanaland. Subsequent to the
peak metamorphism and during the exhumation, the
metamorphic sequences of Sinai were intruded by post-
tectonic granites and the equivalent volcanic rocks (e.g.
Abu-Alam and Stüwe 2009a) and were affected by fluid
influx (Abu-Alam et al. 2010).
The Alaska samples which will be studied here are
pelitic schists from the rim of the Chugach Metamorphic
Complex. This complex is an Eocene metamorphic com-
plex in southern Alaska related to the subduction of the
Pacific plate underneath the North American continent
(Sisson and Hollister 1988; Sisson et al. 1989). This
complex is bound in the north by the Border Range fault
and in the south by the Contact fault (e.g. Plafker et al.
1994). Hudson and Plafker (1982) were the first to study
the tectonic evolution of this complex. Since this contribu-
tion the Chugach Metamorphic Complex is known to have
symmetric geometry with a core made up of migmatitic
gneiss and margins of schists. The studied samples are part
of a suite that were collected by Bruand et al. (2010).
Sisson and Hollister (1988) suggested around 3 kb as an
average pressure value for the entire complex and temper-
atures of 550°C for the external schist zone and 650°C for
the highest grade core of the complex. However, the recent
study by Bruand et al. (2010) reveals that the complex
shows a significant increase of pressure and temperature
from the north (550°C-3 kbar) to the south (<750°C-
13 kbar). Based on a detailed thermodynamic study of
Bruand et al. (op cit), the lower crustal part of the complex
(migmatitic gneiss) was subjected to H
2
O-saturated melting
process. This melt intruded later in the upper crustal rocks
(schists) to form magmatic intrusion bodies e.g. Sanak-
Baranof belt (Hill et al. 1981; Moore et al. 1983; Barker et
al. 1992; Sisson et al. 2003; Farris and Paterson 2007;
Farris and Paterson 2009).
Microstructural descriptions and metamorphic
conditions
In the chosen three examples, chlorite bearing strain caps
form around different porpyroblasts of garnet, staurolite and
muscovite. The mineral assemblages in the studied samples
were analysed at the Institute of Earth Science, Karl-
Franzens-Universität Graz, Austria, using a JEOL JSM-
6310 scanning electron microscope following standard
procedures, operating in EDS/WDS mode at 5 nA beam
current, accelerating voltage 15 kV and duration time is
100 s. The chemical formulae and end-members activities
were calculated using the program AX (http://www.esc.
cam.ac.uk/research/research-groups/holland/ax). The miner-
al abbreviations which will be used in the following
sections are from Holland and Powell (1998). Pressure-
temperature conditions (Tables 1,2and 3) were calculated
by solving independent sets of reactions between mineral
end-members (Appendix 1) by THERMOCALC 330
(Powell and Holland 1988) and the internally consistent
dataset of Holland and Powell (1998).
Bhutan sample is a typical amphibolite facies garnet
mica schist (Figs. 2a, b and 3a). The sample is composed of
garnet, biotite, white mica, chlorite, quartz and sphene.
Parallel mica and quartz crystals define the metamorphic
foliation, which wraps around the garnet porphyroblasts
Strain caps formation mechanism 3
Table 1 Representative mineral analyses of Bhutan sample
Strain shadow Strain cap
Mineral g4 bi4 mu4 g5 bi5 mu5 g10 bi10 mu10 g7 bi7 chl7
SiO
2
37.75 34.49 47.13 37.49 36.41 45.83 38.46 36.39 46.75 37.22 35.8 23.87
TiO
2
0.14 3.21 0.77 0.22 2.56 0.53 0.15 2.31 0.44 0.24 3.28 0.3
Al
2
O
3
20.4 17.99 36.58 21.05 18.42 34.51 21.22 18.44 35.2 20.77 19.19 21.43
Fe
2
O
3
–––––––––0.51 ––
FeO 33.14 20.75 1.12 36.11 21.5 1.38 35.19 20.82 1.87 33.63 20.46 27.92
MnO 1.12 0.12 0.08 0.71 0.04 b.d.l. 1 0.14 b.d.l. 1.45 b.d.l. 0.17
MgO 1.12 7.91 0.55 1.77 7.86 0.68 1.69 8.82 0.83 1.17 8.39 11.51
CaO 6.15 b.d.l. 0.03 3.43 0.11 0.11 3.63 0.14 0.09 5.66 0.07 0.09
Na
2
O 0.06 0.29 1.51 0.03 0.25 1.18 0.07 0.26 1.07 0.05 0.24 0.07
K
2
O 0.1 9.88 9.96 b.d.l. 9.85 10.08 b.d.l. 10.22 10.29 0.02 10 0.24
Totals 99.98 94.64 97.84 100.83 97.04 94.33 101.44 97.54 96.56 100.73 97.49 85.59
O 121111121111121111121114
Si 3.044 2.682 3.051 3.005 2.75 3.086 3.047 2.733 3.081 2.99 2.682 2.63
Ti 0.008 0.188 0.038 0.014 0.145 0.027 0.009 0.13 0.022 0.014 0.185 0.025
Al 1.939 1.649 2.792 1.989 1.64 2.74 1.982 1.633 2.734 1.967 1.695 2.784
Fe
3+
–––––––––0.031 ––
Fe
2+
2.235 1.349 0.061 2.421 1.358 0.078 2.332 1.308 0.103 2.259 1.282 2.573
Mn 0.077 0.008 0.004 0.048 0.002 –0.067 0.009 –0.099 –0.015
Mg 0.135 0.917 0.054 0.212 0.884 0.068 0.2 0.987 0.082 0.14 0.937 1.891
Ca 0.532 –0.002 0.294 0.009 0.008 0.308 0.011 0.006 0.487 0.006 0.01
Na 0.009 0.043 0.19 0.005 0.036 0.155 0.01 0.038 0.137 0.008 0.035 0.014
K 0.011 0.98 0.823 –0.949 0.866 –0.98 0.865 0.002 0.956 0.033
Sum 7.988 7.818 7.019 7.989 7.777 7.028 7.957 7.829 7.031 8 7.781 9.977
X(chl) = Fe
2+
/(Fe
2+
+Mg
2+
) 0.576
Activity of end-members
py 0.00026 0.00074 0.00069 0.00027
gr 0.0062 0.00124 0.0014 0.0046
alm 0.37 0.5 0.44 0.41
phl 0.0248 0.0261 0.034 0.0272
ann 0.073 0.076 0.07 0.06
east 0.03 0.026 0.033 0.034
mu 0.69 0.71 0.69
cel 0.025 0.017 0.015
fcel 0.028 0.02 0.018
pa 0.8 0.13 0.11
clin 0.0106
daph 0.048
ames 0.0152
Pressure 7.2± 0.8 7.1± 0.5 6±0.7 7.2± 0.5
Temp. 614 ±30 564± 45 513 ±35 557 ± 50
The chemical formulae were calculated based on the number of oxygen atoms listed in the table and ignoring H
2
O for hydrated minerals. Activities of end members were calculated with the
program AX. Average pressure and temperature values were calculated with THERMOCALC. Mineral analyses (mu1* and mu3*) are for mica grains along the chlorite cleavages. b.d.l is below
detection limit.
4 T. Abu-Alam, K. Stüwe
Strain cap
Mineral mu7 g9 bi9 chl9 mu9 g11 bi11 chl11 mu11 mu1* mu3*
SiO
2
46.26 36.75 35.88 24.76 46.71 37.21 36.01 23.79 48.27 44.21 44.38
TiO
2
0.58 0.1 2.33 0.41 0.53 0.38 3.27 0.16 0.85 1.78 1.43
Al
2
O
3
36.12 20.44 19.86 22.7 34.83 20.9 18.03 21.87 32.38 34.53 34.53
Fe
2
O
3
–2.11 –– –0.92 –––––
FeO 1.18 31.72 20.62 27.77 1.51 30.3 20.91 27.66 1.79 1.06 1.08
MnO b.d.l. 1.05 0.17 0.13 0.03 2.86 0.05 0.09 b.d.l. 0.06 b.d.l.
