JOURNAL OF PETROLOGY
VOLUME 41
NUMBER 12
PAGES 1673–1719
2000
Contrasting P–T Paths in the Eastern
Himalaya, Nepal: Inverted Isograds in a
Paired Metamorphic Mountain Belt
BEN GOSCOMBE∗ AND MARTIN HAND
GROUP FOR INTEGRATED TECTONIC STUDIES AT DEPARTMENT OF GEOLOGY AND GEOPHYSICS,
ADELAIDE UNIVERSITY, ADELAIDE, S.A. 5005, AUSTRALIA
RECEIVED NOVEMBER 12, 1998; REVISED TYPESCRIPT ACCEPTED APRIL 17, 2000
Petrology and phase equilibria of rocks from two profiles in Eastern
Nepal from the Lesser Himalayan Sequences, across the Main
Central Thrust Zone and into the Greater Himalayan Sequences
reveal a Paired Metamorphic Mountain Belt (PMMB) composed
of two thrust-bound metamorphic terranes of contrasting metamorphic
style. At the higher structural level, the Greater Himalayan Sequences
experienced high-T/moderate-P metamorphism, with an anticlockwise P–T path. Low-P inclusion assemblages of quartz +
hercynitic spinel + sillimanite have been overgrown by peak
metamorphic garnet + cordierite + sillimanite assemblages that
equilibrated at 837 ± 59°C and 6·7 ± 1·0 kbar. Matrix
minerals are overprinted by numerous metamorphic reaction textures
that document isobaric cooling and re-equilibrated samples preserve
evidence of cooling to 600 ± 45°C at 5·7 ±1·1 kbar. Below
the Main Central Thrust, the Lesser Himalayan Sequences are a
continuous (though inverted) Barrovian sequence of high-P/moderate-T metamorphic rocks. Metamorphic zones upwards from the
lowest structural levels in the south are:
Zone A: albite + chlorite + muscovite ± biotite;
Zone B: albite + chlorite + muscovite + biotite + garnet;
Zone C: albite + muscovite + biotite + garnet ± chlorite;
Zone D: oligoclase + muscovite + biotite + garnet ± kyanite;
Zone E: oligoclase + muscovite + biotite + garnet +
staurolite + kyanite;
Zone F: bytownite + biotite + garnet + K-feldspar +
kyanite ± muscovite;
Zone G: bytownite + biotite + garnet + K-feldspar +
sillimanite + melt ± kyanite.
The Lesser Himalayan Sequences show evidence for a clockwise
P–T path. Peak-P conditions from mineral cores average 10·0 ±
1·2 kbar and 557 ± 39°C, and peak-metamorphic conditions
from rims average 8·8 ± 1·1 kbar and 609 ± 42°C in Zones
∗Corresponding author. Present address: 18 Cambridge Road, Aldgate,
Adelaide, S.A. 5154, Australia. E-mail: [email protected]
Extended data set can be found at:
http://www.petrology.oupjournals.org
D–F. Matrix assemblages are overprinted by decompression reaction
textures, and in Zones F and G progress into the sillimanite field.
The two terranes were brought into juxtaposition during formation
of sillimanite–biotite ± gedrite foliation seams (S3) formed at
conditions of 674 ± 33°C and 5·7 ± 1·1 kbar. The contrasting
average geothermal gradients and P–T paths of these two metamorphic terranes suggest they make up a PMMB. The upperplate position of the Greater Himalayan Sequences produced an
anticlockwise P–T path, with the high average geothermal gradient
being possibly due to high radiogenic element content in this terrane.
In contrast, the lower-plate Lesser Himalayan Sequences were
deeply buried, metamorphosed in a clockwise P–T path and display
inverted isograds as a result of progressive ductile overthrusting of the
hot Greater Himalayan Sequences during prograde metamorphism.
thermobarometry; P–T paths; Himalaya; metamorphism;
inverted isograds; paired metamorphic belts
KEY WORDS:
INTRODUCTION
Previous work in the Himalayas interpreted all metamorphic mineral assemblages to have formed in the
Himalayan Metamorphic Cycle during collisional orogenesis in the Tertiary (Pecher, 1989; Vannay & Hodges,
1996). Although one of the youngest metamorphic belts,
use of the Himalaya as a type example of tectonic and
metamorphic processes in collisional orogens is limited
because the tectonometamorphic evolution is complex
and not entirely understood. This study of the Eastern
 Oxford University Press 2000
JOURNAL OF PETROLOGY
VOLUME 41
Himalaya recognizes two thrust-bound units of
contrasting metamorphic style (termed metamorphic
‘terranes’): (1) the Lesser Himalayan Sequences (LHS),
which underwent Barrovian high-P/moderate-T metamorphism and display a clockwise P–T path; (2) at
higher structural levels, the high-T/moderate-P Greater
Himalayan Sequences (GHS), which display an anticlockwise P–T path. Matrix mineral assemblages in both
terranes were formed during SSW- to S-directed ductile
overthrusting during collisional orogenesis. Thus, the two
terranes constitute a Paired Metamorphic Mountain Belt
(PMMB), in a continent–continent collisional orogen
(Armstrong et al., 1992), distinct from subduction-type
paired metamorphic belts in the sense of Miyashiro
(1973). Anticlockwise P–T paths and high average geothermal gradients have not been reported from elsewhere
in the Himalayan orogen and may be restricted to a
portion of the Eastern Himalaya only.
East of the Mount Everest–Dudh Kosi area in Eastern
Nepal (Fig. 1), very few metamorphic studies have been
previously undertaken. All are from the Makalu–Arun
River area (Brunel & Kienast, 1986: Lombardo et al.,
1993; Meier & Hiltner, 1993; Pognante & Benna, 1993),
with none available for the Kangchenjunga region. These
previous metamorphic studies document metamorphic
assemblages in the different structural units and attribute
these to crystallization events on the one clockwise P–T
path. In an attempt to characterize the metamorphism
of the wider Eastern Nepal region, both with petrological
analysis and P–T calculations, two profiles across the
LHS and GHS were mapped and sampled by the author
in 1992 and 1997 (Fig. 1). These are the Makalu profile
along the Arun River and upper Barun Khola (Fig. 2)
and the Kangchenjunga profile along the Tamur River
and Kabeli Khola (Fig. 3). This study reports new petrological data and a comparable set of P–T estimates
using one thermobarometric method [THERMOCALC;
Powell & Holland (1988)], these providing a framework
for the tectonometamorphic evolution of the Eastern
Nepal region. The distribution of metamorphic assemblages does not differ significantly from that given
by previous workers. However, in contrast, the present
study reports significant differences in interpretative P–T
paths and peak metamorphic conditions for both the
LHS (P >3 kbar higher) and GHS (T >50–100°C
higher). Variation in metamorphic conditions through
the Eastern Nepal profiles is shown to be discontinuous,
and is interpreted as reflecting two metamorphic terranes
with contrasting P–T evolution, although formed within
the same metamorphic cycle.
REGIONAL GEOLOGY
The entire Himalayan front is a metamorphic belt that
was metamorphosed and exhumed in the Tertiary
NUMBER 12
DECEMBER 2000
Himalayan Orogeny as a result of collision of the Indian
subcontinent with Asia. The Himalayan metamorphic
belt has been subjected to only one metamorphic cycle.
Along the entire length of the Himalayan front, metamorphic field gradients are inverted, with metamorphic
grade increasing towards higher structural levels (e.g.
LeFort, 1975; Hodges et al., 1988; Pecher, 1989; Inger
& Harris, 1992; Vannay & Hodges, 1996). Isograds are
parallel to the main metamorphic foliation and vary
smoothly across the Main Central Thrust (MCT) Zone
(Table 1), without a break in grade (Hubbard, 1989;
Metcalfe, 1993; Hubbard, 1996). Thus metamorphic
equilibration is interpreted to have been synchronous
with ductile, southward thrusting along the MCT (Pecher,
1989).
Numerous models for the origin of the inverted isograds
in the Himalayas have been discussed in detail by Mohan
et al. (1989), Searle & Rex (1989) and Hubbard (1996).
Some workers (Hodges et al., 1988; Pecher, 1989; Inger
& Harris, 1992; Vannay & Hodges, 1996) proposed that
the Himalayan Metamorphic Cycle is polyphase, with
two metamorphic events (Table 2). Those workers saw
evidence for an Eohimalayan metamorphism spanning
37–24 Ma age (Inger & Harris, 1992; Hodges et al., 1994;
Parrish & Hodges, 1996; Coleman & Hodges, 1998).
This high-P/moderate-T Barrovian event is thought to
be associated with burial and prograde metamorphism
(Inger & Harris, 1992; Vannay & Hodges, 1996). Neohimalayan metamorphism at 22–13 Ma (Hubbard &
Harrison, 1989; Inger & Harris, 1992; Metcalfe, 1993)
is interpreted as a moderate-P/moderate-T peak metamorphic event (Hodges et al., 1988) associated with upthrusting of the GHS (Table 1) along the MCT (Inger
& Harris, 1992; Hodges et al., 1996; Vannay & Hodges,
1996).
The Himalayan metamorphic belt is composed of
laterally extensive sequences bounded by crustal-scale
north-dipping thrusts. The basal unit, the LHS (Table
1), is interpreted as Middle Proterozoic in age (Parrish
& Hodges, 1996) or alternatively as Upper Proterozoic to
Palaeozoic Tethyan meta-sediments (Vannay & Hodges,
1996). The Main Boundary Thrust forms the basal
contact, with the LHS thrust onto the Late Miocene
molasse of the Siwalik Group (LeFort, 1975). The LHS
is overthrust by the GHS at the MCT. The GHS is
continuously exposed along the Himalayan front and
interpreted as a metamorphosed sequence of Late Proterozoic supracrustal rocks and orthogneisses (Bhanot et
al., 1977: Miller & Frank, 1992; Parrish & Hodges, 1996),
along with Late Cambrian granitoid rocks (LeFort et al.,
1986). Although the GHS is thought to have originally
been overthrust by the Tethyan Sequences of Upper
Proterozoic to Eocene age, the two terranes are separated
by a series of extensional detachment faults, the Southern
Tibet Detachment System (STDS) (Table 1; Fig. 1) (Burg
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GOSCOMBE AND HAND
P–T PATHS IN EASTERN HIMALAYA
Fig. 1. Simplified geology of the Eastern Himalaya in Nepal, with location of Makalu (Fig. 2) and Kangchenjunga profiles (Fig. 3) indicated.
(a) Tibetan series sediments. Proterozoic to early Palaeozoic Greater Himalayan Sequences (GHS) composed of (c) Barun Gneiss and
Jannu–Kangchenjunga Gneiss, (d) Namche Migmatite Orthogneiss, (e) Black Gneisses and (f ) Miocene leucogranite. Uppermost Lesser Himalayan
Sequences (LHS) include (b) Himal Group and (g) Kathmandu Group with (h) Num Orthogneiss Formation and (k) Cambrian granites.
Palaeozoic Tethyan rocks of the LHS in the Midlands Group are (i) Kushma Formation, ( j) Ulleri Formation and (l) Seti Formation. Below the
Main Boundary Thrust (MBT) is (m) Molasse of the Siwalik Formation. STDS, Southern Tibet Detachment System; ITSZ, Intra-Tibetan
Suture Zone. Other major thrusts are labelled MCTI and MCTII after Maruo & Kizaki (1981) (MCT, Main Central Thrust). Geology sourced
from Bordet (1961), Shrestha et al. (1984, 1985), Lombardo et al. (1993), Morrison & Oliver (1993) and authors. Regional setting inset after
Harris et al. (1993). E, Mount Everest; M, Makalu; K, Kangchenjunga.
et al., 1984; Burchfiel et al., 1992; Edwards & Harrison,
1997).
The sequence of thrusting and reactivation of crustalscale shear zones during collision is complex. All Tethyan
oceanic crust had been subducted and the Indian subcontinent had collided with Asia by Middle Eocene times
(Sengor, 1990). Prograde burial of the GHS and LHS
occurred during the Middle Eocene to Oligocene by
overthrusting of Tethyan Sequences at the Eohimalayan
Thrust, the precursor to the STDS (Vannay & Hodges,
1996). Peak metamorphic conditions in the Early Miocene resulted in emplacement of leucogranites of
26–17 Ma age into the highest structural levels of the
GHS (Scharer, 1984; Deniel et al., 1987; Copeland et al.,
1990; Hodges et al., 1996, 1998; Edwards & Harrison,
1997; Coleman, 1998). The MCT was initiated at
>22 Ma (Harrison et al., 1995; Hodges et al., 1996,
1998; Edwards & Harrison, 1997; Coleman, 1998) and
remained active throughout metamorphism, resulting in
the inverted metamorphic sequence (Hubbard & Harrison, 1989; Inger & Harris, 1992). Metamorphism immediately below the MCT may be as young as 6–13 Ma
(Macfarlane et al., 1992; Harrison et al., 1997). Isostatic
uplift and extensional collapse of the High Himalayas
was accommodated by episodic movement on extensional
detachments at structural levels above the MCT, at the
STDS, from >22 Ma (Burchfiel et al., 1992; Hodges et
al., 1996, 1998; Vannay & Hodges, 1996; Edwards &
Harrison, 1997; Harrison et al., 1999).
In the region investigated, the High Himalayas of
Eastern Nepal, the GHS is composed of a variety of
distinct gneissic rock units. In the Makalu profile (Fig.
2); the base of the GHS is marked by the principal thrust
of the MCT zone, locally called the Barun Thrust. In
the Kangchenjunga profile the MCT is represented by
the highly sheared informal ‘Biotite Gneiss’ unit. The
overthrust Barun Gneiss (Bordet, 1961; Lombardo et
al., 1993) is composed of high-grade quartzo-feldspathic
gneisses, mafic gneiss and calc-silicates. At higher structural levels is the Namche Migmatite Orthogneiss and
above that the Black Gneisses (Lombardo et al., 1993).
In the Kangchenjunga profile (Fig. 3); the entire GHS
is called the Jannu–Kangchenjunga Gneiss (Mohan
et al., 1989). The Jannu–Kangchenjunga Gneiss is
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NUMBER 12
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Fig. 2. Simplified geological map of Makalu profile and location of samples (prefixed by M in tables). Geology based on Bordet (1961),
Lombardo et al. (1993), Pognante & Benna (1993) and authors.
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GOSCOMBE AND HAND
P–T PATHS IN EASTERN HIMALAYA
Fig. 3. Simplified geological map of Kangchenjunga profile and location of samples (prefixed by K in tables). Geology based on Shrestha et al.
(1984) and authors.
Table 1: Abbreviations
PMMB
Paired Metamorphic Mountain Belt; paired metamorphic belt in a continent–continent collisional orogen environment,
distinct from subduction-type paired metamorphic belts of Miyashiro (1973)
STDS
South Tibet Detachment System
MCT
Main Central Thrust; ductile thrust zone within the upper LHS; principal movement plane at base of GHS (the Barun Trust
and Biotite Gneiss unit in the eastern Nepal region investigated)
MBT
Main Boundary Thrust; basal thrust to Himalayan sequences
GHS
Greater Himalayan Sequences, also called Higher Himalayan Crystallines; Proterozoic sequences from the north Indian
passive margin, including Barun Gneiss, Black Gneisses, Namche Migmatite Orthogneiss and Jannu–Kangchenjunga
Gneiss units in the eastern Nepal region
LHS
Lesser Himalayan Sequences; Palaeozoic sequences mostly of Tethyan origins, including Himal, Kathmandu and
Midlands Groups in the eastern Nepal region
petrologically similar to the Barun Gneiss of the Makalu
profile and both are correlated with each other. At high
structural levels, both the Black Gneisses and Jannu–
Kangchenjunga Gneiss contain leucogranite intrusions
of 22–24 Ma age (Scharer, 1984). Below the Barun Thrust
are the Barrovian metamorphic rocks of the LHS composed of numerous groups and formations (Bordet, 1961;
Lombardo et al., 1993) (Figs 1 and 2). The uppermost
are the undifferentiated Himal Group (Shrestha et al.,
1985) and below that the Num Orthogneiss and Kathmandu Group (Shrestha et al., 1985; Lombardo et al.,
1993) (Figs 1–3). Lowest structural levels of the LHS are
exposed in antiformal tectonic windows (Figs 1–3). These
are thrust-bound units of Tethyan meta-sediments of the
Midland Group, including the Kushma, Ulleri and Seti
Formations (Shrestha et al., 1985).
STRUCTURAL GEOLOGY OF THE
MAKALU AND KANGCHENJUNGA
PROFILES
The lowest structural level of the LHS is a series of thrust
slices of Seti, Ulleri and Kushma Formations of the
Midland Group (Shrestha et al., 1985), exposed in an
antiformal dome in the Kangchenjunga profile (Fig. 3).
The Ulleri Formation is particularly highly sheared and
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JOURNAL OF PETROLOGY
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almost entirely mylonitic. In the Makalu profile, the
principal thrust zone of the MCT and upper bounding
margin of the LHS is a series of thrust sheets within the
Himal Group (Bordet, 1961; Lombardo et al., 1993) (Fig.
2). In the Kangchenjunga profile the principal thrust
zone of the MCT is represented by the 900 m thick,
informal Biotite Gneiss unit (Fig. 3). Biotite Gneiss is a
highly sheared unit of biotite-rich schists and mylonitic
NUMBER 12
DECEMBER 2000
gneisses with numerous pegmatitic and granitic veins and
tight to isoclinal folds. The movement history within the
Biotite Gneiss unit is unknown.
Throughout the entire LHS a strong proto-mylonitic
and schistose metamorphic foliation (S1) is developed
parallel to layering and thrusts, dipping shallowly to the
NE to NW (Fig. 4). L1 stretching lineation is defined
by quartz- and feldspar-aggregate ribbons and aligned
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GOSCOMBE AND HAND
P–T PATHS IN EASTERN HIMALAYA
kyanite, biotite and hornblende grains. L1 is shallowly
(20–30°) north-plunging throughout most of the LHS
and in the Kangchenjunga profile is NE-plunging in the
Himal Group (Fig. 4). S–C relationships, shearbands and
-type porphyroclasts throughout indicate north over
south transport, consistent with shearing in the Himalayan Metamorphic Cycle (Table 2). Aligned micas,
amphiboles and kyanite within the matrix assemblage
suggest that ductile shearing immediately preceded or
accompanied the peak of metamorphism. S1 foliation is
crenulated, with new biotite growth, and is isoclinally to
closely folded around sub-horizontal east–west-trending
axes, normal to L1. F2 folds and crenulations verge to the
south and formed in response to the same tectonic
transport as the S1–L1 fabric. Leucocratic partial melt
segregations occur only at the highest structural levels in
the Himal Group, as thin stromatic layers and elongate
moats parallel to L1 and enclosing garnet or tourmaline
porphyroblasts.