MgO 0.75 1.75 7.99 11.79 0.68 1.33 8.12 11.36 1.37 0.87 0.89
CaO 0.15 5.99 0.08 0.05 0.07 7.08 0.01 0.04 0.02 b.d.l. 0.08
Na
2
O 1.45 0.09 0.26 0.04 1.37 0.04 0.28 0.02 0.84 0.78 0.78
K
2
O 10.21 0.02 10.28 0.12 10.37 b.d.l. 10.01 0.12 10.27 11.04 11.19
Totals 96.75 99.85 97.52 87.76 96.13 100.94 96.71 85.17 95.93 94.35 94.37
O 1112111411121114111111
Si 3.038 2.964 2.692 2.639 3.092 2.969 2.729 2.626 3.196 2.998 3.01
Ti 0.029 0.006 0.132 0.033 0.026 0.023 0.187 0.013 0.043 0.091 0.073
Al 2.797 1.943 1.757 2.853 2.718 1.966 1.611 2.846 2.527 2.761 2.761
Fe
3+
–0.128 –– –0.055 –––––
Fe
2+
0.065 2.14 1.294 2.475 0.083 2.023 1.325 2.553 0.099 0.06 0.061
Mn –0.072 0.011 0.011 0.001 0.193 0.003 0.008 –0.003 –
Mg 0.073 0.211 0.894 1.873 0.067 0.158 0.917 1.869 0.135 0.088 0.09
Ca 0.011 0.518 0.006 0.005 0.005 0.606 0.001 0.005 0.001 –0.006
Na 0.184 0.014 0.038 0.008 0.176 0.006 0.041 0.004 0.107 0.102 0.102
K 0.856 0.002 0.984 0.016 0.876 –0.968 0.018 0.868 0.955 0.969
Sum 7.055 8 7.809 9.914 7.048 8 7.783 9.946 6.983 7.06 7.072
X(chl) = Fe
2+
/(Fe
2+
+Mg
2+
) 0.569 0.577
Activity of end-members
py 0.00096 0.00046
gr 0.0053 0.0093
alm 0.34 0.29
phl 0.0296 0.0256
ann 0.065 0.07
east 0.034 0.028
mu 0.68 0.72 0.63 0.75 0.73
cel 0.034 0.019 0.037 0.019 0.033
fcel 0.03 0.024 0.027 0.013 0.023
pa 0.86 0.14 0.078 0.08 0.578
clin 0.0115 0.0109
daph 0.046 0.051
ames 0.0146 0.0147
Pressure 6.9± 0.4 7.2± 0.6
Temp. 627± 25 600±40
Table 1 (continued)
Strain caps formation mechanism 5
Table 2 Representative mineral analyses of Sinai sample
Strain shadow Strain cap
Mineral bi1 mu1 fsp1 st1 bi2 mu2 fsp2 st2 bi7 chl7 mu7 fsp7 st7 bi9 chl9 mu9 fsp9 st9
SiO
2
34.22 45.31 61.88 27.08 33.91 45.34 62.01 26.29 33.61 24.18 44.49 62.59 27.89 33.61 24.57 44.7 62.68 27.34
TiO
2
3.07 0.97 0.09 1.1 2.37 0.83 b.d.l. 1.26 3.48 0.07 0.71 b.d.l. 1.07 5.22 0.34 1.04 0.05 1.09
Al
2
O
3
19.59 35.11 22.82 51.98 19.82 36.08 23.82 51.05 19.14 21.39 35.85 23.37 53.45 17.43 20.92 34.96 23.6 52.97
Fe
2
O
3
–– –––– ––––– ––––0.08 ––
FeO 17.36 1.11 b.d.l. 12.91 18.79 1.1 b.d.l. 12.04 19.25 25.07 1.09 b.d.l. 12.24 19.18 24.8 2.09 b.d.l. 11.89
MnO 0.1 b.d.l. b.d.l. 0.4 0.1 0.02 0.05 0.38 0.14 0.2 0.03 0.02 0.44 0.13 0.24 0.01 b.d.l. 0.38
MgO 8.52 0.52 0.09 1.61 8.18 0.85 0.26 1.43 8.67 12.61 0.19 b.d.l. 1.36 7.43 13.47 0.9 0.12 1.03
CaO 0.08 0.06 5.29 b.d.l. 0.06 0.13 5.47 0.05 0.04 0.07 0.14 5.13 0.05 0.27 0.06 0.07 5.36 b.d.l.
Na
2
O 0.2 0.93 9.36 b.d.l. 0.11 1.28 9.3 0.04 0.13 b.d.l. 0.99 9.17 0.02 0.16 0.01 0.72 9.54 0.02
K
2
O 10.05 10.34 0.11 b.d.l. 10.28 10.34 0.11 0.01 9.24 0.01 10.68 0.11 b.d.l. 10.09 0.03 10.38 0.11 b.d.l.
Totals 93.24 94.5 99.65 95.08 93.71 95.98 101.06 92.57 93.73 83.59 94.22 100.47 96.54 93.62 84.45 94.98 101.46 94.74
O 11 11 8 46 11 11 8 46 11 14 11 8 46 11 14 11 8 46
Si 2.656 3.047 2.758 7.724 2.641 3.006 2.731 7.684 2.612 2.68 3.009 2.762 7.793 2.639 2.692 3.009 2.745 7.771
Ti 0.179 0.049 0.003 0.237 0.139 0.042 –0.277 0.203 0.006 0.036 –0.224 0.308 0.028 0.053 0.002 0.232
Al 1.792 2.783 1.199 17.479 1.82 2.82 1.237 17.59 1.754 2.795 2.859 1.216 17.606 1.613 2.703 2.774 1.218 17.751
Fe
3+
–– 0.007 ––– 0.003 –––– 0.006 –––0.005 0.008 –
Fe
2+
1.127 0.063 –3.079 1.224 0.061 –2.944 1.251 2.324 0.062 –2.86 1.259 2.273 0.117 –2.828
Mn 0.007 ––0.097 0.007 0.001 0.002 0.095 0.009 0.018 0.002 0.001 0.104 0.009 0.022 0.001 –0.09
Mg 0.985 0.052 0.006 0.684 0.949 0.084 0.017 0.623 1.004 2.083 0.019 –0.566 0.869 2.2 0.09 0.008 0.436
Ca 0.006 0.004 0.252 –0.005 0.009 0.258 0.017 0.003 0.009 0.01 0.242 0.015 0.023 0.007 0.005 0.252 –
Na 0.03 0.121 0.809 –0.017 0.165 0.794 0.023 0.02 –0.13 0.785 0.011 0.024 0.002 0.094 0.81 0.011
K 0.995 0.887 0.006 –1.021 0.875 0.006 0.002 0.916 0.001 0.922 0.006 –1.011 0.005 0.891 0.006 –
Sum 7.781 7.014 5.042 29.3 7.828 7.062 5.049 29.256 7.775 9.916 7.051 5.02 29.183 7.762 9.932 7.042 5.048 29.124
Ca(pl) = Ca/(Ca+Na+K) 0.23 0.24 0.23 0.24
Activity of end-members
phl 0.034 0.0342 0.029 0.0173
ann 0.04 0.054 0.05 0.057
east 0.047 0.045 0.042 0.026
an 0.39 0.4 0.39 0.39
ab 0.76 0.76 0.77 0.76
mst 0.0017 0.0015 0.0013 0.00061
fst 0.4 0.42 0.43 0.51
mu 0.74 0.72 0.79 0.68
cel 0.02 0.014 0.0098 0.0151
fcel 0.024 0.01 0.032 0.02
pa 0.58 0.14 0.68 0.332
clin 0.0194 0.022
daph 0.034 0.027
ames 0.0206 0.0246
Pressure 6± 0.8 6.2± 0.5 6.1±0.3 6 ± 0.7
Temp. 652±30 649± 25 643± 40 635 ±33
The chemical formulae were calculated based on the number of oxygen atoms listed in the table and ignoring H
2
O for hydrated minerals. Activities of end members were calculated with the
program AX. Average pressure and temperature values were calculated with THERMOCALC. b.d.l is below detection limit.
6 T. Abu-Alam, K. Stüwe
Table 3 Representative mineral analyses of Alaska sample
Strain shadow Strain cap
Mineral fsp6 bi6 mu6 fsp7 bi7 mu7 fsp9 bi9 mu9 fsp11 chl11 bi11 mu11 fsp12 chl12 bi12 mu12 fsp14 chl14 bi14 mu14
SiO
2
62.53 34.36 45.88 62.37 34.52 45.27 62.88 34.22 46.21 62.29 24.9 34.37 45.51 60.49 24.13 34.48 45.66 63.59 24.75 34.86 46.15
TiO
2
0.23 4.11 1.62 0.14 4.14 1.35 0.02 4.33 1.63 0.09 0.28 4.38 1.57 0.23 0.07 3.11 1.77 b.d.l. 0.33 4.25 1.88
Al
2
O
3
23.13 18.04 34.35 23.56 18.46 34.95 23.9 17.62 35.99 23.75 21.22 18.39 35.16 24.68 21.28 17.87 35.08 23.22 21.41 18.15 35.32
Fe
2
O
3
– – –– – –– – –– – – –– –– –– –– –
FeO b.d.l. 19.3 1.19 b.d.l. 18.72 0.99 b.d.l. 19.38 1.17 b.d.l. 24.91 19.48 0.93 b.d.l. 24.42 18.71 1.19 b.d.l. 24.45 19.27 1.34
MnO b.d.l. 0.3 b.d.l. 0.05 0.38 b.d.l. 0.06 0.42 0.04 b.d.l. 0.59 0.44 0.1 0.05 0.61 0.42 0.07 0.03 0.71 0.42 0.02
MgO b.d.l. 8.46 1.1 0.22 8.62 0.9 0.01 8.3 0.87 b.d.l. 13.36 8.11 0.66 0.06 12.98 9.17 0.77 b.d.l. 13.5 8.77 1.08
CaO 5 0.04 0.19 5.36 0.1 0.04 5.54 0.06 b.d.l. 5.71 0.04 0.04 0.04 6.97 0.02 0.15 b.d.l. 5.1 b.d.l. 0.07 0.14
Na
2
O 9.47 0.11 0.73 9.64 0.06 0.73 9.23 0.05 0.77 9.06 0.04 0.12 0.71 8.44 0.03 0.05 0.72 9.59 0.04 0.1 0.82
K
2
O 0.12 10.38 11.09 0.13 10.4 11.12 0.13 10.27 11.22 0.11 b.d.l. 10.32 11.13 0.06 0.04 10.47 11.06 0.21 0.13 10.5 10.9
Totals 100.48 95.17 96.18 101.52 95.41 95.36 101.78 94.67 97.92 101.02 85.39 95.68 95.84 100.99 83.57 94.48 96.36 101.73 85.36 96.4 97.65
O 8 11 11 8 11 11 8 11 11 8 14 11 11 8 14 11 11 8 14 11 11
Si 2.76 2.649 3.048 2.737 2.646 3.029 2.747 2.656 3.013 2.739 2.7 2.637 3.029 2.676 2.674 2.674 3.026 2.774 2.682 2.652 3.018
Ti 0.008 0.238 0.081 0.005 0.239 0.068 0.001 0.253 0.08 0.003 0.023 0.253 0.079 0.008 0.006 0.181 0.088 –0.027 0.243 0.092
Al 1.204 1.64 2.69 1.219 1.668 2.757 1.231 1.612 2.766 1.231 2.713 1.663 2.759 1.287 2.78 1.634 2.741 1.194 2.735 1.628 2.722
Fe
3+
0.009 ––0.002 ––0.001 ––0.007 –– –0.001 –– –0.007 –– –
Fe
2+
–1.244 0.066 –1.2 0.055 –1.258 0.064 –2.259 1.25 0.052 –2.263 1.213 0.066 –2.216 1.226 0.073
Mn –0.019 –0.002 0.025 –0.002 0.027 0.002 –0.054 0.029 0.006 0.002 0.057 0.028 0.004 0.001 0.065 0.027 0.001
Mg –0.972 0.108 0.014 0.985 0.09 –0.96 0.085 –2.159 0.928 0.066 0.004 2.144 1.06 0.076 –2.18 0.994 0.105
Ca 0.236 0.003 0.014 0.252 0.008 0.003 0.259 0.005 –0.269 0.005 0.004 0.003 0.33 0.003 0.012 –0.238 –0.006 0.01
Na 0.81 0.016 0.094 0.82 0.009 0.094 0.782 0.008 0.097 0.773 0.008 0.018 0.092 0.724 0.006 0.008 0.093 0.811 0.009 0.014 0.104
K 0.007 1.021 0.94 0.007 1.017 0.95 0.007 1.017 0.933 0.006 –1.01 0.945 0.004 0.006 1.036 0.935 0.012 0.017 1.019 0.909
Sum 5.034 7.809 7.043 5.06 7.795 7.047 5.031 7.798 7.04 5.028 9.923 7.793 7.031 5.036 9.936 7.849 7.03 5.037 9.935 7.809 7.036
Activity of end-members
an 0.37 0.39 0.41 0.43 0.51 0.37
ab 0.77 0.77 0.75 0.74 0.69 0.77
mu 0.68 0.73 0.72 0.75 0.74 0.68
cel 0.04 0.032 0.025 0.027 0.017 0.019
fcel 0.025 0.02 0.019 0.021 0.014 0.013
pa 0.545 0.506 0.454 0.506 0.074 0.078
phl 0.026 0.028 0.025 0.0229 0.037 0.028
ann 0.057 0.05 0.059 0.056 0.056 0.055
east 0.035 0.038 0.032 0.033 0.042 0.036
clin 0.021 0.021 0.022
daph 0.027 0.029 0.025
ames 0.0226 0.023 0.024
Pressure 4± 0.7 3.7±0.4 3.9±0.2 4.8 ±0.8 4± 0.4 4±0.6
Temp. 326±30 351 ±45 391±25 393± 20 417±35 364 ± 43
The chemical formulae were calculated based on the number of oxygen atoms listed in the table and ignoring H
2
O for hydrated minerals. Activities of end members were calculated with the
program AX. Average pressure and temperature values were calculated with THERMOCALC. b.d.l is below detection limit.
Strain caps formation mechanism 7
forming the strain caps. The white mica has a chemical
composition between muscovite and paragonite and occurs
as 0.5×2 mm crystals. Biotite forms subhedral crystals
(0.7×5 mm); compositionally it is between annite, phlog-
opite and eastonite. Quartz shows undulose extinction and
lies interstitially between muscovite, biotite and garnet
crystals. Garnet occurs as equidimensional grains (6×
7 mm) and is with almandine rich. It has spiral inclusion
patterns of quartz, sphene and biotite. Garnets usually have
idiomorphic crystal faces on the sides where they touch the
strain shadow but serrated edges where they are in contact
with the strain caps (Fig. 2a and b). There is a slight
variation in the FeO content between the region where the
garnet rim touches that strain shadow (33.14–36.11 wt.%)
and the region where it is in contact with the strain cap
(30.30–33.63 wt.%) (Table 1and Fig. 4a). Compositional
zoning and shape of the grain surfaces give the garnets the
appearance that their rims were consumed subsequent to the
formation of a concentric zoning pattern (Fig. 3a). The strain
shadows are plagioclase-free areas, which are composed
mainly of quartz with very small amounts of biotite and
muscovite. Chlorite grains that exclusively grow in the strain
caps are compositionally ripidolite with X(chl) = Fe
2+
/
(Fe
2+
+Mg
2+
) = 0.569–0.577 (Table 1). Chlorites often
contain small mica relics which are aligned parallel to the
metamorphic foliation (Fig. 2b). The relic grains have the
same pleochroism as the mica grains in the foliated matrix.
In addition chlorites contain small grains of muscovite with
bi relic
b
d
a
c
Fig. 2 Photomicrographs of different examples of polyphase strain
caps. aExample from the central Himalayan metamorphic complex of
Bhutan. Note that the spiral inclusion pattern within the garnet
porphyroblast is partly consumed during the formation of the strain
cap phase (see also Fig. 3). bBiotite relic in the strain cap region of
Bhutan sample. The chlorite is found only in the strain cap region
while the strain shadows are free of the micas. cExample from the
Taba metamorphic complex of Sinai. The biotite lies in a direct
contact with the staurolite porphyroblast at the strain shadow of Sinai
sample but there is no chlorite between them. The staurolite
porphyroblast is consumed by chlorite at the strain cap and by
plagioclase at the strain shadow dExample from the Chugach
metamorphic complex of southern Alaska. Note the biotite at left
side of the porphyroblast is in direct contact with the muscovite but
without any chlorite growth. The upper face of the muscovite
porphyroblast is a convex face which may indicates that the muscovite
starts to be consumed at the edge between two faces
8 T. Abu-Alam, K. Stüwe
different pleochroism than the mica in the matrix. These
small mica grains which are found along the chlorite
cleavages parallel to crystal direction (001) have different
chemistry (e.g. SiO
2
,TiO
2
and K
2
O) than the mica in the
foliated matrix (Table 1). Careful comparative geothermo-
barometry between the strain cap region and the matrix
based on solving a set of reactions between mineral end-
members (Appendix 1a) showed that there is no local
pressure variation between both regions within the error of
the method. Both regions give P-T conditions around 6–
7.2 kbar and 513–650°C (Table 1), corresponding to the
estimates of Stüwe and Foster (2001).
The chosen sample from Sinai is a mica schist with
coarse staurolite porphyroblasts and a matrix of biotite,
muscovite, plagioclase, quartz and ilmenite (Fig. 2c). The
metamorphic foliation is defined by parallel crystals of
biotite, muscovite, plagioclase, quartz and ilmenite. The
foliation is wrapped around the staurolite porphyroblasts.
Both biotite and muscovite form crystals 0.05 ×0.6 mm in
size with brownish and greenish white color, respectively.
Quartz and plagioclase are anhedral to subhedral elongated
crystals (0.2×0.5 mm). The plagioclase is oligoclase with
Ca(pl) = Ca/(Ca+Na+K) = 0.21–0.25 (Table 2). The strain
caps form around subhedral to euhedral (1×2 mm in size)
staurolite porphyroblasts with cruciform twinning. The
porphyroblasts are consumed by the chlorite growth at the
strain cap region and by the plagioclase growth at the strain
shadow region (Fig. 2c). The staurolite crystals contain
inclusions of quartz, plagioclase and ilmenite. Minor biotite
and muscovite can also be present as inclusions in
staurolite. Long axes of the porphyroblasts are usually
more or less parallel to the metamorphic foliation (Fig. 2c).
Chlorite appears only in the strain caps and has a ripidolite
composition. Chlorite cleavages are parallel to the meta-
morphic foliation while sub-microscopic ilmenite crystals
grow perpendicular to these cleavages. The strain shadow
regions are characterized by depletion in micas and
enrichment in quartz and plagioclase. In few cases, the
biotite of the strain shadow is in direct contact with the
staurolite porphyroblasts (Fig. 2c). This observation is
crucial for our interpretations below. Thermobarometry
(Table 2and Appendix 1b) indicates that this sample
reached peak metamorphic conditions at pressures around
5.8–6.3 kbar and temperatures between 620 and 660°C with
no difference in the conditions between the strain caps and
the strain shadows.
Alaska sample contains coarse muscovite porphyroblasts
embedded in a matrix of biotite, chlorite, plagioclase,
quartz and apatite. The muscovites are colorless crystals
with very weak pleochroism. Two types of muscovite were
recognized; large euhedral porphyroblasts and small sub-
hedral to euhedral crystals in the metamorphic foliation.
The foliation is wrapped around the muscovite porphyro-
blasts. Small muscovite crystals (1 × 3 mm), biotite, quartz
and plagioclase define the metamorphic foliation. Biotite is
(1.5×5 mm) subhedral to euhedral crystals with perfect
cleavages. Plagioclase forms colorless crystals attaining 1 ×
2.5 mm with albite and albite-Carlsbad twinning. Quartz
forms anhedral to subhedral equant crystals (1.2 ×2 mm)
Fig. 3 Cartoons showing the main microstructural features of the
strain caps in the Bhutan and Alaska examples. cTiming relationships
between the deformation and the strain cap mineral growth as
interpreted from the microstructural description
Strain caps formation mechanism 9
and it is characterized by wavy extinction. The muscovite
porphyroblasts attain 5×18 mm in size. They have minor
amount of quartz, plagioclase, muscovite and biotite as
inclusions. The porphyroblasts have a random orientation
within the ground matrix but they are constantly oblique to
the metamorphic foliation (Fig. 2d). Crystal faces of the
muscovite porphyroblasts which are against the strain cap
are always convex (Fig. 2d and 3b). The strain shadows are
composed of small muscovite crystals, biotite, plagioclase
and quartz. The micas in the strain shadows lie in direct
contacts with the muscovite porphyroblasts. The strain caps
have the same mineralogical composition as the strain
shadow, with the exception of the presence of chlorite with
ripidolite composition. The chlorite grains are medium (0.3×
1.5 mm), euhedral yellowish green crystals. The chlorites
have biotite inclusions parallel to the chlorite cleavage. The
strain cap and the strain shadow areas have the same range of
pressure-temperature conditions (Table 3and Appendix 1c)
in the range 3.7–4.8 kbar and 326–417°C.