Above the MCT, in contrast to the LHS, the GHS is
a less schistose sequence of high-grade migmatitic gneissic
rock-types. In the Makalu profile the basal GHS unit is
the Barun Gneiss, composed of metapelitic, quartzofeldspathic and mafic gneisses with subordinate calcsilicates and carbonates. The overlying Namche Migmatite Orthogneiss (Lombardo et al., 1993) is structurally
and petrologically similar to the Barun Gneiss and also
contains subordinate carbonate, calc-silicate and mafic
layers; differing only in a greater proportion of stromatic
partial melt segregations (Pognante & Benna, 1993). The
overlying Black Gneisses are quartzo-feldspathic gneisses
that are more schistose than the remainder of the GHS
(Lombardo et al., 1993), but are otherwise structurally
similar. The nature of the contact between the Namche
Migmatite Orthogneiss and the Black Gneisses is unknown. In the Kangchenjunga profile, the Jannu–
Kangchenjunga Gneiss (Mohan et al., 1989) unit is
petrologically and lithologically similar to and occupies
the same structural level as the Barun Gneiss (Fig. 1),
with which it is tentatively correlated. In all units of the
GHS, the majority of partial melt segregations (typically
garnet bearing) pre-date the metamorphic foliation (S1)
and less common biotite pegmatitic veins post-date S1.
Black Gneisses and the highest structural levels of the
Jannu–Kangchenjunga Gneiss are intruded by numerous
leucogranite plutons and sills.
In the Makalu profile, gneisses of the Barun Gneiss
have a coarse polygonal granoblastic texture with a
strong S1–L1 fabric defined by aligned micas, sillimanite,
hornblende and quartz- and feldspar-aggregate ribbons.
This high-grade shear fabric constitutes the coarse matrix
assemblage that was annealed at the peak of metamorphism (Table 2). S1 dips moderately NW and two
stretching lineation directions are apparent, plunging
shallowly to the NNE and west (Fig. 4). The NNEplunging lineation is defined by aligned peak-metamorphic sillimanite prisms and mineral aggregate ribbons. The -type porphyroclasts and asymmetrical
isoclinal folding document NNE over SSW transport
accompanying the dominant S1–L1 shear fabric. Consequently, this fabric is interpreted to have formed coeval
with the S1–L1 schistose foliation in the underlying LHS.
East-plunging lineations are unique to the Barun
Gneisses, not being recognized in the LHS (Figs 2 and
4). This lineation orientation is possibly due to refolding
by tight to isoclinal F2 folds with shallow NNE to NE
plunges. In the Jannu–Kangchenjunga Gneiss S1 is nearly
flat lying (Fig. 4), of variable intensity and stretching
lineations are only rarely apparent.
S1 is overprinted at low angles, by thin (millimetrescale), fine-grained sillimanite + biotite ± gedrite shearbands (S3) with NE- to ENE-plunging sillimanite and
quartz-aggregate lineations. S3 is common in the Barun
Gneiss and Jannu–Kangchenjunga Gneiss and less common within the uppermost Himal Group in the Kangchenjunga profile. The S3 foliation is interpreted to be
the latest episode of ductile shearing associated with the
MCT Zone. Sense of S3 shear is unknown in the area
investigated. This episode may correspond to the Neohimalayan event (Vannay & Hodges, 1996) (Table 2).
Tight to close mesoscopic F4 folds with shallow NNE to
NE plunge and crenulation cleavages with aligned biotite
are recognized in both the GHS and LHS. These may
be coeval with the large-scale folding giving rise to the
antiformal tectonic windows within the LHS in the
Eastern Himalaya (Figs 1 and 3).
DESCRIPTION OF ROCKS ALONG
THE MAKALU PROFILE
Greater Himalaya Sequences
Black Gneisses
The Black Gneisses are composed of biotite-rich schists
and schistose quartzo-feldspathic gneisses and are intruded by the tourmaline + biotite ± garnet-bearing
Makalu leucogranite. Schists have well-aligned matrix
assemblages of biotite + muscovite + quartz + albite
+ K-feldspar + ilmenite ± sillimanite ± tourmaline
(Table 3). Quartzites and foliated quartzo-feldspathic
gneisses have similar assemblages but with lower modal
proportions of micas and are often rich in tourmaline.
Typically biotite and muscovite are in equilibrium, although coarse muscovites both overgrow biotite and are
boudinaged with interstitial biotite growth. Muscovite
and feldspars both contain fibrolite inclusions and muscovite is overprinted by thin fibrolite seams. Chlorite is
typically retrogressive after biotite and muscovite.
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JOURNAL OF PETROLOGY
VOLUME 41
NUMBER 12
DECEMBER 2000
Fig. 4. Lower hemisphere equal-area projections of structural data from both the Makalu and Kangchenjunga profiles. Ε, mineral and mineralaggregate lineations (L1); Φ, tight to isoclinal fold axes (F2); Χ, poles to bedding and the sub-parallel metamorphic foliation (S1).
Namche Migmatite Orthogneiss
The Namche Migmatite Orthogneiss is a mixed unit
composed predominantly of augen orthogneiss and
highly migmatized quartzo-feldspathic gneisses with
subordinate calc-silicate, carbonate and mafic gneiss.
Both quartzo-feldspathic and metapelitic gneisses have
biotite + K-feldspar + plagioclase + quartz +
ilmenite + sillimanite matrix assemblages (Table 3).
Some metapelites also contain garnet, perthite, antiperthite and coarse-grained muscovites that overgrow
biotite and K-feldspar. Like Barun Gneiss samples, the
coarse foliated matrix assemblage is overprinted by
thin, fine-grained S3 seams of fibrolite + biotite +
muscovite + opaque minerals. Matrix ilmenite is
1680
(foli-QFG)
(foli-QFG)
(QFG)
(quartzite)
M73
M74
M70
M68∗
gn
X-S
S
S
S
S
S
SK
bi
1681
X
(pelite)
(pelite)
(pelite)
(pelite)
(pelite)
(pelite)
(pelite)
(pelite)
(pelite)
(pelite)
M11
M13
M16
M17
M18
M21
M22
M34
M38
M39∗
X
X
XS
X
X
X
X
X
X
Barun Gneiss, metapelites
(pelite)
M78∗
(QFG)
(pelite)
M75
MX
R
SL
X
SR
X
S
S
SR
S
mu
IgnXS C
IgnXS C
IgnXS
IgnXS
XS
Ign S R
IgnXS
IgnXS
XS
IgnXS
S
IgnXS F S-L R F
XS
XS
Namche Migmatite Orthogneiss
(schist)
(schist)
M71∗
(schist)
M69∗
M72∗
(schist)
M60
Black Gneisses
Sample
st
XS
XS
X
X
X
XS
X
XS
X
XS
X
X
XS
X?
X
X?
S
S
XS
kf
XS
XS
X
X
XS
XS
X
IgnXS
X
XS
X
X
XS
X
X
X
X
X-S
S
S
pl
XSC
S
S
XS
Ign S
S
XS
IgnXSC
XS
XSF
Iq S
S
S
SCF
Imu,q S
SK
sill
ky
Table 3: Petrology of aluminous rocks from the Makalu profile
Ign
sp
Ign X
IgnXS
XS
SE
Ichl
XR
R
Rmu
Imu
SR
SR
ilm
XS
IgnXS
Ign X
Ign X
XS
IgnXS
Ign X
XS
Ign S
IgnXS
XS
X
Ign S
IgnXS
Isill, gn IgnXS
IgnXS
X
XS
XS
X
X
X
X
Imu S
S
SR
S
q
E
E
Ibi
XF
R
X
R
E
hm
XS
E
E
mg
R
S
R
SR
R
SR
R
R
SR
R
chl
XC
X
X
X
cd
L
L
L
L
L
gd
gn→green bi
gn→bi, hm→mg,
bi→chl
cd→sill-gd
sill→fine sill,
ilm→hm±mg
gn→bi,
sill→fine sill
cd→gd,
cd→sill, gn→sill
gd→fine sill
sill→fine sill,
kf→mu, bi→mu
bi→ilm, ilm→hm
mu→bi
bi→mu, sill→chl,
ilm→hm
bi→mu, mu→fibr,
bi→chl±ilm
mu→ilm
bi→chl
bi→chl
Reaction textures
act
gn→bi-pl
gn→q-cd,
ap, zr, mon, gn→bi-sill-cd,
graphite
zr, rut
zr, rut
ap, tm
sph
act, ap, tm
tm
sph
tm
tm
Accessory
GOSCOMBE AND HAND
P–T PATHS IN EASTERN HIMALAYA
(pelite)
(pelite)
(pelite)
(pelite)
(pelite)
(pelite)
(pelite)
(pelite)
M47ii
M49
M51b∗
M54
M81
M82
M83
M85
1682
M92i
X
X
X
X
X
X
X
XC
Ign S
IgnXS
XS
S
S
Ign X
XS
XS
IgnXS
Ign S C
IgnXS
bi
R
R
S
mu
(QFG)
(QFG)
(QFG)
(QFG)
(QFG)
(QFG)
M14
M15
M52
M58∗
M92ii
M95
X
X
IgnXS
IgnXS
X
S
IgnXS
R
S
XCR
X
XS
X
X
X
X
X
XS
XS
X
XS
IXS
XG
X
XG
XC
XS
kf
S
sill
X
IgnXS
XS
X
X
XS
IX
XS
XS
X
XS
IgnXS
XG
X
XCS
XS
IgnXS
Imu Ccd
XSC
XS
S
XS
XS
XS
XS
S
X G S C IgnXS
XC
XS
pl
ky
Ign
sp
X
IgnXS
ilm
XS
IgnXS
Ign X
X
Ign X
IgnXS
Ign X
IgnXS
IgnXS
XS
XS
IgnXS
XG
IgnXS
IXS
XS
X
X
Ibi S
IgnXS
Ign X-S
XS
XS
SC
IgnXGS X-S
Ign X
IgnXS
q
X
C
E
X-S
hm
C
X-S
mg
R
R
R
R
R
R
R
chl
X
IgnXC
X
XC
X
X
cd
L
L
S
S
L
gd
ilm→hm
sill→ilm, bi→chl
bi→chl
cd→chl
gn→sill-chl,
ilm→bi→pl-q←gn
gn→chl-mg,
gn→bi-pl±sill
ilm→bi→myr→gn,
Reaction textures
zr, tm
sph, rut
zr
tm
zr, rut
zr, rut, py
rut
ap, zr, sph,
zr, ap
sill→gd
sill→fine sill,
sill, q→g→q→gn
cd→sill, sill→fine
gn→cd
gn→sill-bi-q,
pl→mu
cd→phengite-sill,
cd→sill, gn→cd
ilm→mg-hm
gn→q→gn
gn→chl
gn→green bi,
bi→chl, gn→chl
ap, phengite gn→phengite,
zr, rut
ap
zr
zr
rut
ap, zr, cc, py
Accessory
NUMBER 12
X
X
X
st
VOLUME 41
Barun Gneiss, quartzo-feldspathic gneisses
(pelite)
M47i
X
(pelite)
(pelite)
M44
gn
Sample
Table 3: continued
JOURNAL OF PETROLOGY
DECEMBER 2000
(foli-QFG) X
M95a∗
1683
X
X
X
(quartzite)
M99
(schist)
M111
X
S
Ign S
XS
X
S
Ist S
S
S
S
IgnXS
S
bi
S
S
S
XR
SL
S
S
S
XSC
XS
R
mu
X
X
X
X
PHmu
st
Epl S
S
kf
X
X
X
S
X
XS
XS
X
pl
S-L
sill
X
X
X
present
X
XS
XS
ky
sp
S
ilm
XS
IgnXS
S
XR
IstXS
S
C
hm
C
Ign,chl X S
S
Ibi
Ign,st S Ign,st S
Ist S
S
XS
X
Ist,gn
X
Ist,gn
XS
XS
Ign, pl Ign
XS
q
mg
S
S
R
S
R
S
R
SR
R
chl
cd
gd
bi→mu
ky→mu, bi→chl,
Reaction textures
tm
ap
piemontite
ap
ap
ap
ap,tm
tm
mu across chl
gn→chl-bi-q
bi→q-chl-ilm,
ky→q-mu
bi→chl, pl→mu,
mu across bi
st→chl±mu
gn→chl, bi→chl
gn→chl-hm
mu→chl,
st→mu→chl,
ky→seri←bi
tm, sph, rut ky→chl, ky→mu
ap, zr
Accessory
Samples grouped according to the mapped units (Fig. 2) and arranged from north to south down the page. I, inclusion phase with host indicated; X, peak
metamorphic phase; S, aligned or contained within the dominant fabric in equilibrium with peak metamorphic phases; X-S, coarse aligned matrix; C, corona
phase; K, crenulation phase; G, sub-grains; F, fine foliated seam; L, late phase; R, retrogressive phase; E, exsolution phase; PH, pseudomorphed phase; ∗, analysed
sample. Postscript indicates phase that is retrogressed or contains the inclusion. Reaction textures presented as reactant phases → product phases.
(schist)
M109∗
Himal Group, Zone C
(schist)
M97
Himal Group, Zone D
(schist)
M110
(schist)
M106∗
(schist)
(schist)
M104
M108∗
(schist)
M98∗
Himal Group, Zone E
(schist)
X
gn
M2∗
Himal Group, Zone F
Sample
GOSCOMBE AND HAND
P–T PATHS IN EASTERN HIMALAYA
JOURNAL OF PETROLOGY
VOLUME 41
overgrown by haematite coronas. Mafic gneisses have
quartz–plagioclase + K-feldspar + hornblende +
garnet + biotite + sphene + ilmenite assemblages
(Table 4).
Barun Gneiss
The Barun Gneiss is dominated by layered quartzofeldspathic paragneiss and metapelitic gneisses with
numerous thin bands of mafic gneiss, carbonate and calcsilicate. Quartzo-feldspathic paragneisses and metapelites
have identical assemblages and differ only by higher
proportions of feldspars and less garnet and sillimanite
in the quartzo-feldspathic gneisses. The coarse-grained,
foliated granoblastic matrix consists of garnet + plagioclase + K-feldspar + quartz + sillimanite + biotite
+ ilmenite ± rutile ± cordierite (Table 3) and most
samples contain partial melt segregations. The matrix
assemblage comprises the dominant coarse-grained foliation (S1), which is not significantly reworked apart from
minor fine-grained S3 foliation seams (Fig. 5d). Plagioclase
is oligoclase to andesine and rarely antiperthitic, and Kfeldspar is microcline and commonly perthitic. Kyanite
and staurolite are absent and muscovite is rare and
invariably a late-stage phase, though aligned with S1
(Table 3). Magnetite is rare and haematite commonly
occurs within S1 or as exsolution lamellae within primary
ilmenite grains, but is absent from inclusion assemblages.
In contrast to the LHS, garnet is commonly nearly
devoid of inclusions, which occur only in cores and are
absent from rims. Inclusions are typically biotite + quartz
+ ilmenite ± rutile and rarely also sillimanite and
hercynitic spinel (Fig. 5a) (Table 3). Spinel occurs only
as an inclusion phase and is absent from the matrix.
Fine sillimanite inclusions preserve an earlier, overgrown
crenulation cleavage. Matrix sillimanite is medium
grained, in textural equilibrium with and aligned with
the S1 matrix assemblage. Non-corrosive, fine-grained
sillimanite overgrowths occur on the margins of primary
sillimanite (Fig. 5d) and garnet. Garnet is consumed by
quartz + biotite + sillimanite ± plagioclase symplectites
(Fig. 5b) and overgrown by concentric quartz and garnet
coronas. Cordierite mostly occurs as stretched augen in
the matrix assemblage and rarely as cordierite ± quartz
or cordierite + sillimanite + biotite symplectites that
consume garnet porphyroblast margins. Fine-grained sillimanite + gedrite ± biotite and sillimanite + phengite
aggregates both replace cordierite and comprise the S3
foliation seams that are sub-parallel to, but overprint the
coarse-grained S1 fabric (Fig. 5d).
Mafic gneisses have a well-aligned granoblastic matrix
composed of quartz + labradorite + green hornblende
+ phlogopite + ilmenite; more rarely clinopyroxene
or garnet or both also occur (Table 4). Retrogressive
blue–green hornblende, chlorite or biotite occur on clinopyroxene margins. Garnet is rarely enclosed by coronas
NUMBER 12
DECEMBER 2000
of clinopyroxene or quartz + hornblende symplectites.
Calc-silicate gneisses have a granoblastic matrix of quartz
+ bytownite + meionite + clinopyroxene + orange
phlogopite + sphene + ilmenite ± green hornblende
± calcite (Table 4). Hornblende-free samples contain
forsterite. Matrix microcline is common and garnet is
rare. Epidote and clinozoisite are interstitial and appear
to be post-peak metamorphic phases. Forsterite also
occurs as coronas on meionite. Carbonates have very
similar matrix assemblages but are devoid of primary
hornblende, meionite is less common and wollastonite is
present in meionite-free samples. Wollastonite occurs as
a matrix phase and as coronas on clinopyroxene.
Lesser Himalayan Sequences
Himal Group
The Himal Group consists of thrust slices of aluminous
schists, schistose gneisses and meta-quartzites with subordinate calc-silicate and mafic gneisses. Aluminous
schists and meta-quartzites have coarse foliated matrix
assemblages of quartz + albite + biotite + muscovite
+ ilmenite + kyanite ± garnet ± staurolite (Figs 5c
and 6a and b) (Table 3). Chlorite and muscovite are
typically in equilibrium with the S1 matrix assemblage
but also occur as late-stage grains that overgrow S1 (Fig.
6a). Garnet and staurolite are often poikiloblastic (Fig.
6b) and typically have idiomorphic, inclusion-free overgrowths. Inclusions in both garnet and staurolite consist
of quartz, biotite and ilmenite. Matrix sillimanite, andalusite, cordierite and hercynitic spinel are entirely
absent from the Himal Group. Melt segregations are
also absent except thin leucocratic segregations in gneiss
sample (M95a) at the highest structural level (Zone F, see
below). Kyanite and staurolite margins are retrogressively
replaced by fine muscovite, often with an outer corona
of chlorite.