Interpreted timing relationships
For all three of the samples described above we infer
similar timing relationships between metamorphism, defor-
mation and the timing of formation of the strain caps. In the
Bhutan and Sinai samples, porphyroblasts are interpreted to
have grown syn-tectonically: In the Bhutan sample this is
evidenced by spiral inclusion trails in garnet (Fig. 3a), in
the Sinai samples staurolite is aligned in the foliation, but is
itself undeformed. In both examples, strain caps formed
after prophyroblast growth as the foliation is wrapped
around the porphyroblast without penetrating them. Simi-
larly, the strain caps in the Alaska sample also formed after
porphyroblast growth, but the porphroblast itself formed
prior to deformation: Muscovite porpyroblasts in the
samples from Alaska are statically grown and rotated by
the subsequent deformation (Fig. 3b). However, small
muscovite crystals in the matrix apparently grew syn-
tectonically. In all three samples, the absence of any
deformation features in the chlorites in the strain cap region
indicates that this mineral grew later as a post-tectonic
phase (Fig. 3c). In the Alaska and Sinai samples, the strain
cap chlorite appears to statically replace biotite and
muscovite pseudomorphically. In the samples from Bhutan
chlorite in the strain cap regions also partially replaces
garnet. The convex faces of the muscovite against the strain
cap (Fig. 2d and 3b) may indicate that muscovite
consumption started at the edges of the porphyroblast. In
Bhutan samples, the difference in the pleochroism and the
chemistry (mu1* and mu3*; Table 1) between the small
mica grains (along the chlorite cleavage) and the mica in
the foliated matrix may indicate that the mica grains along
the cleavages have a different origin than the relic grains
and the mica in the matrix. This observation is crucial for
our discussion below.
Strain cap formation mechanism
There are at least three mechanisms that can explain the
formation of new phases in the strain cap region: (a) the
strain cap region may have experienced different P-T
conditions from the matrix during the peak metamorphism,
for example due to shear heating or local pressure
gradients; (b) the strain cap region has different effective
bulk composition from the surrounding matrix; (c) fluid
flow that is preferentially focused parallel to the foliation
planes causing only local adjustment to retrograde meta-
morphism in the strain cap region.
The first mechanism is easily excluded, because the new
phase grown in the strain cap region (i.e. the chlorite) has a
post-tectonic origin and grew later than the peak assem-
blage (Fig. 3c). Moreover, careful thermobarometry showed
that the peak metamorphic conditions are uniform through-
out the rock (Tables 1,2,3and Appendix 1). The second
mechanism is potentially viable as petrographic observa-
tions and element distribution maps (e.g. Fig. 4a) show that
there are indeed differences in the local bulk composition
between the strain cap regions and the strain shadow
region: Most micas are in contact with the porphyroblasts
only in the strain cap region so that the nutrients for the
chlorite forming reaction are all present there. However, in
two of the three investigated examples (Alaska and Sinai),
micas are also observed in direct contact with the
porphyroblasts in the strain shadow region (circled in
Figs. 2c, d and 3b). Thus, the effective bulk composition
required to form the chlorite is also locally presents in the
strain shadow region—where chlorite never grows. Thus
we exclude the local variation in effective bulk composition
also as a formation mechanism to nucleate the strain cap
mineral. However once the strain cap mineral had nucleat-
ed, the effective bulk composition plays an important role
Fig. 4 Thermodynamic pseudosections used for the interpretation of
the strain cap formation. aElement distribution map showing the
change in the effective bulk composition (e.g. FeO) around garnet
porphyroblast. bP-T pseudosection of Sinai sample for average pelitic
bulk (mole%); SiO
2
: 73.18, Al
2
O
3
: 15.04, MgO: 6.53, FeO: 2.62,
K
2
O: 2.39. cIsobaric T-MH
2
O pseudosection for the Sinai sample at
6 kbar for the same bulk composition of (b) and H
2
O: 5.7–10.2. dP-T
pseudosection for the Alaska sample. The bulk composition (mole%)
is SiO
2
: 68.00 Al
2
O
3
: 10.07, MgO: 6.09, FeO: 13.34, K
2
O: 2.5. eP-T
pseudosection for the strain cap region (Table 4) Bhutan sample. fP-T
pseudosection for the strain shadow region (Table 4) Bhutan sample. g
T-MH
2
O pseudosection for the Bhutan sample at a pressure of 7 kbar
for the same bulk composition of (e; strain cap composition) and H
2
O:
1.18–10.62 (mole%). For all pseudosections the white arrows are the
paths which will produce chlorite
10 T. Abu-Alam, K. Stüwe
at p=7kbar
200 µm
~ 37 28 0 %
FeO
a
bi mu st
bi mu st chl
610 615 620 625 630 635
5
6
7
P
kbar
TC
KFMASH (+H O +q)
2
(b) Sinai
P
kbar
TC
KFMASH (+H2O +qz)
4
5
6
7
mu chl
bi muchl st
bi chl
st
bi chl st g
bi chl
cd g
bi chl st cd
bi cd
g
300 350 400 450 500 550
3
0.38
0.37 0.35
0.33
0.31
0.39
bi mu chl g st
bi mu chl g
bi mu chl
chl mode
0.36
(d) Alaska
Mole% HO2
TC
bi mu st chl H O
2
bi mu st H O
2
bi mu st
bi mu st chl
5.7 6.6 7.5 8.4 9.3
610
620
630
P=6 kbar
KFMASH (+q)
(c) Sinai
bi mu g chl H O2
bi mu g H O2
bi mu g
bi mu g chl
NCKFMASH (+q)
500 550 600 650
4
5
6
7
8
4
5
6
7
kbar
bi mu g q H O
2
bi mu g cd q H O
2
chl bi mu g q HO
2
chl bi mu g q
chl ctd bi
mu g q
bi mu g q H O
2
bi mu g cd q H O2
NCKFMASH
chl bi mu g cd q H O2
0.32
0.36 0.28
0.22
0.24
0.26
0.28
chl bi mu g q H O
2
0.1 0.12
0.05 0.09
g
bi
mu
bi mode
g mode
mu mode
H O mode
2
(e) Bhutan (strain cap)
(f) Bhutan (strain shadow)
TC 1.18 3.54 5.9 8.26
560
580
600
620
TC
Mole% HO2
(g) Bhutan (strain cap)
bi st
g
bi cd
g st
HO
2
Strain caps formation mechanism 11
to grow these nucleuses so the growth of strain cap mineral
will be discussed in more detail in the section of element
diffusion below. Finally, the third mechanism—namely
focused fluid flow along foliation planes—may be a
possible model to explain the observations. In order to test
this hypothesis we employ modeling using thermodynamic
pseudosections.
Thermodynamic modeling
For our modeling we employ thermodynamic pseudosec-
tions. We construct thermodynamic pseudosections for
several bulk compositions of the rocks described above to
constrain the conditions of fluid flow around the porphyr-
oblasts. As the strain caps of interest are a microstructure of
millimeter scale and all occur in “normal”pelitic rocks, the
XRF bulk chemistry will be not used here. Average pelitic
bulk composition is used for Alaska sample (e.g. after
Shaw 1956; Ague 1991; Mahar et al. 1997; Keller et al.
2005). For the Sinai sample we use the average bulk of Abu
El-Enen (1995). These two samples (Alaska and Sinai) will
be studied qualitatively. For more quantitative discussion
the effective bulk composition of the Bhutan sample were
calculated by MBC
1.7
(Abu-Alam and Stüwe 2009b,c) for
both strain cap and strain shadow regions based on the
volume proportions and the chemical composition of the
phases (Table 4). The pseudosections were constructed
using THERMOCALC 330 (Powell and Holland 1988) and
the internally consistent dataset of Holland and Powell
(1998). The following a-x models were used: biotite and
garnet (White et al. 2007); muscovite (Coggon and Holland
2002); cordierite and staurolite (Holland and Powell 1998)
and chlorite (Holland et al. 1998). For Alaska and Sinai
samples we constructed a P-T pseudosection (assuming the
rocks were hydrated during their entire metamorphic
evolution, i.e. with H
2
O in excess) and a T-Mole% H
2
O
(henceforth called T-MH
2
O) pseudosection at the pressure
of interest to investigate changes in the hydration regime
during cooling.
For the Sinai sample, the P-T and T-MH
2
O pseudosec-
tions were constructed in the system KFMASH (Fig. 4b
and c). The P-T pseudosection is characterized by presence
of two fields: the divariant field (bi-mu-st-chl) at lower
temperature condition and the trivariant field (bi-mu-st) at
higher temperature condition showing that chlorite forma-
tion due to cooling is possible. At high H
2
O concentration
of the T-MH
2
O pseudosection (Fig. 4c, constructed at
pressure 6 kbar; Table 2), the divariant field (bi-mu-st-chl-
H
2
O) separates the trivariant assemblage (bi-mu-st-H
2
O) at
the higher temperature side from the assemblage (bi-mu-st-
chl) at the lower temperature side. The chlorite-free
assemblage (bi-mu-st) is stable at low water content
(mole% H
2
O: 5.7–7) and chlorite and ultimately also
chlorite+water bearing assemblages become stable at
successive stages of hydration. Thus, the chlorite in the
Sinai samples can also form due to both processes: cooling
and/or hydration. A P-T pseudosection for the sample
from Alaska gives a similar conclusion: Although topo-
logically somewhat different, the P-T pseudosection
(Fig. 4d) shows that chlorite will grow on the expense of
Table 4 The effective bulk composition of Bhutan sample for the strain cap and strain shadow regions as calculated by Modal Bulk Composition
(MBC
1.7
) program (Abu-Alam and Stüwe 2009b,c)
strain cap strain shadow
Mineral g bi mu chl qz g bi mu qz
Volume (Mol.%) 17.79 17.96 26.55 12.03 25.67 5.22 1.25 0.49 93.02
SiO
2
37.06 35.89 47.08 24.14 100 37.9 35.76 46.57 100
TiO
2
0.24 2.96 0.65 0.29 –0.17 2.69 0.58 –
Al
2
O
3
20.70 19.02 34.44 22.00 –20.89 18.28 35.43 –
Fe
2
O
3
1.18 ––– –– –– –
FeO 31.88 20.66 1.49 27.78 –34.81 21.02 1.45 –
MnO 1.78 0.07 0.01 0.13 –0.94 0.1 0.03 –
MgO 1.41 8.16 0.93 11.55 –1.52 8.19 0.68 –
CaO 6.24 0.05 0.08 0.06 –4.40 0.08 0.07 –
Na
2
O 0.06 0.26 1.22 0.04 –0.05 0.26 1.25 –
K
2
O 0.01 10.09 10.28 0.16 –0.03 9.98 10.11 –
H
2
O–2.76 3.73 13.82 ––3.59 3.75 –
Eff. bulk comp. Wt% Mol% Wt% Mol%
SiO
2
37.83 34.54 74.15 76.31
Al
2
O
3
23.74 12.77 8.75 5.3
Fe
2
O
3
0.17 0.06 0 0
FeO 16.07 12.27 11.64 10.01
MnO 0.31 0.24 0.26 0.23
MgO 4.96 6.75 1.38 2.11
CaO 0.95 0.93 1.17 1.29
Na
2
O 0.49 0.43 0.08 0.08
K
2
O 6.14 3.58 1.5 0.98
H
2
O 9.34 28.44 1.07 3.67
The mineral chemistry is average of Table 1.
12 T. Abu-Alam, K. Stüwe
muscovite with cooling and a similar conclusion is true for
hydration (not shown).
The P-T pseudosection for the Bhutan sample is
modeled in the chemical system Na
2
O-CaO-K
2
O-FeO-
MgO-Al
2
O
3
-SiO
2
-H
2
O (NCKFMASH). Figures 4e, g,5
and 6show the mineral equilibria with bulk composition
corresponding to the bulk at the strain cap region (Table 4)
while the lower part of the P-T pseudosection (Fig. 4f)
shows the stability fields of the mineral assemblages with
respect to the bulk composition at the strain shadow. It may
be seen that a steep, near isothermal line separates chlorite-
bearing assemblages at low temperatures from chlorite-free
assemblages at high temperature. The temperature position
of this line changes concerning the change in the bulk
composition. At the strain cap, the boundary between
chlorite-free and chlorite-bearing assemblages is around
610°C while this temperature is shifted to about 570°C at
the strain shadow (Fig. 4e). Absence of chlorite at the strain
shadow region indicates that the Bhutan rock did not pass the
boundary between the chlorite-free and the chlorite-bearing
assemblage (570°C) during the retrograde path. Garnet,
biotite, muscovite and H
2
O modes were added to the field
(chl-bi-mu-g-q-H
2
O). These modes show that garnet, biotite
and H
2
O stability decreases with temperature decrease
(Fig. 4e) while the muscovite modes increase with temper-
ature decrease. This explains the presence of small musco-
vite grains along the chlorite cleavage planes which have
different pleochroism than the mica in the foliated matrix. In
addition the decreasing in the H
2
O mode indicates that the
chlorite growth requires water consuming. By these modes,
the strain cap formation reaction can be written in the system
NCKFMASH+q as g+bi+H
2
O=chl+mu. The transition from
g-bi-mu assemblage to chl-bi-mu assemblage can be seen in
qualitative KFM (K
2
O, FeO, MgO) compatibility diagrams
(Fig. 5). The KFASH+mu+q+H
2
Ogrid(Fig.5)simplifies
the strain cap formation reaction to g+bi=chl (the thick line).
The pseudosections suggest that one interpretation for the
chlorite growth is simply that it grows due to cooling (arrow
on Fig. 4e) and that its formation is largely independent of
pressure.
To show the behavior of the Bhutan rock during
hydration, a T-MH
2
O pseudosection was calculated in the
NCKFMASH system at a pressure of 7 kbar and bulk
composition representing the strain cap region (Fig. 4g).
The assemblage (bi-mu-g) is stable at low water content
and over a wide range of temperatures (560–640°C). The
chlorite bearing field bi-mu-g-chl is stable at low temper-
ature condition (560–612°C) and H
2
O content above 4.72
mole%. At the same H
2
O range but with much higher
temperature (>612°C), the assemblage bi-mu-g-H
2
O
becomes stable in the water saturated region. The assem-
blage bi-mu-g-chl-H
2
O occurs at temperature range of 570–
612°C and H
2
O range of 4.72–10.62 mole%. The pseudo-
section suggests that isothermal hydration is a second
possibility for the formation of the chlorite in the strain
caps.
In summary, there is apparently no conclusive interpre-
tation possible if the strain cap forming process is due to
cooling or fluid infiltration. However, there is one obser-
vation that allows to rule out cooling: In two of the three
chosen examples biotite is in contact with the porphyroblast
without forming chlorite in a region outside the strain cap
(Fig. 2c, d and 3b). In the samples from Sinai and Alaska,
occasional biotite growth in the strain shadow region
illustrates that the effective bulk composition is locally the
same in the strains shadow and in the strain cap region. As
there is no chlorite growing in the strain shadow—and the
temperature history is undoubtedly the same for the entre
thin section—we argue that cooling can be excluded as a
unique formation mechanism of the strain cap chlorite.
Thus fluid infiltration appears to be the dominant mecha-
nism and we will discuss below if the amount of fluid can
be quantified.
The role of fluid
Many authors discussed the role of the water during
metamorphism (e.g. Guiraud et al. 2001). However, few
authors have been able to quantify the amount of water that
has infiltrated a given rock (e.g. Tenczer et al. 2006). Here,
we suggest that the strain caps described above provide a
unique opportunity to quantify the amount of fluid that has
infiltrated the rock and how it varies around the micro-
structure described here. Figure 4c and g show that chlorite
bearing fields can be reached from the chlorite-free fields
simply by increasing the amount of water beyond 7 mole%
and 4.72 mole%, respectively. However, both the chlorite
absent field bi-mu-g and the chlorite present field bi-mu-g-
chl are in the water undersaturated region of the phase
diagram (Fig. 4g). Thus, it is difficult to see how fluids
infiltrating the rock along grain boundaries could have
reached the strain cap region without causing water
undersaturated reaction further afield in the rock. We
therefore suggest that the water infiltration must have been
sufficient to bring the rock from the chlorite absent region
bi-mu-g to the water saturated region bi-mu-g-chl-H
2
O.
This fluid infiltration by continuous metamorphic reaction
allows to quantify the amount of fluid and, in fact, allows to
infer a complete evolution for the water content of the rock.
Figure 6shows the T-MH
2
O pseudosection (at 7 kbar)
for the spectacular strain caps of the Bhutan sample
(Table 4) discussed above. This diagram can be used to
illustrate the near-peak prograde heating path. Prograde
dehydration during heating results in water loss and a path
(black thick arrow in Fig. 6) that tracks along the water
saturation line, because only a small proportion of free
Strain caps formation mechanism 13
500 550 600 650 700
NCKFMASH
Bhutan (strain cap)
Kbar
chl bi mu g q H O2
bi mu g q H O2
bi mu g cd q H O2
chl bi mu g cd q H O2
7
6
5
g
chl
ctd
chl g
bi ctd
ky st
ctd st
g
sill
ky
KFASH (mu+q+H O)2
5
kbar
500 600
550
6
7
FM
K
mu
chl
gbi
bulk composition
550 c
7kbar
FM
K
mu
chl
g
bi
bulk composition
650 c
7kbar
KFMASH
+q+st+ctd+H O
2
KFMASH
+q+st+ctd+H O
2
Fig. 5 Enlargement of Fig. 4e.
The inserts show KFASH (mu
+q+H
2
O) grid to show the reac-
tion g+bi=chl (thick line) and
two qualitative compatibility
diagrams above and below the
reaction. The compatibility dia-
grams are projected from quartz,
staurolite, chloritoid and H
2
O
onto the plane K
2
O-FeO–MgO.