Metamorphic zones cannot be as accurately mapped
as in the Kangchenjunga profile. However, there is a
general progression of increasing metamorphic grade up
sequence from the Arun River, north to the Barun Thrust
at the base of the Barun Gneiss (Fig. 2). The following
metamorphic zones are recognized:
Zone C: garnet+biotite+muscovite+albite+chlorite;
Zone D: garnet+biotite+muscovite+albite+kyanite;
Zone E: garnet+biotite+muscovite+oligoclase+
staurolite+kyanite;
Zone F: garnet+biotite+oligoclase+K-feldspar+
kyanite±muscovite (late sillimanite).
Calc-silicate gneisses have a granoblastic matrix of
plagioclase + meionite + grossular + clinopyroxene
1684
q
pl
(mafic)
X
(cpx layer)
(pl layer)
(mafic)
M84∗
M84∗
M87∗
XG
X
X
X
X
X
X
X
X
Ign X X
X
X
Ipl,bi
XG
1685
(calc-silicate)
(calc-silicate)
(calc-silicate)
(calc-silicate)
(bi-rich layer) X
(calc-silicate)
(calc-silicate)
(calc-silicate)
M50
M80
M28
M24
M24
M57∗
M57∗
M79
X
X
X
X
X
X
Ign X
X
X
XG
X
X
X
X
X
Ign X X
X
X
(calc-silicate)
(calc-silicate)
M48
X
X
X
X
M42∗
(calc-silicate)
(calc-silicate)
(calc-silicate)
M27
M37a
(calc-silicate)
M26
M35
(calc-silicate)
M25
Barun Gneiss, calc-silicate gneisses
(mafic)
(hn layer)
M84∗
(mafic)
M52
M56
(mafic)
M51
Barun Gneiss, mafic gneisses
M77
Namche Migmatite Orthogneiss
Sample
X
X
X
X
X
X
X
X
X
XG
X
X
kf
X
X
X
X
X
X
X
X
X
X
X
me
X
Ipl X
X
X
X
X
gn
X
X
X
X
X
X
X
X
X
X
X
X
Ipl,gn
X
X
C
IX
cpx
XR
R
R
R
R
X
XR
X
XR
R
XR
C
X
XR
Ign X
X-S
X
X
X
hn
X
X
X
X
ol
X-L
X
XC
X-L
X
X-L
X
ep
X
X
X
X
X
X
X
X
R
cc
R
C
R
act/tr
X
X
X
X
Ign
X
X
X
X
X
X
X
X
X
X
sph
Table 4: Petrology of mafic, calc-silicate and carbonate rocks from the Makalu profile
X
X
X
X
X
XC
XC
X
X
X
Ipl X
X
XR
X
Ign X
X-S
X-S
Ipl X
X-S
X
R
R
R
R
bi/phl chl
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
ilm
X
R
X
hm
X
woll
piedmontite
zr
rut, ap
tm
ap
ap
rut
ap
ap
Accessory
fo→ep
cpx→hn
cpx→hn
cpx→hn, ilm→mg
cpx→hn
phl→act
cpx→phl, cpx→act,
cpx→hn, cpx→phl
ilm→mg, cpx→hn
cpx→hn
cpx→hn
cpx→hn
cpx→hn
cpx→chl
bi→chl, gn→chl,
cpx→bi
cpx→hn
gn→q-hn, gn→cpx
bi→chl, ilm→hn
Reaction textures
GOSCOMBE AND HAND
P–T PATHS IN EASTERN HIMALAYA
q
1686
(mafic)
M6
(mafic)
X
X
X
X
X
X
X
X
gn
X
R
XR
hn
XC
X
X
X
Ign X X R
X
X
X
X
X
X
X
X
cpx
X
X
XC
ol
X
X
C
X
X
X
X
X
X-L
ep
X
X
X
X
X
X
X
X
X
cc
C
R
R
R
act/tr
X
X
Icpx X
X
X
X
X
X
X
X
X
X
X
sph
X
X
X
X
X
X
X
R
R
R
bi/phl chl
X
X
X
X
X
ilm
X
X
hm
X
XC
X
woll
ap, zr
rut
ap
ap
ma
ma
Accessory
ep→cpx
cpx→hn, gn→ep
cpx→act
cpx→hn, cpx→chl
fo→hn±bi
cpx→woll, woll→tr
ilm→mg
woll→q-tr,
woll→act, pl→ma
cpx→woll,
cpx→act, bi→ma
me→fo
Reaction textures
Samples grouped according to the mapped units (Fig. 2) and arranged from north to south down the page. I, inclusion phase with host indicated; X, peak
metamorphic phase; S, aligned or contained within the dominant fabric in equilibrium with peak metamorphic phases; X-S, coarse aligned matrix; C, corona
phase; L, late phase; R, retrogressive phase; G, sub-grains or deformed; ∗, analysed sample. Postscript indicates phase that is retrogressed or contains the
inclusion. Reaction textures presented as reactant phases → product phases.
M103
X
X
X
X
X
X
X
me
NUMBER 12
Himal Group, Zones C–E
(calc-silicate) X
M6
X
(calc-silicate) Ign, cpx Ign X X
X
(calc-silicate)
X
X
X
M4∗
X
X
X
X
X
X
X
kf
M1∗
X
X
X
X
X
X
X
X
pl
VOLUME 41
Himal Group, Zone F
(carbonate)
(carbonate)
M57
M53
(carbonate)
M43
(carbonate)
(carbonate)
M41
M79
(carbonate)
M37
(carbonate)
(carbonate)
M33
M79
(carbonate)
M24
Barun Gneiss, carbonates
Sample
Table 4: continued
JOURNAL OF PETROLOGY
DECEMBER 2000
GOSCOMBE AND HAND
P–T PATHS IN EASTERN HIMALAYA
Fig. 5. (a) Hercynite, ilmenite and quartz inclusions in garnet. (b) Plagioclase + quartz symplectite corroding garnet in contact with biotite
corona around matrix ilmenite. (c) Matrix assemblage of kyanite + biotite + garnet + plagioclase + K-feldspar + quartz all in textural
equilibrium, from Zone F. (d) Granoblastic quartz matrix with sillimanite and garnet porphyroblasts, overprinted by fine sillimanite and biotite
in S3 foliation seams (indicated). (Note fine sillimanite overgrowths on primary sillimanite.) All microphotographs in plane-polarized light and
from Barun Gneisses [except (c)]. Field of view 1 mm across in (a) and (b) and 6·5 mm in (c) and (d).
+ green hornblende + sphene + haematite ± epidote
± quartz ± K-feldspar. Mafic gneisses have a wellaligned granoblastic matrix of quartz + andesine +
hornblende + epidote at lowest structural levels (in Zones
C–E) and bytownite + hornblende + biotite + garnet
+ meionite in a mafic calc-silicate from the highest
structural level (Zone F) (Table 4; Fig. 2). Retrogressive
blue–green hornblende coronas enclose clinopyroxene
and hornblende. Epidote is enclosed by the clinopyroxene
coronas, preserving a prograde metamorphic reaction.
Retrogressive epidote occurs as coronas on garnet porphyroblasts.
DESCRIPTION OF ROCKS ALONG
THE KANGCHENJUNGA PROFILE
Greater Himalaya Sequences
Jannu–Kangchenjunga Gneiss
The Jannu–Kangchenjunga Gneiss is composed of migmatized quartzo-feldspathic paragneisses, metapelitic
garnet + cordierite gneisses and augen orthogneiss, with
rare carbonate, meta-quartzite, calc-silicate and mafic
gneiss. Metapelitic and quartzo-feldspathic gneisses have
coarse, granoblastic textures and similar matrix assemblages of quartz + perthitic K-feldspar + plagioclase
+ orange biotite + sillimanite ± garnet ± cordierite
(Table 5). Garnet porphyroblasts typically have few or
no inclusions, these being ilmenite, biotite and quartz,
and inclusion-free garnet overgrowths are recognized.
Peak metamorphic sillimanite occurs both as porphyroblasts and as late-stage grains that overprint the
main foliation. Sillimanite is a common coronal phase,
occurring in aggregates of sillimanite + gedrite ± plagioclase, sillimanite ± biotite or sillimanite + phengite
replacing cordierite and sillimanite ± gedrite ± biotite
symplectites replacing biotite. As in the Barun Gneiss,
porphyroblastic garnet is partially consumed by coronas
and symplectites containing plagioclase, quartz, biotite
and sillimanite (such as Fig. 5b) (Table 5). Fine-grained
S3 seams of sillimanite ± gedrite are sub-parallel to and
overprint the coarse-grained S1 foliation. Kyanite and
staurolite are entirely absent from the Jannu–
Kangchenjunga Gneiss. Muscovite is a rare matrix phase
and typically replaces biotite.
Mafic gneisses have a coarse, granoblastic matrix assemblage of green hornblende + quartz + plagioclase
1687
JOURNAL OF PETROLOGY
VOLUME 41
NUMBER 12
DECEMBER 2000
Fig. 6. (a) Matrix staurolite + kyanite + quartz + biotite assemblage with late chlorite, from schist in Zone E. (b) Syn-tectonic garnet core
with inclusion-free rims in biotite + muscovite + staurolite schist from Zone E. (c) Syn-tectonic garnet core with post-tectonic idiomorphic
garnet overgrowth (indicated) in sericite + biotite schist from Zone B in the Seti Formation. (d) Syn-tectonic garnet overgrown by the biotite
+ sericite S1 metamorphic foliation (indicated), which in turn is overgrown by post-tectonic chlorite (indicated), also from Zone B. All
microphotographs in plane-polarized light and field of view 6·5 mm across.
+ orange biotite. Calc-silicates have assemblages of
plagioclase + meionite + clinopyroxene + forsterite
+ calcite + sphene and late-stage epidote (Table 6). The
Jannu–Kangchenjunga Gneiss unit contains an intensely
deformed biotite + sillimanite-bearing granitic augen
orthogneiss body (Fig. 3). At high structural levels there
are numerous, deformed leucogranite sills containing
tourmaline + biotite ± muscovite ± garnet ± sillimanite.
Lesser Himalayan Sequences
Biotite Gneiss
Fine-grained, biotite-rich schists dominate the wide ductile shear-zone between the Jannu–Kangchenjunga Gneiss
unit and the highest-grade portion of the Himal Group
(Zone G) to the south (Fig. 3). These schists have brown
biotite + plagioclase + K-feldspar + quartz ± garnet
assemblages with fine-grained sillimanite aligned within
the foliation and as inclusions in feldspars and garnet
(Table 5). Granitic pegmatites and veins are particularly
common.
Himal Group
The Himal Group consists almost entirely of metapelite
schists and schistosic gneisses, with uncommon quartzite,
amphibolite and calc-silicate units of only 0·5–2 m thickness and few mappable (<100 m thick) calc-silicate units.
This monotonous sequence of metapelitic schists permits
the mapping of diagnostic metamorphic assemblages,
and these display an increase in metamorphic grade to
higher structural levels (inverted isograds). The following
matrix assemblage zones are recognized, with increasing
grade from south to north (Fig. 3; Table 5):
Zone D: garnet+biotite+muscovite+oligoclase±
kyanite;
Zone E: garnet+biotite+muscovite+oligoclase+
staurolite±kyanite;
Zone F: garnet+biotite+bytownite+K-feldspar+
kyanite;
Zone G: garnet+biotite+bytownite+K-feldspar+
sillimanite±kyanite.
The matrix assemblage consists of coarse granoblastic
phases in textural equilibrium and is dominated by wellaligned (S1) red–brown biotite sheets (Figs 5c and 6a
1688
gn
(pelite)
(schist)
(pelite)
(pelite)
(pelite)
(ortho)
K13
K67a
K67b
K67c
K71a∗
K77
1689
(QFG)
(pelite)
(pelite)
K65b
K65c∗
K65d
X
X
X
X
X
X
PHchl
X
(pelite)
(pelite)
(pelite)
(pelite)
(pelite)
(pelite)
K6
K8a
K8b
K9
K62a
K62b∗
(pelite)
(pelite)
K18a
K18b
K60a
(pelite)
(pelite)
(pelite)
K60b
K60c∗
K60d
Himal Group, Zone E
(pelite)
(pelite)
K1
Himal Group, Zone D
(pelite)
K5
X
X
X
X
X
X
X
X
X
X
Kushma Formation, Zone D
(pelite)
K17
Ulleri Formation, Zone C
(pelite)
K12
Seti Formation, Zones A and B
Sample
X
X-S C
X-S
X
X-S
X-S L
Ikf S K
X-S
S
S
X
S
X-S
X-S
Ipl X
S
XR
S
SK
S
S
S
S
mu
present
st
X-S
X-S
present
Ign X-S Ign X-S present
X
X-S
X-S
X
X-S
X-S
SK
X-S
S
S
X
Ign S
X-S
X-S
X
S
S
XS
S
X-L
S
bi
X
G
kf
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
XG
X
pl
sill
Iky X
X
X
X
Ign X
Ign X
Ign X
Ign X
Ign X
Ign X
X
X
XG
Ign X
Ign X
Ign X
X
X
Ign X
q
X
X
Ign X
present X
present Ign X
X
present Ign X
X
X
X
ky
Table 5: Petrology of aluminous rocks from the Kangchenjunga profile
X
Ign X-S
X
Ign X
Ign X
X
Ign X
X
X
Ign L
Ign X
Ign X
Ign X-S
Ign X
Ign X-S
X
X
Ign X
SK
Ign X-S
X
L
Ign S
ilm
Eilm
X
hm
X
X
Xtimg
X
mg
Rbi
SK
Rgn
Rgn,bi
Rgn
S-L Rgn
X-S L
R
S-L
Rgn S
S-L Rgn
S
L Rst
L Rgn
chl
cd
gd
zr, ap
rut
zr
tm
tm
tm
tm
tm
zr
tm
tm
tm
Accessory
ky→mu
mu laths across bi
foliation
ilm across
chl across bi-mu
foliation
bi laths across mu
foliation
ilm and chl across
foliation
chl across
Reaction textures
GOSCOMBE AND HAND
P–T PATHS IN EASTERN HIMALAYA
(pelite)
(pelite)
K61b∗
K61c
(QFG)
(pelite)
(pelite)
(pelite)
K15cii
K19a
K21a
K21b
K58
(pelite)
(pelite)
(pelite)
(pelite)
(pelite)
(pelite)
(pelite)
K25b
K26a
K26b
K26c
K27
K28b
K29
1690
(pelite)
(pelite)
(pelite)
(pelite)
(pelite)
(pelite)
(pelite)
K35d
K52
K53a
K53b
K53c
K53d
X
X
X
X
X
X
X
X
X
X
X
X Ipl
X
X
X
X
X
X-S
X-S
XF
X-S
X-S
S
X
X-S
X-S
S
X-S
X-S
Ign X-S
Ign X-S
X-S
X-S
X
X-S
X-S
X-S
X-S
X-S
Rfeld
Rbi
Rbi,kf
SR
X-S L
X-S
st
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
Ipl
kf
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
pl
Ign S-L
C
Ign S-L
Cbi F
feld S
Ign,
Ifeld
X-S C F
X-S C
X-S
X-S C
X-S C
S
X X-S
S-L C
S-L F
S-L F
S-L F
S-L F
sill
Ign X
q
X
Ign
X
X
X-S
Ign X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
Ign X
Ign X
X
X
X
Ign X
Ign X
X
X
X
present Ign X
ky
X
X
X
Ign X
X
X
X
Ign X
Ign X
X
Ign X
S
Ign X
Ign X-L
X
X
Ign X-S
Ign X
ilm
S-L
hm
X
X
X
Xtimg
mg
Rbi
Rgn
Rbi
Rbi
R
Rgn
Rgn
X-S
Rbi
X-S L
chl
cd
S-L
L
gd
py
tm
zr
ap
ap, tm
zr, rut
zr
zr
zr
zr
zr, tm
ap, sph
Accessory
foliation
sill inclusion
foliation
gn→sill, sill inc
bi→sill
foliation
sill inclusion
foliation
sill inclusion
sill→fine sill
sill→fine sill
sill→fine sill
sill→fine sill
bi→sill
foliation, gn→chl
sill across
gn→cc-chl
bi-chl
mu laths across
Reaction textures
NUMBER 12
K35
X
X
mu
Ign X-S Ign X-S
bi
VOLUME 41
Biotite Gneiss, MCT
(pelite)
K25a
Himal Group, Zone G
(pelite)
(pelite)
K15ci
Himal Group, Zone F
X
(pelite)
K60e
X
gn
Sample
Table 5: continued
JOURNAL OF PETROLOGY
DECEMBER 2000
(QFG)
(QFG)
(pelite)
(pelite)
(pelite)
(orthogneiss)
(QFG)
(pelite)
(pelite)
(pelite)
(QFG)
K40b
K40c
K41a
K41b
K43
K45a
K45b
K46b
K46d
1691
K47
K48
X-S
X-S
(schlieren)
(schlieren)
(leucogranite)
K41dii
K41di
K41e
X
X
X
X
Rfeld
X
Rbi, phl
X
Rbi, feld
X Rfeld
Rbi
X-S
S-L Cbi
Ign X-S Rbi
Ign X-S
F
X
Rbi
X
C
Ign XC X-L
Rbi, feld
st
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
XRbi
X
kf
X
X
X
X
X
X
X
X
X
X
XC
XC
X
X
X
X
X
C
X
pl
X S-L
S-L F
S-L F
X S-L C
F
X
S-L C F
X S-L F
S-L C F
Ifeld S
F
Ifeld C
S-L C
Ifeld
S-L F
S-L
S-L
sill
ky
X
X
X
X
X
Ign X
X
Ign X
X
X
Ign X
Ign X
Ign X
X
Ign X
X
Ign X
X
IgnXC
X
q
Rbi
X
X
X
Ign
X
Ign
Ign X
Ign X
X
X
Ign
X
X
ilm
F
hm
Xtimg
mg
X PH
X
cd
Rbi
Rbi
Rbi
Rbi
Rbi
Rbi
Rbi
Rbi
Rbi
PH?
Rbi
X?
Rbi,gn
FRgn
chl
F
F
F?
R
R
CF
C
gd
ep
tm, ap, phl
ap, zr
ap, zr
ap, zr
zr
zr, ap
zr
tm
tm, py
rut
rut
zr
tm
zr
Accessory
bi→mu
bi→sill
bi→sill
gn→sill-pl-bi
bi→mu, gn→bi-pl,
gn→gn over
cd→sill-gd
sill-phengite
pseudomorphs
bi→sill-gd-bi
cd→gd-sill,
cd→fine sill±bi
cd→sill-gd±pl,
gn→pl→q→sill-gd
Reaction textures
Samples grouped according to the mapped units (Fig. 3) and arranged in increasing metamorphic grade down page, that is, from south to north. I, inclusion
phase with host indicated; X, peak metamorphic phase; S, aligned or contained within the dominant fabric in equilibrium with peak metamorphic phases; X-S,
coarse aligned matrix; C, corona phase; K, crenulation phase; G, sub-grains; F, fine foliated seam; L, late phase; R, retrogressive phase; PH, pseudomorphed
phase; ∗, analysed sample; present, recognized in hand specimen. Postscript indicates phase that is retrogressed or contains the inclusion. Reaction textures
presented as reactant phases → product phases.
(schlieren)
K41c
X-L C
Rbi
X-S
Rbi
X-L Cbi
mu
X-S
Rbi, cd
Ign X-S S-L C
Rbi, pl
Ign X-S
Leucogranite in Kangchenjunga Gneiss
X
X
X
X
X
XC
X
X
X
X
X-S
(pelite)
(pelite)
(QFG)
K37b
K39
K40a
X
X
X
X
(pelite)
K37a
bi
X
gn
Kangchenjunga Gneiss
K36
(pelite)
Sample
GOSCOMBE AND HAND
P–T PATHS IN EASTERN HIMALAYA
q
pl
kf
me
(mafic)
K71c
X
X
X
X
(mafic)
K65a
X
X
X
X
X
X
(mafic)
X
X
1692
(mafic)
X
X
(calc-silicate)
K24
X
X
X
X
(mafic)
(calc-silicate)
K32a
K32b
(carbonate)
K46c
X
X
X
X
X
X
X
X
X
X
XC
X
X
X
X
X
C
XC
X
R
X-S
X-S
X-S
X X-S
X-S
XS
X
hn
X
X
X
ol
L
X-L
R
R
X-L
X
X-L
X
ep
X
X
R
R
X
cc
act
X
X
X
XC
C
X
X
X
X
X
sph
X
X
X
X
X
X
X-S
X-S
bi/phl
Rbi
Rbi
Rbi
Rbi
chl
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
op
tm
tm
tm
rut
ap
Accessory
pl→me
cpx→hn
op→sph, cpx→hn
hn→sph
cpx→hn
hn→gn
Reaction textures
Samples grouped according to the mapped units (Fig. 3) and arranged in increasing metamorphic grade down page; that is from south to north. I, inclusion
phase; X, peak metamorphic phase; S, aligned or contained within the dominant fabric in equilibrium with peak metamorphic phases; X-S, coarse aligned matrix;
C, corona phase; R, retrogressive phase; #, out of situ. Postscript indicates phase that is retrogressed or contains the inclusion. Reaction textures presented as
reactant phases → product phases.
(mafic)
(mafic)
K38
K46a
X
X
X
X
X
cpx
NUMBER 12
Kangchenjunga Gneiss
(mafic)
K28a
X
C
gn
VOLUME 41
Himal Group, associated with Zone G metapelites
(carbonate)
K15a#
Himal Group, associated with Zone F metapelites
K61a
Himal Group, associated with Zone E metapelites
K63
Kushma Formation, associated with Zone D metapelites
(mafic)
(mafic)
K11
K64
Ulleri Formation, associated with Zone C metapelites
(calc-silicate)
K71b
Seti Formation, associated with Zone A-B metapelites
Sample
Table 6: Petrology of mafic; calc-silicate and carbonate rocks from the Kangchenjunga profile
JOURNAL OF PETROLOGY
DECEMBER 2000
GOSCOMBE AND HAND
P–T PATHS IN EASTERN HIMALAYA
and b). Garnet porphyroblasts typically have few or no
inclusions, those observed being typically quartz, biotite
and opaque minerals and, at highest grades (Zone G),
aligned fine-grained sillimanite. Kyanite occurs as wellaligned porphyroblasts in textural equilibrium with the
matrix assemblage. In Zones F and G there is latestage sillimanite aligned with S1 and thin, fine-grained
sillimanite ± gedrite ± biotite ± haematite S3 seams.
At highest grades (Zone G), sillimanite occurs as porphyroblasts in textural equilibrium with a matrix assemblage that is almost devoid of kyanite (Table 5).
Partial melt segregations are recognized in metapelites
in Zone G. In the high-grade zones, kyanite occurs both
as laths in equilibrium with the matrix assemblage and
as inclusions in garnet porphyroblasts. Primary sillimanite
and biotite grains are overgrown by coronas of finegrained sillimanite (Table 5).
Mafic gneisses also illustrate an increase in metamorphic grade from south to north, with matrix assemblages of quartz + plagioclase + green hornblende
+ epidote + sphene associated with Zones C–E metapelites and quartz + plagioclase + green hornblende
+ biotite ± clinopyroxene associated with Zone G
metapelites (Table 6). At highest grades epidote and
sphene are late-stage phases, with sphene forming coronas
on opaque minerals and hornblende. Calc-silicates in
Zones F and G have granoblastic matrix assemblages of
quartz + plagioclase + clinopyroxene + calcite +
sphene + orange biotite ± grossular with late-stage
blue–green hornblende and epidote (Table 6).
sphene ± biotite assemblages, rarely with small latestage epidote grains (Table 6). The basal Ulleri Formation
is a mixed zone containing metapelitic schists with biotite
+ muscovite + quartz + K-feldspar ± plagioclase ±
garnet assemblages and late-stage chlorite laths overprinting the main foliation (Table 5).
Seti Formation
The Seti Formation is a monotonous sequence of finegrained, pale grey–green quartz–sericite schists with rare
mafic schists, calc-silicate and quartzite units. Biotite
occurs in the quartz–sericite schists (Zone A); both within
the fine-grained foliation and as aligned (L1) blocky laths
that overgrow the sericite foliation. Garnet is common
in the highest structural levels (Zone B) and occurs as
large snowball poikiloblasts with post-tectonic idiomorphic overgrowths [Fig. 6c; see also Meier & Hiltner
(1993)]. Chlorite is mostly a late-stage phase, such as
coarse laths overprinting the sericite foliation (Fig. 6d)
and retrogressive after garnet. Feldspars are nearly absent
except for rare albite. Mafic schists have quartz +
plagioclase + green hornblende ± epidote assemblages
with garnet coronas on hornblende (Table 6). Deformed
biotite + muscovite + tourmaline leucogranite bodies
are emplaced in the core of the antiformal dome
(Fig. 3).
Kushma Formation
The Kushma Formation is dominated by meta-quartzites
and metapelitic schists with subordinate mafic schists.
Metapelitic schists have a coarse granoblastic matrix with
well-aligned micas and assemblages of garnet + quartz
+ albite + biotite + muscovite + tourmaline ± kyanite
(Table 5), constituting Zone D assemblages. Garnet
porphyroblasts are typically poikiloblastic and snowball
garnets, commonly with inclusion-free rims. Late-stage
crenulation cleavages with axial planar muscovite +
biotite ± chlorite are common. All other chlorite is
retrogressive after garnet and biotite. Mafic schists have
quartz + plagioclase + blue–green hornblende + ilmenite + rutile assemblages.
Ulleri Formation
The Ulleri Formation is a 3–4 km thick unit of intensely
sheared granitic orthogneiss (Fig. 3), with perthite augen
in a granoblastic matrix of quartz + plagioclase + Kfeldspar + muscovite with minor biotite and sieve garnet.
Mafic schists of 1 m width are at very low angles to the
S1 foliation and interpreted as transposed mafic dykes.
These have quartz + plagioclase + hornblende ±
MINERAL CHEMISTRY
Analyses from the Makalu profile were performed on the
Cameca SX50 electron microprobe at the University of
Tasmania, and from the Kangchenjunga profile on the
Camebax microbeam at the University of Cape Town.
For both, operating voltage was 15 kV and current was
20 nA for all phases except micas (10 nA) and feldspar
(15 nA), and beam radius was 2·2 m for most phases
and 4·4 m for micas and feldspars. The following natural
silicates were used as standards and checked periodically
throughout sessions. For garnet, staurolite and aluminosilicates: Kakanui pyrope, Kakanui hornblende, rutile
and Chromite 52NL-11; for micas, amphiboles and pyroxenes: Kakanui hornblende, rutile and rhodonite; for
feldspars: Lake County plagioclase, Nuni albite, Orthoclase-1 and Kakanui pyrope. The range in mineral
chemistry, within each rock-type group in the Makalu
profile, is summarized in Table 7. Representative mineral
analyses are presented in Appendix C and the complete
dataset of mineral analyses used is in the Journal of Petrology
website at http://www.petrology.oupjournals.org. Mineral end-member activities were calculated largely after
1693
JOURNAL OF PETROLOGY
VOLUME 41
Powell & Holland (1985); the formulations used are
summarized in Appendix D.
THERMOCALC RESULTS
Average P–T conditions of equilibration of core and rim
assemblages were determined by the method of Powell
& Holland (1985, 1988). Calculations were performed
using the 1992 thermodynamic dataset and computer
program THERMOCALC v2·0b (Powell & Holland,
1988); the results are presented in Appendix A. All results
satisfy the 2 test and errors from THERMOCALC,
incorporating typical uncertainty for each mineral endmember activity and errors in the thermodynamic dataset, average ±44°C and ±1·0 kbar (Appendix A). Where
there are insufficient independent reactions to calculate
average P–T loci, the intersection of solutions for average
P and average T calculations by THERMOCALC have
been used to define equilibration conditions (Appendix
A; Table 8). The best constrained P–T loci of both the
maximum-T and maximum-P conditions preserved in
each sample are summarized in Table 8 and represented
in Fig. 7. Almost all samples (except M95a) were calculated by THERMOCALC. As a result of using this one
method and this one internally consistent thermodynamic
dataset, the resultant P–T calculations are considered
directly comparable with each other.
There were insufficient mineral end-members in the
garnet + biotite quartzo-feldspathic gneiss assemblages,
from the Black Gneisses, Biotite Gneiss and Jannu–
Kangchenjunga Gneiss samples, to calculate a set of
independent reactions and obtain average P–T results by
THERMOCALC. Estimates of the equilibration conditions of these rock units, based on the stability field of
matrix assemblages, are discussed below (Figs 7–9) and
presented in Table 8. Average P–T results by THERMOCALC were, however, obtained from all other rock
units. These results cluster into three distinct groups,
corresponding to peak metamorphic conditions in the
Barun Gneiss, Barun Gneiss samples that have re-equilibrated during cooling, and peak metamorphic conditions in the Himal Group samples (Fig. 7; Table 8).
Peak metamorphic results from Barun Gneiss metapelites
average 837 ± 59°C and 6·7 ± 1·0 kbar (Table 8),
indicating an average geothermal gradient of 36 ± 4°C/
km. These peak results are identical to the phase stability
field of Barun Gneiss assemblages (Fig. 8), and thus
considered an accurate estimate of the peak of metamorphism (Fig. 7).
THERMOCALC results from Barun Gneiss rims are
of lower T (T = 24–56°C) and slightly lower P (P =
0·3–1·4 kbar) than core results (Fig. 7; Appendix A),
although little significance can be attributed to core vs
rim results because T and P are of similar magnitude
NUMBER 12
DECEMBER 2000
to the errors (typically ±44°C and ±1·0 kbar) on these
results (Appendix A). Barun Gneiss garnets preserve
zoning of contrasting style to garnets from the LHS; with
flat Ca patterns and increasing Fe and Mn and decreasing
Mg from core to rims (Fig. 10). These patterns are typical
of high-grade garnets with retrograde zoning caused by
diffusion (Tracy, 1982; Tuccillo et al., 1990). Consequently, rim calculations may reflect only cation decoupling; where Ca remained relatively immobile and
Fe–Mg continued to exchange during cooling (e.g. Tuccillo et al., 1990). As a result, the assemblages evolved
along isopleths of the reaction
3 anorthite = grossular + 2 aluminosilicate+quartz,
giving rise to the low, positive P/T, core to rim P–T
array (Fig. 7). It is important to realize that this P–T
array does not necessarily coincide with the P–T path
experienced by the rocks (e.g. Frost & Chacko, 1989).
Calc-silicate and mafic gneiss samples from the Barun
Gneiss give results clustered around an average of 674
± 33°C and 5·7 ± 1·1 kbar. These results are similar
in pressure, but significantly lower in temperature (T =
75–125°C) than the phase stability fields of the matrix
assemblages (Fig. 9). Consequently, these samples are
interpreted to have been re-equilibrated at some stage
subsequent to the peak of metamorphism. The calculated
P–T loci represent either conditions at cessation of cation
exchange (Harley, 1992) or re-equilibration in a later
metamorphic event (see below). Re-equilibration of matrix mineral compositions in a later metamorphic event,
without pervasive recrystallization of the matrix phases,
has been documented in the Ungava Orogen (St-Onge
& Ijewliw, 1996) and Zambezi Belt (Goscombe et al.,
1998). Unlike the calcareous rock-types, metapelites and
quartzo-feldspathic gneisses in the Barun Gneiss did not
experience later re-equilibration and peak metamorphic
mineral compositions were apparently preserved.
Himal Group samples equilibrated at P–T conditions
entirely distinct from all Barun Gneiss samples (Fig. 7).
Calculations from syn-kinematic garnet cores vs postkinematic rims commonly document heating (T typically 4–36°C and one sample of 106°C) accompanying
decompression (P up to 2·0 kbar) from near peak-P to
peak-T conditions (Fig. 7). Himal Group samples from
Zones D–F preserve peak metamorphic conditions averaging 609 ± 42°C and 8·8 ± 1·1 kbar, with the
highest pressures recorded in the Kangchenjunga profile
(Table 8). These results indicate an average geothermal
gradient of 20 ± 2°C/km. The maximum-P conditions
indicated by garnet core analyses, where significantly
different from those for rims, average 557 ± 39°C and
10·0 ± 1·2 kbar in Zones D–F (Table 8). Heating
accompanying decompression immediately before the
peak T of metamorphism, as suggested by these P–T
calculations, is typical of collisional orogens controlled
1694
0·03–0·06
0·00–0·01
X(Mn,M2)
X(Fe
0·44–0·72∩
Fe2O3 wt % 0·22–2·21
1695
0·10
0·68–0·93
0·69–0·92∩
X(me)
X(cc)
Scapolite
Carbonate
0·24–0·64
0·10
0·06–0·07
0·59–0·72
0·43–0·51
0·18–0·65∩
0·30–0·45∩
1·19–2·09∩
0·26–0·44ℜ
0·98–0·99
0·80–0·84
0·23–0·54
0·32–0·44
0·43–0·53
0·14–0·98
0·26–0·43
0·40–0·53∩† 0·41–0·49
0·61–0·65ℜ
0·26–0·55ℜ
0·06–0·07ℜ
0·80–0·83∩
0·21–0·39
0·36–0·43
0·34–0·51
0·23–1·51ℜ
0·27–0·44
0·41–0·49†
0·61
0·43–0·52
0·41–0·56
0·06–0·29
0·23–0·38
0·63–0·67
0·42–0·58
0·11–0·12
0·05–0·22
0·25–0·45
0·41–0·56∩
0·06–0·29∩
0·27–0·38ℜ
0·64–0·69ℜ
0·41–0·58∩
0·01–0·22ℜ
0·21
0·11–0·23ℜ
0·85–0·88ℜ
0·00–0·02∩
0·02–0·11∩
0·11–0·15ℜ
0·03–0·13∩
0·72–0·83ℜ
rim
0·78
0·49–0·50
0·61–0·74
0·49
2·53–4·27
0·23–0·42
0·07–0·09
0·79ℜ
0·59–0·60ℜ
0·60–0·69∩
0·49
1·12–4·32
0·27–0·43ℜ
0·07–0·10
0·48–0·96
0·33–0·35ℜ∗
0·45–0·96
0·03–0·05
0·02–0·10
0·10–0·12
0·44–0·47∩
0·36–0·44ℜ
rim
0·26–0·34
0·03–0·05
0·02–0·10
0·10–0·11
0·43–0·48
0·37–0·43
core
calc-silicate
0·51–0·61
0·08–0·10
0·33–0·46
0·66–0·69
0·54–0·67
0·08–0·13
0·06–0·13
core
0·56–0·60†
0·08–0·10
0·38–0·54ℜ
0·65–0·68∩
0·54–0·64
0·15
0·06–0·13
rim
biotite–muscovite schist
Black Gneisses
∗In mafic schists.
†Retrograde or late-stage phase; all other phases are primary matrix phases.
a
Ten per cent of FeOtotal calculated as Fe2O3; b100% of FeOtotal calculated as Fe2O3; c15% of FeOtotal calculated as Fe2O3; d60% of FeOtotal calculated as Fe2O3. Arrows
indicate typical direction of change of component from core to rim.
For micas; (A1,T1) signifies aluminium cations in T1 tetrahedal sites and (Na,A) signifies sodium cations in A sites (Powell & Holland, 1985).