Filled circles show the bulk
composition of the strain cap
region
6.49 8.86
500
520
540
560
580
600
620
bi mu g chl H O
2
bi mu g H O
2
bi mu g
Mole%
HO
2
bi mu g chl
bi mu g pa bi mu g pa chl
bi mu g pa
chl H O
2
P=7kbar
x(chl)
NCKFMASH (+q)
Bhutan (strain cap)
** Fig. 7a
0.59
0.57
0.55
0.51
0.47
Strain
shadow
Fig. 7b (ii)
Strain cap
Fig. 7b (i)
13.58 15.94 18.30
11.22
T C
Fig. 6 T-MH
2
O pseudosection
for the strain cap region of
Bhutan sample at 7 kbar and for
the same bulk composition of
Fig. 4e and H
2
O: 6.49–20.66
(mole%). The black arrow is
heating-dehydration path up to
the asterisk followed by cooling
path then hydration one. The (i)
and (ii) are references to Fig. 7
14 T. Abu-Alam, K. Stüwe
water is likely to remain along grain boundaries (Guiraud et
al. 2001) on traversing a series of phase fields with
progressive dehydration, ultimately the peak assemblage
bi-mu-g-H
2
O. Following Stüwe and Foster (2001) and our
own estimates, we assume that the metamorphic peak was
reached within this field at around 630°C (as marked by the
asterisk). Onset of cooling will “freeze”the water content
of the assemblage to that of the metamorphic peak as the
remaining free fluid is used up with incipient re-hydration.
As such, the water content of the rock at the peak
conditions is likely to have been just below 9.8 mole%
H
2
O. Subsequent cooling will preserve this water content
(black down-temperature part of arrow in Fig. 6). At some
stage during the cooling evolution fluid infiltration occurred
and the dry assemblage was hydrated. In order to reach the
fluid bearing field bi-mu-g-chl-H
2
O, this hydration event
could have occurred anywhere between 510°C and 610°C.
The temperature at which this occurred can be closer
constrained by the composition of the chlorite in the strain
cap region: Contouring the relevant assemblage for X(chl)
isopleths shows that the hydration is likely to have occurred
around 550°C. This temperature constrains the additional
amount of fluid infiltration to be about 8.5 mole% on top of
the 9.8% of water locked in the solid phase assemblage.
However, this fluid infiltration can only have occurred in
the strain cap region, as the effective bulk composition
defined by the contact paragenesis biotite-muscovite-
garnet-quartz also occurs in other microstructural geome-
tries without reaction to chlorite (Fig. 6and Table 1). Such
a strong gradient in fluid presence can only occur if the
fluid flow occurred preferentially along the foliation only
(Fig. 7). Once the fluid is added to the strain cap region and
the chlorite starts to nucleate, elements start to diffuse from
and to the strain cap region to allow chlorite growth.
Elements diffusion process that occurs parallel to the
chlorite growth is discussed below.
Elements diffusion
In the following section we will interpret the evolution of
the strain cap microstructure using phase equilibria explic-
itly involving chemical potentials to show the diffusion of
elements from and to the strain cap region. Quantitative μ-μ
diagrams were calculated using THERMOCALC 330 and
the internally consistent data sets of Holland and Powell
(1998). Chemical potential relationships around the reaction
g+bi=chl (the P-T grid of Fig. 5) are illustrated in the
simpler chemical system KFMASH+quartz+H
2
O for the
minerals biotite, muscovite, garnet and chlorite. In this
system the main controlling chemical potential is likely to
be μ
K2O
(White et al. 2008). In addition to K
2
O, there will
also be chemical potential gradients in MgO and FeO. The
presence of quartz and H
2
O as excess phases fixes μ
SiO2
and μ
H2O
to be constant across the diagrams. The value of
μ
Al2O3
varies across the diagram, and is effectively a
passive variable, its value controlled via Gibbs-Duhem
relationships. Allowing μ
Al2O3
to vary freely between
different equilibria without considering diffusion of Al
2
O
3
is equivalent to considering Al
2
O
3
to be immobile.
Figures 8a, b and c show the chemical potential
relationships between K
2
O, FeO and MgO at the low
temperature side of the reaction g+bi=chl, at this condition
the chl-bi-mu assemblage is the stable one (solid lines)
while the dash lines are the metastable assemblage (g-bi-
mu). The stable and the metastable points are connected by
the line for the common assemblage bi-mu. At higher
temperature conditions, the edge of the garnet grain along
with coexisting muscovite and/or biotite would lie on the g-
mu or g-bi lines or, most likely, at the g-bi-mu point
(Fig. 8a, b and c). The interior of the garnet porphyroblast
could lie within the garnet one phase field, or lie at its edge.
a
b
(ii)
(i)
Fig. 7 Cartoon illustrates the focused fluid infiltration that caused the
preferential growth of chlorite in the strain cap region. ais before
chlorite growth and bis during chlorite growth. The shading
corresponds to the concentration of fluid in the bulk rock with light
grey= 9.8 mole% of water, dark grey=18.3 mole% of water in bulk;
medium grey= between 9.8 and 18.3 mole%
Strain caps formation mechanism 15
The garnet-absent assemblage distal from the garnet grains
would contain either biotite-muscovite prior the nucleation
of chlorite or chlorite-biotite-muscovite if chlorite had
nucleated. Figures 8b and c show that the chlorite-bearing
lines and the bi-mu line, lie at higher μ
K2O
than the
corresponding garnet-bearing equilibria. Thus, there would
tend to be diffusion of K
2
O from the strain cap region to the
garnet porphyroblast, once chlorite had nucleated. At the
same time the FeO and the MgO diffuse from the garnet
porphyroblast to the strain cap region (Fig. 8).
Consider an edge of garnet crystal located at contact
with muscovite and biotite crystals, as replacement pro-
ceeds and the chlorite nucleated, the original point of the
g-bi-mu will move along the bi-mu line towards the chl-bi-
mu point. If the garnet grain at the contact with a biotite or
a muscovite grain, the mineral equilibria lay at the line g-bi
or g-mu. Once the chlorite nucleated the equilibria will
move along the lines g-bi or g-mu till g-bi-mu point then
along the bi-mu line toward the stable assemblage chl-bi-
mu. While minimal addition of K
2
O is involved in driving
μ
K2O
up (Fig. 8b and c) to the g-mu or to the g-bi lines if
the newly exposed garnet lay within the garnet one phase
field. Once the garnet grain is consumed, the gradient
across the muscovite and biotite will be flattened out, with
final position being the chl-bi-mu point. In case that the
chlorite cannot be nucleated, the strain cap region will lose
the K
2
O toward the garnet porphyroblast and it will gain
FeO and MgO from the garnet and the equilibria will be at
the stable part of the bi-mu line in Fig. 8. However loss of
K
2
O from the cap region would occur regardless of where
on the bi-mu line the cap region equilibria is located and
this would rapidly drive the assemblage towards the chl-bi-
mu point.
Conclusion
Strain caps are defined as a region that is depleted in quartz
and enriched in micas occurring on opposite sides of the
rigid body, in the quarters orthogonal to the strain shadow.
In examples from the Bhutan Himalaya, the Chugach
metamorphic complex of Alaska and the Taba metamorphic
complex of Sinai chlorite grows exclusively in the strain
cap region around porphyroblasts of garnet, muscovite and
staurolite, respectively. Thermodynamic modeling shows
that chlorite may form due to both, cooling or hydration but
the microstructural position allows to conclude that the
Fig. 8 Calculated μ-μdiagrams in KFMASH calculated at 550°C
and 7 kbar. These P-T conditions are on the lower-Tside of the
reaction g+bi=chl. Each diagram is drawn with quartz and H
2
Oin
excess. In each diagram the stable equilibria are shown as solid lines
and the metastable equilibria are shown as dashed lines
µMgO (KJmol-1)
µ
M
g
O
(
KJmol-1
)
µFeO (KJmol-1)
-678.56
-339.6
-339.59
-339.58
-678.55
bi
bi mu
g
chl
-678.555
KFMASH (+q +H2O) at 550 C, 7 kbar
µFeO (KJmol
-1
)
µK
2
O (KJmol
-1
)
µK
2
O (KJmol
-1
)
KFMASH (+q +H2O) at 550 C, 7 kbar
bi
mu
g
bi
chl
mu
mu
-339.60 -339.59
bi
mu
g
bi
mu
chl
KFMASH (+q +H2O) at 550 C, 7 kbar
-339.595
-920.40
-920.30
-920.32
-920.34
-920.36
-920.38
-678.56 -678.55-678.555
-920.40
-920.30
-920.32
-920.34
-920.36
-920.38
direction of MgO and FeO diffusion
direction of FeOd
iffusion
directiono
f K O diffusion
2
directiono
f MgO diffusion
directionof KOdiffusion
2
16 T. Abu-Alam, K. Stüwe
polyphase strain caps must form due to focused fluid
infiltration along the foliation planes of a pre-existing fabric
postdating to the metamorphic peak. A local gradient in
fluid concentration that includes at least 8.5 mole%
difference in fluid mode of the bulk over a distance
between strain cap and strain shadow must be maintained
in order to avoid chlorite growth elsewhere in the
microstructure. As such, our study presents a quantification
of differences in fluid concentration infiltrating rocks on a
centimeter scale.
Acknowledgements We thank Johann G. Raith for his constructive
suggestions and efficient editorial handling of the manuscript.
Appendix 1
The independent reaction sets between the end-members as
calculated by THERMOCALC. Sections a, b and c are for
Bhutan, Sinai and Alaska, respectively.