X(ep)
Epidote
Na+K
0·29–0·46
Al2O3 wt %
Hornblendec XFe
0·24–0·44
1·86–2·29
XFe
Cpx
0·41–0·53
0·61–0·64
0·25–0·54
XFe
0·05–0·09∩†
0·21–0·55∩†
0·67–0·69
0·46–0·58∩
0·11–0·17∩
Chloritec
0·05–0·09
X(Na,A)
0·67–0·70
X(Al,T1)
0·27–0·58
0·48–0·58
XFe
Biotitec
XFe
0·10–0·19
X(ab)
K-feldsparb
Muscovited
0·12–0·37
Fe2O3 wt % 0·38–1·40
Fe2O3 wt %
0·84–0·88
0·00–0·02
0·00–0·13
0·04–0·16
0·03–0·16
0·69–0·84
core
0·15
0·56–0·86ℜ
0·00–0·03
0·02–0·04ℜ
0·12–0·29∩
0·18–0·23∩
0·53–0·65ℜ
rim
ZnO wt %
0·49–0·76
0·00–0·02
0·02–0·04
0·13–0·31
0·19–0·24
0·50–0·63
core
XFe
0·18–0·38
0·27–0·43ℜ
0·00–0·02
0·06–0·12ℜ
0·09–0·18∩
0·04–0·06∩
0·70–0·80ℜ
0·26–0·42
,M1)
0·14–0·27
X(Mg,M2)
3+
0·03–0·09
XFe
0·60–0·78
X(Fe,M2)
X(Ca,M2)
rim
core
core
rim
aluminous schist
mafic
metapelite
calc-silicate
Himal Group
Barun Gneiss
Plagioclaseb X(an)
Sillimaniteb
Kyaniteb
Staurolitea
Cordierite
Garnet
Phase
Table 7: Summary of mineral compositional ranges for rock-type groups in the Makalu profile
GOSCOMBE AND HAND
P–T PATHS IN EASTERN HIMALAYA
JOURNAL OF PETROLOGY
VOLUME 41
NUMBER 12
DECEMBER 2000
Table 8: Best estimates of metamorphic conditions of equilibration of samples from both profiles
Sample
Rock type
Peak-T conditions
P (kbar)
Peak-P conditions
T (°C)
P (kbar)
Method∗
T (°C)
Lesser Himalayan Sequences
Seti Formation, Zone B
K71a
gn–bi–mu–chl schist
7·4±1·9
561±11
8·3 assumed
490±32
8·5±1·9
561±12
(1) P–T, T & P
465±30
(1) T
Himal Group and Ulleri Formation, Zone C
M109
gn–ky schist
K65c
gn–bi–mu–chl schist
10·3±1·1
10·0 assumed
587±14
(1) P–T
Himal Group and Kushma Formation, Zones D and E
M106
gn–ky–st schist
8·4±1·0
650±20
9·7±1·1
646±21
(1) P–T
M108
gn–ky–st schist
8·0±1·0
662±21
9·7±0·8
556±20
(1) P–T
K61b
gn–ky schist
9·7±1·1
596±13
K60c
gn–ky schist
8·8±1·5
537±73
9·1±1·5
521±69
K62b
gn–ky schist
9·6±1·1
596±14
8·9±1·1
608±28
9·5±1·1
574±37
average
(1) P–T
(1) P–T
(1) P–T
Himal Group, Zone F
M1
calc-silicate
9·8±0·9
570±37
M2
gn–ky schist
8·0±1·2
612±95
9·9±1·9
600±95
(1) P–T
(1) P & T, (2) P
M4
calc-silicate
8·7±1·0
542±56
10·7±1·1
506±31
(1) P–T & T, (2) P
M95a
gn–ky gneiss
8·6±1·0
713±50
(2) P & T
Biotite Gneiss, MCT in Kangchenjunga profile
All samples
sill–gn–kf–bi±gd
>5·0–7·5
>670–750
(3)
>4·0–7·5
>670–710
(3)
Greater Himalayan Sequences
Black Gneisses
All samples
sill–mu–kf–bi(±cd)
Barun Gneiss and Namche Orthogneiss
M39
gn–cd–sill metapelite
6·8±1·0
832±76
(1) P–T
M51B
gn–cd–sill metapelite
6·5±0·8
823±60
(1) P–T
M58
gn–cd–sill QFG
7·0±0·6
877±55
(1) P–T
M78
gn–sill metapelite
6·5±1·6
818±47
(1) P–T
6·7±1·0
837±59
Average
All samples
metapelite & mafic
>5·5–7·0
>780–820
(3)
Barun Gneiss samples that have been re-equilibrated
M42
calc-silicate
5·7 assumed
695±18
(1) T
M57
calc-silicate
5·7 assumed
700±30
(1) T
M84
gn–hn–cpx–bi mafic
5·4±1·2
664±37
(1) P–T
M87
gn–hn–cpx–bi mafic
6·0±1·1
635±47
(1) P–T
5·7±1·1
674±33
Average
Jannu–Kangchenjunga Gneiss
All samples
QFG & calc-silicate
>4·0–7·0
>710
(3)
S3 assemblages, Barun Gneiss, Jannu–Kangchenjunga Gneiss and Himal Group Zone F
All samples
sill–gd–bi±gn
>4·5–7·0
>610–750
(3)
∗Methods: (1) THERMOCALC results (Appendix A); average P–T calculations used preferentially; with average T or average
P calculations used in few samples as indicated. (2) Conventional geothermobarometry (Appendix B). (3) Estimate based
on phase stability field of matrix assemblage (see text). Error on average for rock group; is the average of result errors.
Makalu profile samples indicated by prefix M and Kangchenjunga profile by prefix K.
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GOSCOMBE AND HAND
P–T PATHS IN EASTERN HIMALAYA
Fig. 7. Summary of results of P–T calculations (Table 8) from both the Makalu and Kangchenjunga profiles. P–T estimates are largely calculated
using THERMOCALC v2.0b (Powell & Holland, 1988) (Appendix A), with some by conventional geothermobarometry (Appendix B). Errors
not included on figure for clarity; THERMOCALC errors average ±35°C and ±1·05 kbar in the Barun Gneisses and ±39°C and ±1·20 kbar
in the Himal Group (Appendix A) and geothermobarometry errors are assumed to be of the order of ±50°C and ±1 kbar (Kohn & Spear,
1990). Closed symbols are cores, open are rims, small symbols are samples constrained by an assumed P and stars are average P–T loci (Table
8). Labelled metamorphic zones are as defined in the text and Fig. 8. Mineral equilibria from literature are: (1) Fe-staurolite stability field in
FASH (Bickle & Archibald, 1984); (2 and 4) theoretical relations in KFMASH (Powell & Holland, 1990); (3) experimental reactions in tholeiite
compositions with quartz–fayalite–magnetite (QFM) buffer (Spear, 1981); (5) experimental garnet-in reaction for mid-ocean ridge basalt of XFe =
40% composition (Green & Ringwood, 1967); (6) Grant (1985); (7) Bohlen & Dollase (1983); (9) theoretical relations in FMASH (Hensen &
Green, 1973; Baker et al., 1987).
by uplift and erosion (England & Thompson, 1984; Spear
et al., 1984).
As is typical at mid-amphibolite facies conditions, all
Himal Group and Midland Group garnet porphyroblasts,
except those closest to the MCT (i.e. sample M95a,
which underwent retrograde diffusion), still preserve
growth zoning from syn-kinematic cores to idiomorphic
post-kinematic rims (Fig. 10). Garnet compositional profiles are typical of growth zoning with increasing Fe and
Mg and decreasing Ca and Mn towards rims (Tracy,
1982; Loomis, 1983; Tuccillo et al., 1990). Thus core to
rim P–T trajectories are considered at least broadly
representative of changing P–T conditions during garnet
growth, although T and P are of magnitude not much
greater than the errors on these results (Appendix A).
Core and rim calculations are interpreted to represent a
portion of the prograde P–T path as argued by Spear et
al. (1984) and St-Onge (1987). In this interpretation,
garnet cores forming by syn-tectonic crystallization at
the same time as the matrix assemblage, possibly near
peak-P conditions, and idiomorphic and inclusion-free
rims and overgrowths represent post-kinematic garnet
growth at peak-T conditions.
The continued growth of garnet on a heating and
decompressional P–T path, as suggested by the textural
relationships and thermobarometry, at first glance
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Fig. 8. Petrogenesis of metapelitic and aluminous schist samples from both Makalu and Kangchenjunga profiles. Phase stability fields indicated
are of matrix assemblages in the metamorphic zones discussed in the text. Arrows indicate transition to new mineral parageneses as indicated
by mineral reaction textures (Tables 3 and 5). Stars are average P–T loci from the Barun Gneiss (Fig. 7; Table 8). Dashed lines indicate mineral
equilibria in calcareous rocks for reference (Fig. 9). Dots are invariant points. Mineral equilibria from the literature are as in Fig. 7; (8) Chatterjee
& Froese (1975); (10) Windom & Boettcher (1976) and Huckenholz et al. (1981); (11) Spear & Cheney (1989); (12) chlorite–muscovite stability
in KMASH (Seifert, 1970; Bird & Fawcett, 1973); (13) K-feldspar solvus in NaKASH (Morse, 1970); (14) theoretical relations in NaKASH
(Thompson, 1974).
appears contrary to the general notion that increasing P
favours garnet growth. This, however, can be rationalized
by comparison with the calculations of Vance & Mahar
(1998), made for similar bulk compositions to the samples
of this study. The low T and high P of formation of LHS
samples, and the typical heating and decompression paths
documented (P/T = –0·057 to –0·023), would result
in >5–10% increase in modal garnet, and thus garnet
overgrowths. Anorthite component in matrix plagioclase
increases towards rims (Table 7), indicating decompression coupled with garnet growth (e.g. St-Onge,
1987).
Phase stability fields for garnet + kyanite + staurolite
schist and epidote-bearing, garnet-free amphibolite
assemblages typical of Zones D and E define conditions
encompassing the THERMOCALC results from these
zones (Figs 7–9). Within the errors of the THERMOCALC calculations, all results are consistent with the
phase stability field of the matrix assemblage of the
respective sample. Consequently, these average P–T
results are considered plausible estimates of the equilibration conditions of the matrix assemblages in LHS
samples. Further, peak-T results display a spread broadly
consistent with an inverted metamorphic field gradient
(Fig. 2; Tables 3 and 8). However, because of structural
slicing (Fig. 2) and difficulty in assigning calcareous
samples to the metapelite metamorphic zones (Table 8),
in conjunction with the errors on P–T calculations, the
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GOSCOMBE AND HAND
P–T PATHS IN EASTERN HIMALAYA
Fig. 9. Petrogenesis of mafic, calc-silicate and carbonate samples from both Makalu and Kangchenjunga profiles. Phase stability fields indicated
are of calcareous matrix assemblages associated with the metapelitic metamorphic zones discussed in the text. Arrows indicate transition to new
mineral parageneses as indicated by mineral reaction textures (Tables 4 and 6). Stars are average P–T loci from the Barun Gneiss (Fig. 7; Table
8). Dashed lines indicate mineral equilibria in aluminous rocks for reference (Fig. 8). Mineral equilibria from the literature are as in Figs 7 and
8 and: (15) Binns (1968) and Moody et al. (1983); (16) experimental relations with QFM buffer (Liou, 1973); (17) greenschist–amphibolite facies
transition (Liou et al., 1974; Ghent et al., 1979); (18) Metz (1976) and Kase & Metz (1980); (19) Ellis (1978); (20) plagioclase-out in olivine tholeiite
(Green & Ringwood, 1967); (21) Valley & Essene (1980); (22) relations in CaMS (Yoder, 1976).
THERMOCALC results on their own are not sufficiently
accurate to define an inverted metamorphic field gradient.
GEOTHERMOBAROMETRY
Conventional geothermometers and geobarometers have
been used to constrain, in part, samples with insufficient
mineral end-members to calculate average P–T loci by
THERMOCALC (Appendix B). Calc-silicate sample
(M4) from the Himal Group is constrained by the
Fe-end-member garnet–hornblende–plagioclase geobarometer (Kohn & Spear, 1990) in conjunction with
THERMOCALC average T results. Similarly, garnet +
kyanite schist sample (M2) is constrained by the
Mg-end-member garnet–biotite–muscovite–plagioclase
geobarometer (Hoisch, 1991) and THERMOCALC average T results (Appendix A). A garnet + kyanite quartzofeldspathic gneiss sample (M95a) in Zone F is constrained
by the garnet–plagioclase–kyanite geobarometer
(Perchuk et al., 1985) and garnet–biotite geothermometer
of Dachs (1990), with activities calculated by method of
Hoinkes (1986) (Appendix B). Typical errors for these
methods of ±50°C and ±1 kbar are assumed (Dachs,
1990; Kohn & Spear, 1990). These geothermobarometeric results are very similar to THERMOCALC
results from other Himal Group samples and are
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DECEMBER 2000
Fig. 10. Compositional profiles across garnets in metapelitic schists and gneisses. (a) GHS samples; M39 and M58 from Barun Gneiss and M78
from Namche Migmatite Orthogneiss. (b) LHS samples; M109, M108 and M95A from the Himal Group, Zones C, E and F, respectively, and
K71a from the Seti Formation (Zone B). Horizontal scale is arbitrary.
consistent with the phase stability field of the matrix
assemblage (Fig. 8; Tables 3 and 4), and thus considered
reliable.
P–T EVOLUTION OF THE GREATER
HIMALAYAN SEQUENCES
Barun Gneiss and Namche Migmatite
Orthogneiss
Peak metamorphic conditions in the Barun Gneiss are
tightly constrained by both the stability field of matrix
assemblages and THERMOCALC results. Garnet +
cordierite + sillimanite assemblages in metapelites with
moderate XMg, such as the Barun Gneiss samples (Table
7), indicate an upper limiting P of 7·5 kbar (Powell &
Holland, 1990). Garnet + clinopyroxene amphibolite
assemblages indicate temperatures >800°C (Spear, 1981)
and pressures >5·5 kbar (Green & Ringwood, 1967) (Figs
8 and 9). Partial melting is pervasive and muscovite is
absent from metapelites and quartzo-feldspathic gneisses,
suggesting temperatures >685°C (Chatterjee & Froese,
1975; Harte & Hudson, 1979). These constraints are
consistent with peak metamorphic THERMOCALC
results averaging 837 ± 59°C and 6·7 ± 1·0 kbar (Table
8).
Sequential mineral parageneses in the Barun and
Jannu–Kangchenjunga Gneisses document an anticlockwise P–T loop restricted to within the sillimanite
field (Fig. 11a and b). Quartz + sillimanite + biotite
+ hercynitic spinel ± ilmenite inclusion assemblages
(Table 3) indicate pre-peak metamorphic conditions of
<4·0 kbar and >750°C (Fig. 8; Bohlen & Dollase, 1983;
Grant, 1985; Vielzeuf & Montel, 1994). Prograde spinel
was not stabilized at higher P by high f O2, because
prograde and matrix Fe–Ti oxides are typically ilmenite
± rutile pairs (Table 3), suggesting reducing conditions
(Powell & Sandiford, 1988). However, stabilization of
spinel inclusions at P >4·0 kbar by ZnO cannot be
discounted (Nichols et al., 1992). Because no other phase
in the assemblage partitions ZnO, the topology of spinel
phase relations is not significantly altered by increasing
ZnO contents (Fig. 11a and b) (Nichols et al., 1992; Hand
et al., 1994). Consequently, for any ZnO content, spinelbearing assemblages remain at lower pressures than the
1700
GOSCOMBE AND HAND
P–T PATHS IN EASTERN HIMALAYA
peak metamorphic garnet + sillimanite + cordierite
fields (Hand et al., 1994; Scrimgeour & Hand, 1997) (Fig.
11a and b). Thus the overgrowing of spinel + sillimanite
+ biotite inclusion assemblages by garnet and ultimately
peak-metamorphic cordierite + biotite + sillimanite +
garnet assemblages is consistent with an up-P prograde
path (Fig. 11b). Garnet overgrowths and coronas on
garnet are common (Tables 3 and 5). At relatively high
temperatures and low pressures, the increase in the
proportion of garnet is consistent with increasing P or
isobaric cooling (Vance & Mahar, 1998).
Immediately post peak-metamorphic reaction textures
indicate partial consumption of garnet by matrix cordierite, and subsequent partial replacement of cordierite
by sillimanite + biotite (Tables 3 and 5). Such a sequence
formed in response to either hydration or, as indicated
by P–T pseudosections derived for similar composition
metapelites (Scrimgeour & Hand, 1997), a low P/T
decompressive cooling path or decompression followed
by isobaric cooling (Fig. 11a). Both of these retrograde
P–T vectors could potentially be preceded by either a
clockwise or anticlockwise P–T loop and so do not
constrain the progade P–T vector (Fig. 11a).
Sillimanite inclusions in garnet are common, indicating
that alumino-silicate was not totally consumed during
prograde metamorphism, and because kyanite is entirely
absent from inclusion and matrix assemblages, the prograde path was logically restricted to the sillimanite field
(Fig. 11a). With no petrological or thermobarometry
evidence for prograde metamorphism at high P, and
early spinel parageneses and subsequent increases in the
proportion of garnet both suggesting an up-P prograde
path, an anticlockwise P–T loop is considered the most
plausible for the GHS, particularly in conjunction with
the isobaric retrograde path documented below (Fig. 11a
and b).
Peak metamorphic conditions were terminated by nearisobaric cooling. Matrix cordierite is replaced by sillimanite + gedrite-bearing assemblages and sillimanite
+ biotite assemblages (see above), indicating isobaric or
up-P cooling across the following reactions with low,
positive P/T (Hensen & Green, 1973; Spear &
Rumble, 1986; Baker et al., 1987; Xu et al., 1994) (Figs
8 and 11a and b):
cordierite + garnet → gedrite + sillimanite + quartz
K-feldspar + cordierite + garnet → sillimanite +
biotite.
Numerous post-peak metamorphic reaction textures
contain sillimanite (Tables 3 and 5). Kyanite is entirely
absent, suggesting little or no increase in P during cooling.
THERMOCALC results from matrix assemblages document a linear array from peak metamorphic conditions
through re-equilibrated samples ranging to 600 ± 45°C
Fig. 11. Phase relationships pertinent to Barun Gneiss and Jannu–
Kangchenjunga Gneiss metapelites. (a) Qualitative P–T pseudosection
in KFMASH with excess K-feldspar + quartz + melt, modified after
Scrimgeour & Hand (1997). Solid arrows constrained by sequential
mineral parageneses discussed in text; cordierite coronas on garnet
replaced by sillimanite + biotite aggregates. Shaded P–T paths indicate
alternative hypothetical prograde paths discussed in text: (i) constrained
by spinel + sillimanite + biotite inclusions; (ii) highest pressure P–T
path possible, constrained by common sillimanite and absence of
kyanite inclusions. (b) Qualitative P–T pseudosection in KFMASHT
+ Zn modified after Hand et al. (1994) is considered to account for
the possibility of Zn-bearing hercynite inclusions. Gedrite-forming
reaction in FMASH is after Spear & Rumble (1986) and Xu et al.
(1994). Solid arrow is P–T path constrained by the sequence of mineral
parageneses: from sillimanite + spinel + biotite inclusion assemblages,
overgrown by garnet and ultimately garnet + sillimanite + cordierite
+ biotite matrix assemblages, which are in turn overgrown by S3
gedrite + sillimanite assemblages.
and 5·7 ± 1·1 kbar, defining a near-isobaric cooling path
with P/T slope of 0·0035 kbar/°C (Fig. 7).
S3 foliation seams with sillimanite ± gedrite ± biotite
± haematite assemblages are common in the Barun and
Jannu–Kangchenjunga Gneiss units of the GHS, and are
also present in the highest-grade Zones (F and G) of the
underlying Himal Group. Garnet growth associated with
S3 is indicated by overgrowths on garnet porphyroblasts,
enveloping S3 with fine, aligned sillimanite and biotite
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JOURNAL OF PETROLOGY
VOLUME 41
inclusions preserved. Kyanite is absent from S3 assemblages, indicating that the Barun Gneiss was not in
the kyanite field during formation of the S3 foliation (Fig.
8). Peak, post-peak and S3 parageneses all formed in the
sillimanite field, indicating that the retrograde trajectory
of the Barun Gneiss did not pass through the kyanite
field, further confirming a near-isobaric cooling path
(Fig. 8). S3 assemblages indicate conditions of 610–750°C
and 4·5–7 kbar (Hensen & Green, 1973; Spear & Rumble,
1986; Baker et al., 1987) (Fig. 8). These P–T conditions
of S3 formation are compatible with THERMOCALC
results from re-equilibrated mafic and calc-silicate rocks
in the Barun Gneiss, which average 674 ± 33°C and
5·7 ± 1·1 kbar (Fig. 7; Table 8).
S3 foliations are interpreted to have formed during the
latest phase of ductile movement in the MCT, during
the Neohimalayan metamorphic event. Re-equilibration
of the mineral chemistry of matrix assemblages in calcsilicate and mafic rocks in the Barun Gneiss is interpreted
to be due to this later metamorphic event. S3 assemblages
are identical in both the GHS and upper LHS (Fig. 8),
suggesting both terranes were juxtaposed and at the same
crustal level during these latest ductile movements of the
MCT. In the GHS, S3 conditions must have been preceded by a near-isobaric cooling trajectory (see above),
and in the LHS by a near-isothermal decompression
path from peak metamorphic conditions in the kyanite
field (Fig. 8). Subsequent to S3, both metamorphic
terranes broadly evolved together along the same decompression and cooling path. Numerous retrogressive
chlorite-, epidote- and muscovite-forming reactions document continued cooling subsequent to formation of the
S3 fabric (Tables 3–6; Figs 8 and 9). Retrogressive kyanite
or andalusite has not been recognized and the exact
barometric response during retrogression subsequent to
S3 is unknown.
Black Gneisses
Coexisting K-feldspar and muscovite, sillimanite-bearing
assemblages (Table 3) and the presence of partial melt
segregations in the Black Gneisses are consistent with
conditions centred on 670–710°C and 4–7·5 kbar (Chatterjee & Froese, 1975; Fig. 8). Cordierite is documented
in the Black Gneisses (Pognante & Benna, 1993), limiting
pressures to <6 kbar (Powell & Holland, 1990; Fig. 8).
These are similar pressures, but lower temperatures, than
experienced in the Barun Gneiss (Fig. 8). As in the
underlying Barun Gneiss and Namche Migmatite Orthogneiss, kyanite is entirely absent from all Black
Gneisses parageneses. Further, reaction textures containing secondary sillimanite (Table 3) and late-stage
andalusite documented by Pognante & Benna (1993)
indicate a retrogressive P–T path with low, negative
NUMBER 12
DECEMBER 2000
P/T. Pervasive sillimanite needle inclusions within
muscovite and feldspars suggest that conditions immediately before the peak of metamorphism were in the
sillimanite field.
Although the prograde P–T path in the Black Gneisses
is largely unconstrained, these rocks are thought to have
experienced an anticlockwise P–T path similar to the
underlying Barun Gneiss, but nested within it at lower
temperatures (Fig. 12). This is because, collectively, the
entire section from Barun Gneiss to Black Gneisses is
a contiguous metamorphic terrane, with temperatures
decreasing upwards (i.e. not inverted isograds), that remained at moderate pressures throughout much of its
metamorphic history. Partial melts generated in the hightemperature basal units of this terrane (i.e. in the Barun
Gneiss) were emplaced as leucogranites within the Black
Gneisses at higher structural levels (Patiño Douce &
Harris, 1998; Harrison et al., 1999).
Jannu–Kangchenjunga Gneiss
Peak metamorphic conditions in the Jannu–
Kangchenjunga Gneiss are tightly constrained by the
stability field of matrix assemblages, these being similar
to the Barun Gneiss. Garnet + cordierite + sillimanite
metapelites indicate an upper limiting P of 7·5 kbar
(Powell & Holland, 1990) (Figs 8 and 9). Sphene in
mafic and calc-silicate gneisses indicate pressures >5 kbar
(Spear, 1981). Partial melting is pervasive, muscovite is
absent and K-feldspar is perthitic in metapelites and
quartzo-feldspathic gneisses, suggesting temperatures
>700°C (Chatterjee & Froese, 1975; Harte & Hudson,
1979). This is supported by meionite + plagioclase +
clinopyroxene + forsterite assemblages in calc-silicate,
indicating temperatures >750°C (Metz, 1976; Ellis, 1978;
Kase & Metz, 1980). Like the Barun Gneiss, Jannu–
Kangchenjunga Gneiss samples preserve sillimanite +
gedrite + biotite reaction textures replacing cordierite,
and all retrograde parageneses are restricted to the sillimanite field (Table 5). Consequently, the Jannu–
Kangchenjunga Gneiss is interpreted to have experienced
an isobaric cooling path similar to the Barun Gneiss.
Spinel inclusions have not been recognized and the
prograde P–T path cannot be constrained.
P–T EVOLUTION OF THE LESSER
HIMALAYAN SEQUENCES
Himal and Midland Groups
The LHS is a continuous Barrovian sequence that comprises the Himal Group and underlying Midland Group.
The petrology and structural evolution of the LHS is
very similar in both Makalu and Kangchenjunga profiles.
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GOSCOMBE AND HAND
P–T PATHS IN EASTERN HIMALAYA
Matrix assemblages span a wide range of high-P amphibolite facies conditions typical of the Barrovian series
and metamorphic isograds can be mapped (Figs 2 and
3). Metamorphic grade in both profiles increases towards
higher structural levels and thus the sequence preserves
an inverted metamorphic field gradient as is common in
the LHS throughout the length of the Himalayas (LeFort,
1975; Hodges et al., 1988; Pecher, 1989; Inger & Harris,
1992; Vannay & Hodges, 1996). In the Kangchenjunga
profile, the following metapelitic matrix assemblage zones
are recognized, from south to higher structural levels in
the north (Table 5):
Zone A: albite+chlorite+muscovite±biotite (Seti
Formation);
Zone B: albite+chlorite+muscovite+biotite+garnet
(Seti Formation);
Zone C: albite+muscovite+biotite+garnet±chlorite
(Ulleri Formation);
Zone D: oligoclase+muscovite+biotite+garnet+
kyanite (Kushma Formation);
Zone E: oligoclase+muscovite+biotite+garnet+
staurolite+kyanite;
Zone F: bytownite+biotite+garnet+K-feldspar+
kyanite±muscovite;
Zone G: bytownite+biotite+garnet+K-feldspar+
sillimanite+melt±kyanite.
Only Zones C–F are recognized in the Makalu profile
(Fig. 2). Estimates of peak-metamorphic conditions by
THERMOCALC from both profiles also confirm a rise
in T up-section (Fig. 7; Table 8). The average peakmetamorphic (peak-T) conditions, represented by rim
results from garnet + kyanite ± staurolite assemblages
in Zones D–F, are 609 ± 42°C and 8·8 ± 1·1 kbar
(Table 8). Peak-metamorphic conditions by geothermobarometry from a garnet + kyanite + K-feldspar
metapelitic gneiss in Zone F are 713 ± 50°C and 8·6
± 1·0 kbar (Table 8), and are entirely consistent with
the phase stability field of this sample. Zone G is represented only in the Kangchenjunga profile, with sillimanite ± kyanite + garnet + K-feldspar + melt
+ biotite and muscovite-absent assemblages, suggesting
temperatures >700°C (Thompson, 1974; Chatterjee &
Froese, 1975) and pressures (8 kbar at the kyanite–
sillimanite transition (Fig. 8). The absence of perthitic
K-feldspar sets a maximum limiting T of 770°C and
absence of cordierite suggests P >6·5 kbar (Fig. 8). These
conditions are supported by calc-silicate assemblages (Fig.
9).
At lower structural levels, in the Midlands Group, an
inverted metamorphic sequence from Zone A to Zone
C is preserved, which is contiguous with and progresses up
into the inverted metamorphic sequence in the overlying
Himal Group (Figs 3 and 8). Matrix assemblages document an increase in metamorphic grade from >475°C
above the biotite isograd in Zone A, across the garnet
isograd into Zone B and to >550°C (Fig. 8) in Zone D.
The only barometric constraints available for Zones A–C
are the presence of kyanite and absence of sillimanite
(Table 5), indicating pressures >5 kbar (Fig. 8).
In contrast to the GHS, the Himal Group contains
matrix kyanite and, in the highest-grade zones, kyanite
inclusions, thus attesting to prograde metamorphism entirely within the kyanite field. The peak of metamorphism
is interpreted to be post-tectonic (post-S1) and coeval
throughout all zones in the LHS. In Zones A and B,
snowball garnet growth is syn-tectonic (Powell & Vernon,
1979) and garnet is also deformed within the S1 foliation
(Fig. 6d). Biotite porphyroblasts (Zone A) and idiomorphic garnet overgrowths in all zones (Fig. 6c) formed
at the peak of metamorphism subsequent to S1. A portion
of the prograde P–T path is preserved by garnet growth
zoning, with core and rim P–T calculations indicating
decompression accompanying heating (Fig. 7). Peak pressures experienced were at least 10 kbar as indicated by
THERMOCALC results from mineral cores, implying
considerable burial of the LHS and crustal thickening
during the Himalayan Metamorphic Cycle. Peak-metamorphic conditions indicated by rim THERMOCALC
results were at slightly lower pressures, with most samples
between 7·4 and 8·8 kbar.
Sequential mineral parageneses and reaction textures
constrain portions of what are interpreted to be clockwise
P–T paths in LHS zones (Fig 8 and 9). Decompression
accompanying peak metamorphism is constrained to be
wholly within the kyanite field in Zones A–E and to have
passed into the sillimanite field only in the higher-grade
Zones F and G (Fig. 8). In Zone G, matrix assemblages
equilibrated near the kyanite–sillimanite transition (Table
5). In these samples, the near absence of matrix kyanite
and presence of kyanite inclusions in garnet imply prograde metamorphism from the kyanite field into the
sillimanite field at the peak of metamorphism (Fig. 8).
Furthermore, rare aligned sillimanite inclusions in garnet
rims are due to continued garnet growth, within the
sillimanite field, at the peak of metamorphism. Garnet
is also consumed by metamorphic reactions and even
entirely enclosed by coronas of matrix plagioclase. Thus
most garnet growth occurred early, during prograde
metamorphism. This is also evident in the lower-grade
zones, where garnet porphyroblasts are enclosed by the
S1 foliation, and idiomorphic (peak-metamorphic) garnet
overgrowths are relatively thin and rare (Tables 3 and
5).
Late-stage S3 foliation seams with sillimanite ± gedrite
± biotite assemblages are developed in metapelites in
Zones F and G (Table 5) and formed at conditions of
610–750°C and 4·5–7 kbar, identical to S3 in the GHS.
The GHS and all LHS zones are interpreted to have
evolved along a broadly similar retrogressive P–T path
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JOURNAL OF PETROLOGY
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subsequent to S3. Numerous retrograde biotite-, muscovite- and chlorite-forming reaction textures are recorded in all LHS zones and the GHS, documenting
cooling to at least 400°C (Figs 8 and 9; Tables 3–6). No
aluminosilicate phases were formed in post-S3 retrograde
parageneses, thus the barometric response during cooling
is unconstrained.
NUMBER 12
DECEMBER 2000
km. The differing style of metamorphism of the upperplate GHS, with respect to the lower-plate LHS, is further
supported by the contrasting interpretative P–T paths in
these two metamorphic terranes (Fig. 8). Four tectonic
models are discussed below, and their compatibility with
observations of the structural and metamorphic evolution
of Eastern Nepal is evaluated.
Model A: inverted metamorphic sequence
Biotite Gneiss (Main Central Thrust)
Biotite Gneiss is a highly sheared unit of biotite-rich
schistose gneisses representing the principal movement
zone of the MCT, at the boundary between the Himal
Group and the Jannu–Kangchenjunga Gneiss. Assemblages contain K-feldspar and are muscovite free,
and partial melt segregations and granitic veins are
common, indicating temperatures >670°C (Chatterjee &
Froese, 1967; Thompson, 1974). K-feldspar is not perthitic, limiting temperatures to <750°C (Morse, 1970).
Sillimanite is common, gedrite rare and kyanite and
cordierite are entirely absent, restricting pressures to
5–7·5 kbar (Hensen & Green, 1973; Baker et al., 1987;
Powell & Holland, 1990). These conditions are similar
to those in the uppermost Himal Group (Zone G), except
for the complete absence of kyanite, suggesting slightly
lower P in the Biotite Gneiss unit. Sillimanite is found
in inclusion, peak and retrograde parageneses, suggesting
these rocks remained in the sillimanite field throughout
shearing in the MCT.
These assemblages and conditions of formation are
very similar to S3 parageneses in the underlying Himal
Group and throughout the GHS. Consequently, the S3
foliation is correlated with the movement phase in the
MCT represented by Biotite Gneiss assemblages. S3 metamorphic conditions are correlated with the moderate-T/
moderate-P Neohimalayan metamorphic event (Hodges
et al., 1988; Pecher, 1989; Vannay & Hodges, 1996).
Although the MCT may have experienced multiple
movement phases, complete recrystallization and a single
pervasive foliation suggest that the Biotite Gneiss assemblages represent the main, and latest, phase of movement. The sense of movement during this period is
unknown from the region investigated.
DISCUSSION
There is no petrological evidence for the GHS having
ever been buried deeper than the 6·5 kbar experienced
at the peak of metamorphism. Furthermore, peak metamorphism was on a high average geothermal gradient
of 36 ± 4°C/km. This high-T/moderate-P metamorphism is incompatible with the high-P/moderate-T
Barrovian metamorphism in the underlying LHS, with
average geothermal gradients centred on >20 ± 2°C/
An inverted metamorphic sequence is documented from
throughout the entire LHS, up to the base of the GHS.
This feature is well recognized along the entire length of
the Central Himalayan metamorphic front (e.g. LeFort,
1975; Hodges et al., 1988; Pecher, 1989; Inger & Harris,
1992; Vannay & Hodges, 1996). Conventional models
of Himalayan metamorphism also include the GHS as
the uppermost portion of this inverted metamorphic
sequence. However, in the Eastern Nepal region investigated, metamorphism of the GHS was due to an
entirely distinct average geothermal gradient and the
GHS experienced a contrasting P–T path with respect to
the underlying LHS. Thus the GHS cannot be considered
continuous with, nor part of, the inverted metamorphic
sequence displayed in the LHS. In the GHS, peak-T
conditions also record the maximum depth of burial
attained (Fig. 8). Peak metamorphic conditions in the
GHS are interpreted to have been attained at the same
time as peak-P conditions, before the peak of metamorphism, in the LHS (Fig. 12; Table 2). This is because
matrix assemblages formed at the same time in both
metamorphic terranes, in association with south-directed
ductile overthrusting. These matrix assemblages crystallized at peak-P conditions (before peak-T) in the LHS
and at peak-T conditions in the GHS.
There are numerous models to explain the inverted
metamorphic sequence in the Himalayas. These have
been discussed in detail by LeFort (1975), Mohan et al.
(1989), Searle & Rex (1989) and Hubbard (1996) and
fall into two classes: (1) overthrusting of a hot slab, or (2)
syn- to post-metamorphic overturning of isograds by
ductile overthrusting (Mohan et al., 1989). These are
not mutually exclusive and both processes are possibly
involved in Himalayan metamorphism. Because metamorphic isograds are parallel to both the main metamorphic foliation and the MCT and associated thrusts,
most models propose peak metamorphism synchronous
with overthrusting in the MCT Zone. In the LeFort
(1975) model, lower-crustal, high-grade metamorphic
zones are thrust over the upper-crustal lower-grade zones
at a rate faster than the relaxation of isotherms and the
resultant inverted sequence preserved by rapid uplift.
The anticlockwise P–T path in the GHS can potentially
be rationalized as being the overthrust upper-plate in a
modified version of this model (see Model D below).
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P–T PATHS IN EASTERN HIMALAYA
However, the distinctly higher average geothermal gradient of this metamorphic terrane needs to be explained.
Model B: partially reworked pre-Himalayan parageneses
Contrasting metamorphic styles of the GHS and LHS
can potentially be resolved by a hypothetical scenario in
which the GHS preserve pre-Himalayan metamorphic
assemblages. For example, the high-T/moderate-P metamorphism of the GHS may have occurred during a
Proterozoic metamorphic event in the Indian plate before
collision and was juxtaposed against the LHS during the
Himalayan Metamorphic Cycle. Such a scenario requires
minimal reworking of the older assemblages during the
Himalayan Metamorphic Cycle. Collisional belts
incorporating basement terranes preserving older
assemblages,
which
experienced
only
partial
re-equilibration with minimal recrystallization of new
assemblages, have been recognized (St-Onge & Ijewliw,
1996; Goscombe et al., 1998) and support the plausibility
of such a model.
Barnicoat & Treloar (1989) reported that relic preHimalayan tectonometamorphic and magmatic events
are extensive in the GHS. However, there is no geochronological evidence in the Nepal Himalayas for Proterozoic metamorphic mineral assemblages. Nor is there
conclusive structural evidence for reworked older fabrics
in the GHS of the area investigated. Matrix mineral
assemblages formed in association with the main S1–L1
fabric in both metamorphic terranes. This metamorphic
fabric was coeval and tectonically equivalent in both
terranes and associated with ductile overthrusting during
the Himalayan Metamorphic Cycle.
Similarly, S3 is precluded from being interpreted as the
only expression of the Himalayan Metamorphic Cycle
reworking older parageneses in the GHS, because identical and similarly late-formed S3 parageneses are also
developed in the LHS. Further, in these hypothetical
scenarios, the GHS would not necessarily be a hot slab
during the Himalayan Metamorphic Cycle, rendering
development of an inverted metamorphic sequence in
the LHS unlikely.
Model C: polyphase Himalayan Metamorphic Cycle
A similar hypothetical scenario to Model B can be
proposed in which the Himalayan Metamorphic Cycle
involved two or more metamorphic events, which are
variably preserved in different thrust-bound units in the
orogen. For example, high-T/moderate-P parageneses
in the GHS may have formed in an earlier event in the
Himalayan Metamorphic Cycle and were insignificantly
reworked by a later event in which the LHS was metamorphosed. Such a scenario is not considered a plausible
explanation for the contrasting metamorphic styles recognized, for the same reasons as discussed for Model
B. That is, the main metamorphic fabrics in both metamorphic terranes are considered essentially coeval and
associated with the same south-directed ductile overthrusting.
Two metamorphic events were recognized in the Himalayan Metamorphic Cycle by previous workers from
throughout the Himalayas (Hodges et al., 1988; Pecher,
1989; Inger & Harris, 1992; Vannay & Hodges, 1996).