Section a:
1) 3east + 6q = py + phl + 2mu
2) phl + east + 6q = py + 2cel
3) 2ann + mu + 6q = alm + 3fcel
4) py + 3east + 3fcel = alm + 3phl + 3mu
3) 2east + 6q = py + mu + cel
4) alm + 6east + 18q = 4py + 3mu + 3fcel
5) ann + 3east + 12q = 2py + mu + 3fcel
6) alm + 3east + 6q = 2py + ann + 2mu
7) 2alm + 3phl + 3east + 18q = 5py + 6fcel
8) ann + east + 6q = alm + 2cel
9) py + ann = alm + phl
10) 3ann + 3east + 18q = 2py + alm + 6fcel
11) 2phl + ames + 6q = py + 2cel + clin
12) alm + 6ann + 3ames + 18q = 4py + 6fcel + 3daph
13) 2py + 7fcel + daph = 4ann + 3east + 19q + 4H
2
O
14) 2east + ames + 6q = py + 2mu + clin
15) 5alm + 2phl + 3ames + 6q = 6py + 2mu + 3daph
16) py + 6ann + 3ames + 18q = 4alm + 6fcel + 3clin
17) 3alm + 12fcel + 2clin = 7ann + 5east + 32q + 8H
2
O
18) 4alm + 3fcel + 3ames = 4py + 3mu + 3daph
19) py + 3fcel = alm + 3cel
20) 6ann + 3ames + 18q = 2py + alm + 6cel + 3daph
Section b:
1) 18phl + 2mst + 13mu = 31east + 46q + 4H
2
O
2) 2phl + mu + 2pa = 3east + 2ab + 3q + 2H
2
O
3) 62phl + 6fst + 39mu = 8ann + 93east + 138q +
12H
2
O
4) 30phl + 8ann + 7mu + 92pa = 45east + 92ab + 6fst +
80H
2
O
5) 16phl + 8ann + 78pa + 21q = 24east + 78ab + 6fst +
66H
2
O
6) 8ann + 62pa + 45q = 62ab + 6fst + 8mu + 50H
2
O
7) 41east + 13clin + 47q = 41phl + 6mst + 40H
2
O
8) phl + 3mu + clin = 4east + 7q + 4H
2
O
9) 41mu + 18clin = 41east + 2mst + 80q + 68H
2
O
10) 41mu + 31clin = 41phl + 8mst + 33q + 108H
2
O
11) 33east + 47mu + 46clin = 80phl + 14mst + 156H
2
O
12) 5phl + 2mst + 13cel = 18east + 46q + 4H
2
O
13) 4mst + 41cel + 5clin = 41east + 127q + 28H
2
O
14) 3cel + clin = 2phl + east + 7q + 4H
2
O
15) 41cel + 18clin = 41phl + 2mst + 80q + 68H
2
O
16) 80east + 47cel + 46clin = 127phl + 14mst + 156H
2
O
17) 127mu + 46clin = 47east + 14mst + 80cel + 156H
2
O
18) 41mu + 13clin + 47q = 6mst + 41cel + 40H
2
O
19) 2mst + 18cel = 13east + 5mu + 46q + 4H
2
O
20) 2mu + cel + clin = 3east + 7q + 4H
2
O
21) 41ames + 47q = 6mst + 28clin + 40H
2
O
22) 3mu + 5clin = 3phl + 4ames + 7q + 4H
2
O
23) 47mu + 13clin + 33ames = 47phl + 14mst + 156H
2
O
24) 28mu + 31ames + 13q = 28phl + 10mst + 104H
2
O
25) 80mu + 46daph = 47ann + 14fst + 33fcel + 156H
2
O
26) 41mu + 13daph + 47q = 6fst + 41fcel + 40H
2
O
27) 2fst + 31fcel = 13ann + 18mu + 46q + 4H
2
O
28) 4fcel + daph = 3ann + mu + 7q + 4H
2
O
29) 41fcel + 18daph = 41ann + 2fst + 80q + 68H
2
O
30) 41mu + 31daph = 41ann + 8fst + 33q + 108H
2
O
31) 6fst + 93cel = 31phl + 8ann + 54mu + 138q + 12H
2
O
32) 3cel + 2pa = phl + 2ab + 2mu + 3q + 2H
2
O
33) 8ann + 45cel + 92pa = 15phl + 92ab + 6fst + 38mu +
80H
2
O
34) 4phl + 8ann + 54pa + 57q = 54ab + 6fst + 12cel +
42H
2
O
35) 3east + 3fcel = 2phl + ann + 3mu
36) 12phl + 2fst + 13fcel = 7ann + 18east + 46q + 4H
2
O
37) 26phl + 2fst + 21mu = 39east + 8fcel + 46q + 4H
2
O
38) phl + 3fcel = ann + 3cel
39) 3fcel + 2pa = ann + 2ab + 2mu + 3q + 2H
2
O
40) mst + 4fcel = fst + 4cel
41) 13mst + 36fcel = 26east + 9fst + 10mu + 92q + 8H
2
O
42) 2fst + 26cel = 13east + 5mu + 8fcel + 46q + 4H
2
O
43) east + cel = phl + mu
44) 6fst + 62cel = 8ann + 31east + 23mu + 138q + 12H
2
O
45) 23phl + 6fst + 39cel = 8ann + 54east + 138q + 12H
2
O
Section c:
1) 3cel + ann = 3fcel + phl
2) 3fcel + 2pa = 2ab + 2mu + ann + 3q + 2H
2
O
3) 4fcel + daph = mu + 3ann + 7q + 4H
2
O
4) 12cel + 3daph = 3mu + 4phl + 5ann + 21q + 12H
2
O
5) 9cel + daph = mu + 5fcel + 3phl + 7q + 4H
2
O
6) 20fcel + 9clin = 5mu + 15phl + 4daph + 35q + 20H
2
O
7) 4cel + clin = mu + 3phl + 7q + 4H
2
O
8) 5cel + daph = 5fcel + clin
9) 5phl + 3daph = 5ann + 3clin
Strain caps formation mechanism 17
10) 9mu + 8phl + 3daph = 5ann + 12east + 21q + 12H
2
O
11) 3mu + phl + clin = 4east + 7q + 4H
2
O
12) 15mu + 5ann + 8clin = 20east + 3daph + 35q +
20H
2
O
13) 2ab + 5fcel + 2daph = 2pa + 5ann + 11q + 6H
2
O
14) 6ab + 15cel + 6daph = 6pa + 5phl + 10ann + 33q +
18H
2
O
15) 2ab + 15cel + 2daph = 10fcel + 2pa + 5phl + 11q +
6H
2
O
16) 2ab + 5fcel + 3clin = 2pa + 5phl + daph + 11q +
6H
2
O
17) 2ab + 5cel + 2clin = 2pa + 5phl + 11q + 6H
2
O
18) 9fcel + clin = mu + 5cel + 3ann + 7q + 4H
2
O
19) 12fcel + 3clin = 3mu + 5phl + 4ann + 21q + 12H
2
O
20) 20cel + 9daph = 5mu + 15ann + 4clin + 35q + 20H
2
O
21) 3fcel + 3east = 3mu + 2phl + ann
22) 3fcel + 6pa + 4phl = 6ab + ann + 6east + 9q + 6H
2
O
23) mu + 2pa + 2phl = 2ab + 3east + 3q + 2H
2
O
24) 8ab + 5mu + 3clin = 8pa + 5phl + 9q + 4H
2
O
25) 9cel + 14pa + 2phl = 14ab + 11mu + 3clin + 2H
2
O
26) 5cel + 6pa = 6ab + 5mu + clin + 2q + 2H
2
O
27) 3cel + 2pa = 2ab + 2mu + phl + 3q + 2H
2
O
28) 8ab + 5mu + 3daph = 8pa + 5ann + 9q + 4H
2
O
29) 9fcel + 14pa + 2ann = 14ab + 11mu + 3daph + 2H
2
O
30) 5fcel + 6pa = 6ab + 5mu + daph + 2q + 2H
2
O
31) 45cel + 70pa + 10ann = 70ab + 55mu + 9clin + 6daph +
10H
2
O
32) 2ab + 5cel + 3daph = 2pa + 5ann + clin + 11q + 6H
2
O
33) 2ab + 15fcel + 2clin = 10cel + 2pa + 5ann + 11q +
6H
2
O
34) 6ab + 15fcel + 6clin = 6pa + 10phl + 5ann + 33q +
18H
2
O
35) 5ann + 3ames = 2phl + 3east + 3daph
36) phl + ames = east + clin
References
Abu-Alam TS, Stüwe K (2009a) Exhumation during oblique trans-
pression: an example from the Feiran-Solaf region, Egypt. J
Metamorph Geol 27:439–459
Abu-Alam TS, Stüwe K (2009b) MBC (1.7): a visual basic code to
calculate bulk composition of rocks from chemistry and volume
proportions of the phases. Geol Soc America (GSA)-Abstract
with programs 41, pp 500
Abu-Alam TS, Stüwe K (2009c) The MBC (1.7): a visual basic
program to calculate bulk composition of rocks from chemistry
and the volume proportions of the phases. Mitt Österr Mineral
Ges 155, p 17
Abu-Alam TS, Stüwe K, Hauzenberger C (2010) Calc-silicates from
Wadi Solaf region, Sinai, Egypt. J Afr Earth Sci 58:475–488
Abu El-Enen MM (1995) Geological, geochemical and mineralogical
studies on the metamorphic rocks between Wadi Um-Maghra and
Wadi Tweiba area, southeastern Sinai, Egypt. Ph.D. thesis,
Mansoura Univ, Egypt. 172 pp
Abu El-Enen MM, Zalata AA, El-Metwally AA, Okrusch M (1999)
Orthogneisses from the Taba metamorphic belt, SE Sinai, Egypt:
witnesses for granitoid magmatism at an active continental
margin. N Jahrb Mineral Abh 175:53–81
Abu El-Enen MM, Will TM, Okrusch M (2004) P–T evolution of the
Pan-African Taba metamorphic belt, Sinai, Egypt: constraints
from metapelitic mineral assemblages. J Afr Earth Sci 38:59–78
Ague JJ (1991) Evidence for major mass transfer and volume strain
during regional metamorphism of pelites. Geology 19:855–858
Barker F, Farmer GL, Ayuso RA, Plafker G, Lull JS (1992) The
50 Ma granodiorite of the eastern Gulf of Alaska: Melting in an
accretionary prism in the forearc. J Geophys Res 97:6757–6778
Bruand E, Gasser D, Stüwe K, Beyssac O (2010) Metamorphism of
the Chugach Metamorphic Complex, (Alaska). New pressure
estimates question the ridge subduction context. Geophys Res
Abst 12:EGU2010–12235
Coggon R, Holland TJB (2002) Mixing properties of phengitic micas
and revised garnet-phengite thermobarometers. J Metamorph
Geol 20:683–696
Cosca MA, Shimron A, Caby R (1999) Late Precambrian metamor-
phism and cooling in the Arabian-Nubian Shield: petrology and
40
Ar/
39
Ar geochronology of metamorphic rocks of the Elat area
(southern Israel). Precambrian Res 98:107–127
Dasgupta S, Ganguly J, Neogi S (2004) Inverted metamorphic
sequence in the Sikkim Himalayas: crystallization history, P–T
gradient and implications. J Metamorph Geol 22:395–412
Eliwa HA, Abu El-Enen MM, Khalaf IM, Itaya T, Murata M (2008)
Metamorphic evolution of Neoproterozoic metapelites and
gneisses in Sinai, Egypt: insights from petrology, mineral
chemistry and K–Ar age dating. J Afr Earth Sci 51:107–122
England P, Le Fort P, Molnar P, Pecher A (1992) Heat sources for
Tertiary metamorphism and anatexis in the Annapurna-Manaslu
region, Central Nepal. J Geophys Res 97:2107–2128
Etchecopar A, Malavieille J (1987) Computer models of pressure
shadows: a method for strain measurement and shear-sense
determination. J Struct Geol 9:667–677
Farris DW, Paterson SR (2007) Physical contamination of silicic
magmas and fractal fragmentation of xenoliths in Paleocene
plutons on Kodiak Island, AK. Can Mineral 45:107–129
Farris DW, Paterson SR (2009) Subduction of a segmented ridge
along a curved continental margin: variations between the
western and eastern Sanak-Baranof belt, southern Alaska.