The Eohimalayan event involved prograde Barrovian
metamorphism in response to overthrusting and burial
during collision in the Middle Eocene to Oligocene. The
Early–Middle Miocene Neohimalayan event is characterized by moderate-T/moderate-P conditions in the
sillimanite field. Together these metamorphic events comprise only portions (crystallization events) of the one
metamorphic cycle (Passchier & Trouw, 1996), with a
clockwise P–T path. High-P/moderate-T parageneses do
not occur in the GHS. Thus metamorphism of the GHS
cannot be simply resolved within the above scheme
without modification (see Model D), because the prograde
path was presumably of high-T/low-P type.
Despite the inability of this model to account for the
style of metamorphism in the GHS, both metamorphic
events are recognized in the higher-grade zones of the
LHS. Matrix mineral assemblages formed at peak-P
conditions are of high-P/moderate-T type and consistent
with the Eohimalayan event. Sillimanite-bearing S3 assemblages formed during decompression through the
peak of metamorphism and constitute the Neohimalayan
event. Similarly, two crystallization events are recognized
in the GHS; peak metamorphic matrix assemblages associated with S1–L1 fabrics and a later event represented
by S3 assemblages. Conditions of formation of S3 assemblages are the same in both the GHS and LHS, and
correlated with the Neohimalayan event. Thus a twophase Himalayan Metamorphic Cycle is recognized in
both metamorphic terranes, although currently accepted
metamorphic conditions of both events are consistent
with metamorphism of the LHS only. However, the
prograde (Eohimalayan) phase is of contrasting metamorphic style in the two metamorphic terranes and
metamorphism of the GHS in eastern Nepal is of a type
unrecognized elsewhere in the GHS of the Himalaya.
New geochronology suggests Barrovian metamorphism
in the LHS may be as young as Late Miocene–Pliocene
(Harrison et al., 1997), lending support to a polymetamorphic model. However, such a scenario requires
an explanation for the absence of pervasive reworking
of the overlying GHS in this younger event. It is plausible
that the GHS slab may have been minimally structurally
reworked in a Late Miocene–Pliocene event, with reworking being restricted to development of the S3 foliation
and re-equilibration of calcareous rock-types without
significant recrystallization. In such a scenario, the S3
foliation developed in the uppermost LHS must have
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formed immediately after and be essentially coeval with
the proposed Late Miocene–Pliocene main metamorphic
foliation in the LHS.
Model D: Paired Metamorphic Mountain Belt (preferred
model)
Most aspects of the juxtaposed, but contrasting, eastern
Nepal metamorphic terranes can be resolved as a Paired
Metamorphic Mountain Belt (PMMB) model analogous
to the Acadian metamorphism in the Appalachians (Armstrong et al., 1992). The bases of this model are as follows:
(1) different P–T paths are experienced at different
crustal levels in a collisional orogen, particularly with
respect to being upper-plate vs lower-plate components
in a crustal-scale overthrusting system (Chamberlain &
Karabinos, 1987; Shi & Wang, 1987; Armstrong et al.,
1992).
(2) Metamorphism is dynamic and intimately associated with progressive crustal-scale overthrusting, resulting in an inverted metamorphic sequence (Hubbard,
1996).
(3) Different average geothermal gradients are generated by different tectonic processes within the respective
metamorphic terranes. These are suggested to be (a)
burial of an anomalously radiogenic GHS generating
high-T/moderate-P metamorphism as modelled by Sandiford & Hand (1998) and previously suggested for the
Himalayas by Pinet & Jaupart (1987), and (b) burial and
crustal overthickening (England & Thompson, 1984) in
the LHS.
A possible model for the tectonometamorphic evolution
of the eastern Nepal region investigated, combining these
elements, is outlined below (Fig. 12).
The GHS was buried to moderate crustal levels by
low-angle overthrusting of the Tethyan Sequences from
the north, along a proposed Eohimalayan Thrust (Vannay & Hodges, 1996). Prograde metamorphism of the
GHS accompanied crustal thickening during collisional
orogenesis, with peak-T conditions being reached at the
maximum P experienced (Fig. 12a). The high average
geothermal gradient of GHS metamorphism cannot be
a conductive response to crustal thickening only (Sandiford & Powell, 1991). A possible explanation for the
high-T/moderate-P metamorphism is an internally derived heat source resulting from a particularly radiogenic
GHS (Sandiford & Hand, 1998). A high content of
radiogenic elements in the GHS has previously been
suggested as responsible for the high average geothermal
gradient in this terrane (Pinet & Jaupart, 1987). The
numerous leucogranite sills and plutons in the uppermost
GHS are a consequence of the high average geothermal
gradient and not the cause, as they are derived by partial
melting from within the GHS (Patiño Douce & Harris,
1998). Furthermore, most leucogranite bodies are in the
NUMBER 12
DECEMBER 2000
highest structural levels and lowest-grade portion of the
GHS. Nelson et al. (1996) have reported deep seismic
evidence for a zone of partial melting and granite melt
pooling below the STDS, in the mid-crust, 100–200 km
north of the currently exposed GHS. This indicates that
high-grade metamorphism and partial melting in the
GHS is a continuing process in which these hot rocks
are being continually fed south like a conveyor belt as a
result of overthrusting at the MCT, resulting in burial
metamorphism and isograd inversion of the lower-plate
LHS rocks.
The GHS progressively overrode the lower-plate LHS,
giving rise to high-P/moderate-T Barrovian metamorphism, with the burial phase of the LHS occurring
at the same time as the peak of metamorphism in
the GHS (Fig. 12b). Prograde metamorphism in both
metamorphic terranes constitutes the Eohimalayan event
and coincides with ductile, north over south thrusting
and development of the pervasive main metamorphic
fabric (S1–L1) in both terranes. This prograde (burial)
phase culminated in peak-T conditions in the upperplate GHS (Fig. 12a) and peak-P conditions in the lowerplate LHS (Fig. 12b), with coarse-grained S1–L1 matrix
assemblages crystallized at these respective conditions
(Table 2). Progressive southward transport of the upperplate GHS accompanied cooling of this terrane while it
remained at essentially the same relative crustal level,
and concomitant progressive southward transport in the
LHS resulted in inversion of the metamorphic isograds,
this being further enhanced by the overriding hot GHS
slab. Prograde metamorphism of the LHS was terminated
by rapid decompression through the peak of metamorphism and in the highest-grade zones, progressed
into the sillimanite field (Fig. 12c).
It would be predicted by this PMMB model that peak
metamorphic conditions in the GHS should slightly predate peak metamorphic (peak-T) conditions in the underlying LHS (Fig. 12b and c). This indeed seems to be
the case, with peak metamorphism in the GHS being
generally accepted as occurring at >22 Ma (Guillot et
al., 1994; Harrison et al., 1995; Coleman & Hodges,
1998) and associated with emplacement of leucogranites
in the High Himal with ages in the range 26–17 Ma
(Scharer, 1984; Deniel et al., 1987; Hubbard & Harrison,
1989; Copeland et al., 1990; Inger & Harris, 1992;
Metcalfe, 1993; Hodges et al., 1996; Edwards & Harrison,
1997; Coleman, 1998). Peak metamorphic conditions in
the LHS are associated with overthrusting at the MCT
and variously interpreted to be at 22–16 Ma (Hodges et
al., 1996; Coleman, 1998; Coleman & Hodges, 1998),
before cooling ages of 15–13 Ma (Macfarlane et al., 1992;
Vannay & Hodges 1996) and possibly as late as 13–6 Ma
(Macfarlane et al., 1992; Harrison et al., 1997). Postpeak, S3 foliations with identical assemblages in both
1706
GOSCOMBE AND HAND
P–T PATHS IN EASTERN HIMALAYA
Fig. 12. Interpretative schematic representation of the preferred paired metamorphic belt model for the tectonometamorphic evolution of the
Eastern Nepal region investigated. Age range of Eohimalayan and Neohimalayan metamorphic events and leucogranites as reported in the
literature (see text). Dashed line is >600°C geotherm at the end of each indicated period. Dots follow the evolution of currently exposed and
sampled metamorphic zones. Double arrows indicate pervasive ductile shearing and S1 foliation formation, and single arrow indicates movement
on thrusts or extensional detachment faults. Highlighted portions of the interpretative P–T paths are for the period indicated; Β, average P–T
loci at specific metamorphic periods in the different terranes (Fig. 7).
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JOURNAL OF PETROLOGY
VOLUME 41
metamorphic terranes formed in association with shearing in the MCT (Figs 8 and 12c), implying that the
terranes were juxtaposed and occupying approximately
the same crustal level (>21 km depth; Fig. 8) at this
time. These moderate-T/moderate-P parageneses constitute the Neohimalayan event (Pecher, 1989; Vannay &
Hodges, 1996), of >22–13 Ma age (Hubbard & Harrison,
1989; Inger & Harris, 1992; Metcalfe, 1993; Table 2).
Neohimalayan metamorphic conditions were immediately preceded by decompression from peak-metamorphic conditions in the LHS and by essentially isobaric
cooling in the GHS (Fig. 8).
Rapid decompression and extrusion of the entire metamorphic core of the Himalayas, contemporaneous with
contraction (Fig. 12c and d), has been proposed by Burg
et al. (1984), England & Molnar (1993), Hodges et al.
(1993, 1996) and Vannay & Hodges (1996). Extrusion
of deep-seated rocks is accommodated by extensional
detachments, during episodic periods of orogenic collapse
in response to lithospheric overthickening (Hodges et al.,
1996). The contrasting P–T trajectories, immediately
before S3, in the upper- and lower-plate terranes in
the area investigated could plausibly be explained by
reactivation of the MCT as an extensional detachment
(Fig. 12c). However, extensional movements during the
Miocene were accommodated by the STDS above the
GHS (Burchfiel et al., 1992; Hodges et al., 1996; Vannay
& Hodges, 1996; Edwards & Harrison, 1997), and extensional reactivation of the MCT, at the base of the
GHS, has not been documented.
Although a PMMB model may explain much of the
distribution of palaeo-metamorphic conditions in the
investigated portion of eastern Nepal, it remains a working hypothesis. In particular, the sequence of thrusting
and extensional movements on major structures in the
region needs to be established and the proposed mechanism for the high average geothermal gradient in the
GHS has not been tested. Furthermore, comparably high
average geothermal gradients and anticlockwise P–T
paths, such as proposed for the portion of the GHS
investigated, have not been reported from elsewhere in
the GHS. Consequently, the model discussed here may
not be generally applicable throughout the entire Himalayan metamorphic front. Studies of deeply eroded,
older metamorphic belts elsewhere have shown them
to be heterogeneous, being composed of metamorphic
terranes of contrasting metamorphic style [i.e. Appalachians (Armstrong et al., 1992), Zambezi Belt (Goscombe et al., 1998), Damara Orogen (Miller, 1983),
Ubendian Belt (Ring, 1993), Grenville Province (Rivers
et al., 1989) and Ungava Orogen (St-Onge & Ijewliw,
1996)], suggesting that it may be unlikely to expect
comparable metamorphic profiles throughout the entire
1600 km length of the Central Himalaya metamorphic
front.
NUMBER 12
DECEMBER 2000
ACKNOWLEDGEMENTS
Thanks are due to Mel Lambourne for the inspiration,
and to Ned Stolz, Wolf Zwolsman, Lia, Prem, Kancha
Lama, Ram Lama, Rhim Bardo Lama, Pemba Dike
Sherpani and Shiring Nema Sherpani for their great
company and help amongst those bonnie hills. Constructive comments by Drs K. Hodges, T. Hoisch, C.
M. Wilson, S. Harley and S. Sorensen, and discussions
with Drs R. Oliver and M. Sandiford, greatly improved
this paper, and their efforts were very much appreciated.
Dr Wieslaw Jablonski at the University of Tasmania and
Dick Rickards at the University of Cape Town are
thanked for their help with electron microprobe analyses.
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1711
T (°C)
P (kbar)
766±57
M78-rim
789±56
M51B-rim
5·7±0·7
6·5±0·8
6·5±1·1
6·8±1·0
6·5±0·9
7·0±0·6
5·1±2·0
6·5±1·7
f
1·08
1·14
1·66
1·56
1·37
0·88
1·58
1·32
1712
635±47
600±45
M87-core
M87-rim
561±11
587±14
584±13
596±14
592±13
596±13
521±69
537±73
K71a–overgrowth
K65c-core
K65c-rim
K62b-rim
K61b-core
K61b-rim
K60c-core
K60c-rim
8·8±1·5
9·1±1·5
9·7±1·1
9·3±1·1
9·6±1·1
10·1±1·1
10·3±1·1
7·4±1·9
8·0±1·0
9·7±0·8
10·7±1·1
0·804
0·795
0·068
0·082
0·104
0·103
0·177
−0·105
0·80
0·36
0·84
0·73
0·47
0·86
0·39
0·71
0·73
1·18
0·89
0·90
0·86
0·83
0·93
0·89
0·97
0·55
0·85
0·087
0·192
0·040
0·018
0·556
0·654
0·652
0·561
0·542
0·270
0·290
alm–py–gr–phl–ann–east–naph–mu–an–ab–q–ky–H2O
alm–py–gr–phl–ann–east–naph–mu–an–ab–q–ky–H2O
alm–py–gr–phl–ann–east–naph–mu–an–ab–clin–daph–ames–q–ky–H2O
alm–py–gr–phl–ann–east–naph–mu–an–ab–clin–daph–ames–q–ky–H2O
alm–py–gr–phl–ann–east–naph–mu–an–ab–clin–daph–ames–q–H2O
alm–py–gr–phl–ann–east–naph–mu–an–ab–clin–daph–ames–q–H2O
alm–py–gr–phl–ann–east–naph–mu–an–ab–clin–daph–q–H2O
alm–py–phl–ann–east–naph–mu–cel–pa–clin–daph–ames–q–H2O
alm–gr–phl–ann–mu–cel–mst–fst–an–ky–ames–q–H2O
alm–gr–east–mu–cel–fst–an–ky–ames–q–H2O
alm–gr–phl–ann–mu–cel–mst–fst–an–ky–ames–q–H2O
alm–gr–phl–ann–mu–mst–fst–an–ky–ames–q–H2O
alm–an–ab–di–hed–cats–hb–parg–ed–me–q–H2O–CO2
alm–py–gr–an–ab–parg–ed–di–hed–cats–q–H2O
alm–py–gr–an–ab–parg–ed–di–hed–cats–q–H2O
alm–py–gr–an–ab–parg–ed–anth–di–hed–cats–q–H2O
alm–py–gr–an–ab–parg–ed–anth–di–hed–cats–q–H2O
alm–py–gr–phl–ann–an–ab–hb–parg–ed–di–cats–q–H2O
alm–py–gr–phl–ann–an–ab–hb–parg–di–cats–q–H2O
alm–py–crd–fcrd–phl–ann–naph–an–ksp–ab–sill–q–H2O
alm–py–crd–fcrd–phl–ann–east–naph–ksp–ab–sill–q–H2O
alm–py–gr–crd–fcrd–phl–ann–an–ksp–sill–q–H2O
alm–py–crd–fcrd–phl–ann–ksp–sill–q–H2O
alm–py–gr–crd–fcrd–sill–ksp–an–ab–phl–ann–naph–east–q–H2O
alm–py–crd–fcrd–sill–ksp–an–ab–phl–ann–naph–q–H2O
alm–py–gr–sill–phl–ann–mu–cel–an–ksp–q–H2O
alm–py–gr–sill–phl–ann–mu–cel–an–ksp–q–H2O
Mineral end-members used
NUMBER 12
Himal Group, Kangenjunga profile
662±21
M108-rim
8·4±1·0
650±20
556±20
646±21
M106-core
M106-rim
506±31
M4-core∗
M108-core
9·7±1·1
570±37
9·8±0·9
553±36
M1-rim
9·4±0·9
5·7±1·1
6·0±1·1
5·9±1·3
5·4±1·2
0·853
0·869
0·869
0·833
0·882
0·848
0·920
0·950
cor
VOLUME 41
M1-core
Himal Group, Makalu profile
645±43
664±37
M84-rim
M84-core
Barun Gneiss samples that have been re-equilibrated, Makalu profile
806±84
823±60
832±76
M39-core
M51B-core
821±72
M58-rim
M39-rim
877±55
M58-core
Barun Gneiss, Makalu profile
818±47
M78-core
Namche Migmatite Orthogneiss, Makalu profile
Sample
Table A: Pressure–temperature conditions, derived using THERMOCALC v2.0b (Powell & Holland, 1988)
APPENDIX A
JOURNAL OF PETROLOGY
DECEMBER 2000
400°C
5·8±1·02
Average T results (°C)
Sample
f
7·0 kbar
8·3±1·87
8·0±1·84
6·8±1·17
500°C
1713
683±31
568±34
M57-core∗
M57-rim∗
643±95
637±94
471±30
492±30
M2-core
M2-rim
M109-core
M109-rim
561±12
K71a-over
1·1
1·1
0·4
0·5
0·6
0·5
1·7
1·6
1·5
1·1
562±12
543±11
492±32
470±31
620±95
625±95
537±56
614±32
726±30
720±18
0·6
0·7
0·5
0·4
0·3
0·2
1·7
1·4
1·4
1·1
f
0·5
0·8
0·4
f
560±12
540±12
489±32
467±32
604±95
608±95
543±60
645±33
760±29
752±17
9·0 kbar
8·6±2·07
8·3±2·03
7·9±1·32
600°C
0·4
0·7
0·7
0·2
0·2
0·3
1·8
1·4
1·3
1·0
f
0·3
0·6
0·1
f
alm–phl–ann–east–naph–mu–cel–pa–daph–ames–q–H2O
alm–phl–ann–east–naph–mu–cel–pa–daph–ames–q–H2O
alm–py–phl–ann–east–ky–ames–q–H2O
alm–py–phl–ann–east–ky–ames–q–H2O
alm–py–phl–ann–naph–mu–pa–ky–q–H2O
alm–py–phl–ann–naph–mu–pa–ky–q–H2O
alm–py–an–hb–di–hed–cats–me–q–H2O
hb–ftr–anth–di–hed–cats–ann–naph–an–ab–ksp–cc–q–H2O–CO2
hb–ftr–anth–di–hed–ann–naph–an–ab–ksp–cc–q–H2O–CO2
hb–ftr–anth–di–hed–phl–ann–naph–an–ab–ksp–cc–q–H2O–CO2
Mineral end–members used
alm–phl–ann–east–naph–mu–cel–pa–daph–ames–q–H2O
alm–phl–ann–east–naph–mu–cel–pa–daph–ames–q–H2O
alm–gr–ann–an–mu–ky–q–H2O
Mineral end-members used
f, 2 test; cor, correlation coefficient. Mineral end-member abbreviations after Powell & Holland (1988).