Tectonophysics 464:100–117
France-Lanord C, Le Fort P (1988) Crustal melting and granite
genesis during the Himalayan collision orogenesis. Trans R Geol
Soc Edinburgh 79:183–195
Gansser A (1964) Geology of the Himalayas. Wiley Interscience,
London
Grujic D, Casey M, Davidson C, Hollister L, Kündig R, Pavlis T
(1996) Ductile extrusion of the higher Himalayan Crystalline in
Bhutan: evidence from quartz microfabrics. Tectonophysics
260:21–44
Guiraud M, Powell R, Rebay G (2001) H
2
O in metamorphism and
unexpected behaviour in the preservation of metamorphic
mineral assemblages. J Metamorph Geol 19:445–454
Heimann A, Eyal Y, Eyal M, Foland KA (1995) Thermal events and
low temperature alteration in the Precambrian schistose dykes
and their host rocks in the Elat area, southern Israel:
40
Ar/
39
Ar
geochronology. In: Baer G, Heimann A (eds) Physics and
Chemistry of Dykes. Balkema, Rotterdam, pp 281–292
Hill M, Morris J, Whelan J (1981) Hybrid granodiorites intruding the
accretionary prism, Kodiak, Shumagin and Sanak Islands,
Southwest Alaska. J Geophys Res 86:10569–10590
Hodges KV, Bowring S, Davidek K, Hawkins D, Krol M (1988)
Evidence for rapid displacement on Himalayan normal faults and
18 T. Abu-Alam, K. Stüwe
the importance of tectonic denudation in the evolution of
mountain ranges. Geology 26:483–486
Hodges KV, Parrish RR, Searle MP (1996) Tectonic evolution of the
central Annapurna Range, Nepalese Himalayas. Tectonics
15:1264–1291
Holland TJB, Powell R (1998) An internally consistent thermody-
namic dataset for phases of petrological interest. J Meta Geol
16:309–343
Holland TJB, Baker JM, Powell R (1998) Mixing properties and
activity-composition relationships of chlorites in the system
MgO-FeO-Al
2
O
3
-SiO
2
-H
2
O. Eur J Mineral 10:395–406
Hubbard MS (1989) Thermobarometric constraints on the thermal
history of the Main Central thrust zone and Tibetan slab, eastern
Nepal Himalaya. J Metamorph Geol 7:19–30
Hubbard MS, Harrison TM (1989)
40
Ar/
39
Ar constraints on deforma-
tion and metamorphism in the Main Central Thrust zone and
Tibetan Slab, eastern Nepal Himalaya. Tectonics 8:865–880
Hudson T, Plafker G (1982) Paleogene metamorphism of an
accretionary flysch terrane, eastern Gulf of Alaska. Geol Soc
Amer Bull 93:1280–1290
Keller LM, Abart R, Schmid SM, De Capitani C (2005) Phase
relations and chemical composition of phengite and paragonite in
pelitic schists during decompression: a case study from the
Monte Rosa Nappe and Camughera–Moncucco Unit, Western
Alps. J Petrol 46:2145–2166
Kröner A, Eyal M, Eyal Y (1990) Early Pan-African evolution of the
basement around Elat, Israel, and Sinai Peninsula revealed by
single-zircon evaporation dating, and implication for crustal
accretion rates. Geology 18:545–548
Le Fort P (1975) Himalaya: the collided range. Present knowledge
about the continental arc. Am J Sci 275A:1–44
Le Fort P (1981) Manaslu leucogranite: a collision signature of the
Himalaya. A model for its genesis and emplacement. J Geophys
Res 86:10545–10568
Mahar EM, Baker JM, Powell R, Holland TJB, Howell N (1997) The
effect of Mn on mineral stability in metapelites. J Metamorph
Geol 15:223–238
Moore JC, Byne T, Plumley PW, Reid M, Gibbons H, Coe RS (1983)
Paleogene evolution of the Kodiak Islands, Alaska: consequences
of ridge-trench interaction in a more southerly latitude. Tectonics
2:265–293
Passchier CW, Trouw RAJ (1996) Micro-tectonics. Springer Verlag,
New York, 289
Passchier CW, Trouw RAJ (2005) Micro-tectonics. Springer Verlag,
New York, 366
Plafker G, Moore JC, Winkler GR (1994) Geology of the southern
Alaska margin: the Geology of North America, G-1. The
Geology of Alaska. Geol Soc Am 12:389–449
Powell R, Holland TJB (1988) An internally consistent thermody-
namic dataset with uncertainties and correlations: 3. Application,
methods, work examples and a computer program. J Metamorph
Geol 6:173–204
Shaw DM (1956) Geochemistry of pelitic rocks. Part III: Major
elements and general geochemistry. Bull Geol Soc Am 67:919–
934
Shimron AE (1980) Proterozoic island arc volcanism and sedimenta-
tion in Sinai. Precambrian Res 12:437–458
Sisson VB, Hollister LS (1988) Low-pressure facies series metamor-
phism in an accretionary sedimentary prism, southern Alaska.
Geology 16:358–361
Sisson VB, Hollister LS, Onstott TC (1989) Petrologic and age
constraints on the origin of a low-pressure/high-temperature
metamorphic complex, southern Alaska. J Geophys Res
94:4392–4410
Sisson VB, Poole PR, Harris NR, Burner HC, Pavlis TL, Copeland P,
Donelick RA, McLelland WC (2003) Geochemical and geochro-
nologic constraints for the genesis of tonalite-trondhjemite suite
and associated mafic intrusive rocks in the eastern Chugach
Mountains, Alaska: a record of ridge-transform subduction:
geology of a transpressional orogen developed during ridge-
trench interaction along the North Pacific Margin. Geol Soc Am
Spec Pap 371:293–326
Stüwe K, Foster D (2001)
40
Ar/
39
Ar, pressure, temperature and fission
track constraints on the age and nature of metamorphism around
the main central thrust in the eastern Bhutan Himalaya. J Asia
Earth Sci 19:85–95
Takagi H, Ito M (1988) The use of asymmetric pressure shadows in
mylonites to determine the sense of shear. J Struct Geol 10:347–
360
Tenczer V, Stüwe K, Barr TD (2001) Pressure anomalies around
cylindrical objects in simple shear. J Struct Geol 23:777–788
Tenczer V, Powell R, Stüwe K (2006) Evolution of H
2
O content in a
polymetamorphic terrane: the Plattengneiss Shear Zone (Koralpe,
Austria). J Metamorph Geol 24:281–295
Trouw RAJ, Tavares FM, Robyr M (2008) Rotated garnets: a
mechanism to explain the high frequency of inclusion trail
curvature angles around 90° and 180°. J Struct Geol 30:1024–
1033
Vannay JC, Hodges KV (1996) Tectonometamorphic evolution of the
Himalayan metamorphic core between Annapurna and Dhaula-
giri, central Nepal. J Metamorph Geol 14:635–656
Vidal P, Cocherie A, Le Fort P (1982) Geochemical investigations of
the origin of the Manaslu leucogranite (Himalaya, Nepal).
Geochim Cosmochim Acta 46:2279–2292
White RW, Powell R, Holland TJB (2007) Progress relating to
calculation of partial melting equilibria for metapelites. J
Metamorph Geol 25:511–527
White RW, Powell R, Baldwin JA (2008) Calculated phase equilibria
involving chemical potentials to investigate the textural evolution
of metamorphic rocks. J Metamorph Geol 26:181–198
Strain caps formation mechanism 19