∗Calculated for XCO2 = 0·5 and XH2O = 0·5.
542±12
K71a-core
Himal Group, Kangchenjunga profile
522±54
M4-rim∗
Himal Group, Makalu profile
680±18
M42-core∗
Barun Gneiss samples that have been re-equilibrated, Makalu profile
5·0 kbar
7·9±1·67
0·9
7·7±1·76
K71a–over
1·1
0·7
K71a-core
Himal Group, Kangchenjunga profile
M2-rim
f
Average P results (kbar)
Himal Group, Makalu profile
Sample
GOSCOMBE AND HAND
P–T PATHS IN EASTERN HIMALAYA
Rock-type
1714
calc-silicate
gn–ky schist
gn-ky gneiss
gn–ky gneiss
gn–ky gneiss
gn–ky gneiss
M4-rim
M2-core
M95a-core
M95a-rim
M95a-core
M95a-rim
Himal Group, Makalu profile
Sample
641±50
701±50
650±50
710±50
2·7±1
2·3±1
4·2±1
4·6±1
8·8±1
8·3±1
500°C
6·1±1
6·5±1
9·9±1
9·3±1
600°C
gn–pl–ky (Perchuk et al., 1995)
gn–bi (Dach, 1990; activity, Hoinkes, 1986)
gn–bi–mu–pl–q (Hoisch, 1991; Mg)
gn–hn–pl (Kohn & Spear, 1990; Fe#B)
Calibration
VOLUME 41
632±50
692±50
7·7±1
7·3±1
400°C
8·0 kbar
4·0 kbar
6·0 kbar
P result (kbar)
T result (°C)
Table B: Geothermobarometry results from the Makalu profile
APPENDIX B
JOURNAL OF PETROLOGY
NUMBER 12
DECEMBER 2000
1715
0·72
X(Fe)
0·00
0·00
8·02
OH
Sum
0·00
0·00
0·00
Na
K
0·87
8·02
0·00
0·10
0·37
0·17
0·81
0·10
Mn
2·40
0·00
0·09
2·06
Fe2+
Mg
0·01
Fe3+
2·00
0·00
0·00
2·98
99·46
0·00
0·00
0·03
1·12
3·06
2·45
35·33
0·03
20·83
Ca
0·00
1·98
Cr
Al
2·97
98·68
Total
0·00
0·00
H 2O
Si
0·02
K 2O
Ti
0·02
1·07
CaO
Na2O
1·42
6·83
MnO
MgO
0·25
30·79
21·04
Al2O3
Fe2O3
0·05
Cr2O3
FeO
0·01
0·00
0·00
36·60
37·19
SiO2
rim
core
TiO2
M58
M58
Sample:
gn
gn
Mineral:
0·81
8·01
0·00
0·00
0·00
0·73
0·42
0·07
1·84
0·04
1·95
0·00
0·00
2·96
99·16
0·00
0·00
0·02
8·51
3·55
1·07
27·48
0·70
20·71
0·05
0·05
37·02
core
M87
gn
0·84
8·04
0·00
0·00
0·00
0·68
0·36
0·12
1·91
0·05
1·95
0·00
0·00
2·97
99·39
0·00
0·01
0·00
7·86
2·99
1·75
28·41
0·82
20·58
0·01
0·00
36·95
rim
M87
gn
0·88
8·04
0·00
0·00
0·01
0·19
0·34
0·01
2·52
0·01
1·99
0·00
0·00
2·97
99·23
0·00
0·01
0·03
2·12
2·76
0·17
36·85
0·23
20·63
0·00
0·04
36·39
core
M106
gn
0·89
8·03
0·00
0·00
0·00
0·19
0·32
0·05
2·50
0·00
2·01
0·00
0·00
2·96
99·08
0·00
0·00
0·03
2·17
2·60
0·72
36·52
0·00
20·79
0·03
0·05
36·17
rim
M106
gn
0·78
8·05
0·00
0·00
0·00
1·29
0·33
0·29
1·18
0·09
1·91
0·00
0·00
2·96
101·21
0·00
0·00
0·00
15·47
2·84
4·34
18·09
1·49
20·83
0·05
0·07
38·02
core
M1
gn
Table C: Representative matrix mineral analyses from the Makalu profile
APPENDIX C
0·79
8·05
0·00
0·00
0·00
1·31
0·32
0·30
1·17
0·08
1·92
0·00
0·00
2·95
101·10
0·00
0·00
0·01
15·71
2·72
4·50
17·93
1·43
20·87
0·00
0·01
37·93
rim
M1
gn
0·84
33·43
4·00
0·01
0·01
0·00
0·62
0·02
3·37
0·04
17·60
0·00
0·11
7·65
98·04
2·11
0·02
0·02
0·00
1·47
0·10
14·17
0·20
52·52
0·00
0·50
26·92
core
M106
st
st
0·85
33·38
4·00
0·01
0·00
0·00
0·60
0·02
3·36
0·04
17·48
0·01
0·12
7·74
98·30
2·12
0·02
0·01
0·00
1·43
0·10
14·18
0·20
52·34
0·03
0·56
27·32
rim
M106
cd
0·40
11·03
0·00
0·00
0·03
0·00
1·19
0·02
0·79
0·03
3·96
0·00
0·00
5·01
96·36
0·00
0·01
0·15
0·01
7·54
0·17
8·91
0·40
31·78
0·00
0·00
47·40
core
M58
cd
0·37
11·02
0·00
0·00
0·03
0·00
1·24
0·02
0·72
0·04
3·97
0·00
0·00
5·00
96·56
0·00
0·01
0·14
0·01
7·94
0·22
8·15
0·55
32·00
0·00
0·00
47·54
rim
M58
sill
—
3·00
0·00
0·00
0·00
0·00
0·00
0·00
0·00
0·01
2·00
0·00
0·00
0·99
98·34
0·00
0·01
0·00
0·02
0·01
0·00
0·00
0·38
61·74
0·01
0·01
36·16
core
M58
GOSCOMBE AND HAND
P–T PATHS IN EASTERN HIMALAYA
3·46
3·67
19·12
Al2O3
8·94
0·00
0·02
8·31
0·00
0·19
9·86
3·99
100·33
MnO
MgO
CaO
Na2O
K 2O
H 2O
Total
3·10
1716
0·19
1·05
0·00
Fe2+
Mn
0·53
X(Fe)
2·00
2·00
9·71
OH
Sum
0·92
0·03
0·95
Na
K
1·01
0·50
9·68
0·02
0·00
0·93
0·00
Mg
Ca
2·70
0·50
9·69
2·00
0·80
0·07
0·00
1·08
0·00
1·08
0·19
1·68
0·09
0·49
9·73
2·00
0·84
0·07
0·00
1·10
0·00
1·07
0·19
1·66
0·09
2·71
99·05
3·96
8·73
0·44
0·03
9·72
0·00
16·92
3·32
18·62
1·58
35·74
rim
M106
bi
0·67
9·66
2·00
0·91
0·03
0·00
0·62
0·01
1·26
0·22
1·82
0·14
2·65
100·51
3·95
9·40
0·23
0·00
5·49
0·14
19·86
3·90
20·30
2·41
34·83
core
M71
bi
0·64
9·68
2·00
0·91
0·02
0·00
0·70
0·01
1·26
0·22
1·76
0·13
2·67
100·54
3·95
9·43
0·16
0·00
6·14
0·22
19·77
3·88
19·60
2·25
35·15
rim
M71
bi
0·29
8·97
2·00
0·69
0·28
0·00
0·05
0·00
0·02
0·03
2·81
0·02
3·07
100·12
4·57
8·25
2·21
0·00
0·48
0·00
0·39
0·65
36·36
0·43
46·77
core
M106
mu
0·30
8·98
2·00
0·69
0·26
0·00
0·07
0·00
0·03
0·05
2·81
0·02
3·05
99·25
4·52
8·19
2·04
0·00
0·68
0·00
0·60
1·00
35·88
0·31
46·03
rim
M106
mu
0·36
8·96
2·00
0·86
0·08
0·00
0·07
0·00
0·04
0·06
2·73
0·03
3·09
100·63
4·55
10·21
0·63
0·03
0·71
0·00
0·74
1·23
35·09
0·52
46·91
core
M71
mu
0·45
8·96
2·00
0·86
0·09
0·00
0·06
0·00
0·05
0·07
2·74
0·02
3·07
100·87
4·55
10·25
0·72
0·00
0·62
0·00
0·88
1·46
35·31
0·44
46·63
rim
M71
mu
0·47
17·90
8·00
0·00
0·00
0·01
2·22
0·00
1·98
0·35
2·79
0·01
2·54
100·43
11·54
0
0
0·05
14·34
0·04
22·76
4·46
22·77
0·07
24·39
core
M106
chl
0·45
17·26
2·00
0·00
0·18
2·70
2·14
0·02
1·74
0·31
1·02
0·18
6·97
99·16
1·99
0·01
0·60
16·72
9·52
0·17
13·80
2·71
5·75
1·60
46·20
core
M87
hn
0·51
17·32
2·00
0·14
0·18
2·13
1·99
0·02
2·09
0·37
1·49
0·03
6·88
99·38
1·98
0·75
0·62
13·15
8·85
0·14
16·53
3·24
8·33
0·28
45·52
core
M42
hn
0·48
17·79
2·00
0·43
0·31
1·97
2·05
0·09
1·93
0·34
2·50
0·10
6·07
99·22
1·96
2·19
1·04
12·03
8·97
0·68
15·05
2·95
13·86
0·83
39·67
core
M1
hn
NUMBER 12
0·00
1·01
0·18
1·68
1·70
Al
Fe3+
2·66
0·20
2·65
0·21
Si
99·11
3·97
8·25
0·51
0·05
9·56
0·08
17·02
3·34
18·90
1·63
35·80
core
M106
bi
VOLUME 41
Ti
98·70
3·94
9·53
0·14
0·00
15·82
3·27
16·69
Fe2O3
FeO
18·74
35·03
35·19
SiO2
rim
core
TiO2
M58
M58
Sample:
bi
bi
Mineral:
Table C: continued
JOURNAL OF PETROLOGY
DECEMBER 2000
1717
0·00
0·00
0·00
0·20
0·76
0·01
0·00
0·00
0·00
4·98
Fe2+
Mn
Mg
Ca
Na
K
Sr
Ba
OH
Sum
—
0·00
Fe3+
X( Fe)
1·20
Al
99·37
Total
0·00
0·00
H2O
Ti
0·00
BaO
2·81
0·00
SrO
Si
0·10
0·00
MgO
K2O
11·15
0·00
MnO
4·19
0·00
FeO
8·83
0·05
Fe2O3
CaO
0·00
22·86
Al2O3
Na2O
0·00
0·01
—
4·98
0·00
0·00
0·01
0·01
0·42
0·55
0·00
0·00
0·00
0·00
1·51
0·00
2·48
98·67
0·00
0·00
0·26
0·20
4·71
0·04
0·00
0·08
27·90
54·31
63·31
TiO2
core
core
SiO2
M87
M58
Sample:
pl
pl
Mineral:
—
5·65
0·00
0·00
0·01
0·00
0·70
0·97
0·00
0·00
0·00
0·00
1·89
0·00
2·08
99·49
0·00
0·00
0·23
0·02
0·76
19·36
0·01
0·04
0·00
0·08
34·37
0·02
44·57
core
M42
pl
—
4·97
0·00
0·00
0·01
0·01
0·80
0·15
0·00
0·00
0·00
0·00
1·14
0·00
2·86
99·15
0·00
0·04
0·25
0·09
9·27
3·18
0·00
0·02
0·00
0·01
21·81
0·01
64·47
core
M106
pl
—
5·00
0·00
0·00
0·01
0·00
0·19
0·81
0·00
0·00
0·00
0·01
1·76
0·00
2·22
99·12
0·00
0·00
0·24
0·05
2·16
16·23
0·00
0·02
0·00
0·17
32·34
0·02
47·88
core
M1
pl
—
4·97
0·00
0·00
0·00
0·83
0·14
0·00
0·00
0·00
0·00
0·00
1·01
0·00
2·99
98·25
0·00
0·00
0·00
13·96
1·57
0·09
0·00
0·07
0·00
0·03
18·36
0·00
64·18
core
M58
kf
—
5·01
0·00
0·01
0·01
0·90
0·10
0·00
0·00
0·00
0·00
0·00
1·01
0·00
2·98
98·61
0·00
0·32
0·31
15·01
1·05
0·05
0·00
0·00
0·00
0·03
18·32
0·01
63·48
core
M87
kf
—
5·00
0·00
0·00
0·00
0·92
0·08
0·00
0·00
0·00
0·00
0·00
1·03
0·00
2·97
99·39
0·00
0·19
0·18
15·42
0·91
0·01
0·00
0·06
0·00
0·04
18·71
0·00
63·86
core
M71
kf
—
5·12
0·00
0·00
0·01
0·01
0·20
1·09
0·00
0·00
0·00
0·00
1·66
0·00
2·15
94·28
0·00
0·00
0·21
0·16
2·11
20·33
0·00
0·02
0·04
0·00
28·31
0·00
43·09
core
M42
me
—
8·01
1·00
0·00
0·00
0·00
0·00
2·03
0·01
0·01
0·00
0·50
2·47
0·01
2·98
100·43
1·92
0·00
0·00
0·00
0·00
24·27
0·09
0·18
0·00
8·53
26·93
0·24
38·24
core
M1
ep
0·42
3·99
0·00
0·00
0·00
0·00
0·02
0·89
0·56
0·01
0·41
0·08
0·10
0·01
1·91
99·06
0·00
0·00
0·00
0·01
0·31
21·44
9·72
0·31
12·63
2·71
2·29
0·31
49·30
core
M87
cpx
0·44
3·99
0·00
0·00
0·00
0·00
0·01
1·01
0·51
0·02
0·40
0·04
0·04
0·00
1·96
99·39
0·00
0·00
0·00
0·00
0·10
24·36
8·89
0·59
12·38
1·52
0·85
0·07
50·54
core
M42
cpx
0·23
4·01
0·00
0·00
0·00
0·00
0·04
0·94
0·64
0·02
0·19
0·12
0·18
0·01
1·87
101·04
0·00
0·00
0·00
0·00
0·56
23·62
11·62
0·62
6·00
4·20
4·00
0·22
50·19
core
M1
cpx
GOSCOMBE AND HAND
P–T PATHS IN EASTERN HIMALAYA
JOURNAL OF PETROLOGY
VOLUME 41
NUMBER 12
DECEMBER 2000
APPENDIX D
Table D: End-member activities were calculated as follows, largely after Powell & Holland
(1985); abbreviations after Powell & Holland (1988)
Assuming ideal solid solution
Amphibole
ahb
= 37·9XVA(XCaM2)2(XMgM3)3(XMgM1)(XAlM1) (XSiT1)3(XAlT1)
aparg
= 64XNaA(XCaM2)2(XMgM3)3(XMgM1)(XAlM1)(XSiT1)2(XAlT1)2
aed
= 9·48XNaA(XCaM2)2(XMgM3)3(XMgM1)2(XSiT1)3(XAlT1)
aftr
= XVA(XCaM2)2(XFeM3)3(XFeM1)2(XSiT1)4
aanth
= XVA(XMgM2)2(XMgM3)3(XMgM3)2(XSiT1)4
Cordierite
acrd
= (XMgM2)2
afcrd
= (XFeM2)2
Feldspar
aab
= (XNaA)
aksp
= (XKA)
White mica
amu
= 4XKA(XVM2)(XAlM1)2(XAlT1)(XSiT1)(XSiT2)2
acel
= 4XKA(XVM2)(XAlM1)(XMgM1)(XSiT1)2(XSiT2)2
apa
= 4XNaA(XVM2)(XAlM1)2(XAlT1)(XSiT1) (XSiT2)2
Biotite
aphl
= 4XKA(XMgM1)(XMgM2)2(XSiT1)(XAlT1)(XSiT2)2
aann
= 4XKA(XFeM1)(XFeM2)2(XSiT1)(XAlT1)(XSiT2)2
aeast
= XKA(XAlM2)(XMgM2)2(XAlT1)2(XSiT2)2
anaph
= 4XNaA(XMgM1)(XMgM2)2(XSiT1)(XAlT1)(XSiT2)2
Chlorite
aames
= (XMgM2)4(XAlM1)2(XAlT1)2
Pyroxene
adi
= (XCaM2)(XMgM1)(XSi)2
ahed
= (XCaM2)(XFeM1)(XSi)2
acats
= (XCaM2)(XAlM1)(XSi)(XAl)
Staurolite
amst
= (XMgM2)2(XAlM1)9(XSiT1)4
afst
= (XFeM2)2(XAlM1)9(XSiT1)4
Scapolite
ame
= (XAlT3)3(XCaA)4
Oxides
aq = asill = aky = 1
Assuming non-ideal solid solution
Plagioclase
= an(XCaA)
an after Hoisch (1990)
apy
= (pyXMgM2)3(XAlM1)2(XSiT1)3
py after Berman (1990)
aalm
= (almXFeM2)3(XAlM1)2(XSiT1)3
alm after Berman (1990)
agr
= (grXCaM2)3(XAlM1)2(XSiT1)3
gr after Berman (1990)
aan
Garnet
1718
GOSCOMBE AND HAND
P–T PATHS IN EASTERN HIMALAYA
Mineral
Abbreviation
Mineral
Abbreviation
Actinolite
act
Magnetite
mg
Akermanite
ak
Margarite
ma
Almandine
alm
Meionitic scapolite
me
Andalusite
and
Microcline
mic
Antiperthite
aper
Monazite
mon
Apatite
ap
Monticellite
mo
Biotite
bi
Muscovite
mu
Calcite
cc
Myrmekite
myr
Chlorite
chl
Olivine
ol
Chloritoid
ctd
Opaque
op
Clinopyroxene
cpx
Perthite
per
Clinozoisite
clz
Phlogopite
phl
Cordierite
cd
Plagioclase
pl
Diopside
di
Pyrite
py
Dolomite
dol
Quartz
q
Epidote
ep
Rutile
rut
Ferrosilite
fs
Scapolite
scap
Fibrolite
fibr
Sericite
seri
Forsterite
fo
Sillimanite
sill
Garnet
gn
Sphene
sph
Gedrite
gd
Staurolite
st
Haematite
hm
Stilpnomelane
stil
Hercynitic spinel
sp
Ti-magnetite
timg
Hornblende
hn
Tourmaline
tm
Ilmenite
ilm
Tremolite
tr
K-feldspar
kf
Wollastonite
woll
Kyanite
ky
Zircon
zr
1719