partial melting and decompression of the thor-odin dome, shuswap metamorphic core complex, canadian...

Partial melting and decompression of the Thor-Odin dome,

Shuswap metamorphic core complex, Canadian Cordillera

Britt H. Norlander a,*, Donna L. Whitney a, Christian Teyssier a, Olivier Vanderhaeghe b

aDepartment of Geology and Geophysics, University of Minnesota, Minneapolis, MN 55455, USAbUniversite Henri Poincare Nancy 1, Geologie et Gestion des Ressources Minerales et Energetiques, BP 239 54506,

Vandoeuvre-les-Nancy Cedex, France

Received 15 March 2001; accepted 18 December 2001

Abstract

The Thor-Odin dome region of the Shuswap metamorphic core complex, British Columbia, contains migmatitic rocks

exhumed from the deep mid-crust of the Cordilleran orogen. Extensive partial melting occurred during decompression of the

structurally deepest rocks, and this decompression path is particularly well recorded by mafic boudins of silica-undersaturated,

aluminous rocks. These mafic boudins contain the high-temperature assemblages gedrite + cordierite + spinel + corundum+

kyanite/sillimaniteF sapphirineF hogbomite and gedrite + cordierite + spinel + corundum+ kyanite/sillimanite + garnetF staur-

staurolite (relict)F anorthite. The boudins are interlayered with migmatitic metapelitic gneiss and orthogneiss in this region.

The mineral assemblages and reaction textures in these rocks record decompression from the kyanite zone (P> 8–10 kbar) to

the sillimanite–cordierite zone (P< 5 kbar) at Tf 750 jC, with maximum recorded temperatures of f 800 jC. Evidence forhigh-temperature decompression includes the partial replacement of garnet by cordierite, the partial to complete replacement of

kyanite by corundum+ cordierite + spinel (hercynite)F sapphirineF hogbomite symplectite, and the replacement of some

kyanite grains by sillimanite. Kyanite partially replaced by sillimanite, and sillimanite with coronas of cordieriteF spinel are

also observed in the associated metapelitic rocks. Partial melt from the surrounding migmatitic gneisses has invaded the mafic

boudins. Cordierite reaction rims occur where minerals in the boudins interacted with leucocratic melt. When combined with

existing structural and geochronologic data from migmatites and leucogranites in the region, these petrologic constraints

suggest that high-temperature decompression was coeval with partial melting in the Thor-Odin dome. These data are used to

evaluate the relationship between partial melting of the mid-crust and localized exhumation of deep, hot rocks by extensional

and diapiric processes. D 2002 Elsevier Science B.V. All rights reserved.

Keywords: Omineca Belt; Petrology; Geothermobarometry; Decompression; Partial melting; Metamorphic core complex

1. Introduction

Recent geophysical studies of thickened crust in

modern orogenic settings have recognized the exis-

tence of zones in the mid-crust containing approxi-

mately 20% partial melt (Nelson et al., 1996; Schilling

and Partzsch, 2001). These data are consistent with

observations from the cores of exhumed orogens

where the widespread occurrence of migmatites indi-

cates a high percentage of melt was present in the mid-

crust (Brown, 1994). High-grade metamorphic rocks in

the migmatitic cores of exhumed orogens are typically

0024-4937/02/$ - see front matter D 2002 Elsevier Science B.V. All rights reserved.

PII: S0024 -4937 (02 )00075 -0

* Corresponding author. Tel.: +1-612-624-8557; fax: +1-612-

625-3819.

E-mail address: [email protected] (B.H. Norlander).

www.elsevier.com/locate/lithos

Lithos 61 (2002) 103–125

characterized by clockwise P–T paths, reflecting bur-

ial and heating during crustal thickening, followed by

decompression and cooling. In some cases, geochro-

nologic and structural evidence suggest a temporal link

between partial melting and exhumation (Lister and

Baldwin, 1993; Brown and Dallmeyer, 1996; Foster

and Fanning, 1997; Vanderhaeghe and Teyssier, 1997).

Evaluating the connection between partial melting

and decompression is complicated by the positive

feedback between the two processes. Dehydration-

melting reactions for typical crustal lithologies have a

positive slope in P–T space, and thus partial melting

may occur either by an increase in temperature (e.g.,

heating during burial) or decrease in pressure (e.g.,

decompression as a result of unroofing) (Fig. 1). Path

A (Fig. 1) demonstrates that rocks in the mid-crust

can cross dehydration-melting reactions during a

decrease in pressure without significant change in

temperature. Therefore, partial melting in the crust

can occur as a result of unroofing of high-grade rocks.

Alternatively, H2O-saturated partial melting can

occur during burial and heating (Path B, Fig. 1). A

drastic reduction in strength is expected in a partially

molten layer in the crust (Arzi, 1978; Van der Molen

and Paterson, 1979; Vanderhaeghe and Teyssier,

2001), and could lead to the exhumation of high-grade

rocks by late-orogenic collapse (Vanderhaeghe and

Teyssier, 1997). There is abundant evidence for a

temporal link between granite emplacement and the

development of metamorphic core complexes beneath

low-angle detachment zones (Crittenden et al., 1980,

and references therein; Lister and Baldwin, 1993;

Foster and Fanning, 1997; Vanderhaeghe et al.,

1999b). The localization of strain in orogenic hinter-

lands may be enhanced by the presence of melt

(Hollister and Crawford, 1986; Hollister, 1993). In

addition, significant partial melting in the crust may

result in the formation of diapirs which accommodate

decompression during the buoyant rise of felsic melt

(Schuilling, 1960; Thompson et al., 1968; Howard,

1980; Faure and Cottereau, 1988; Calvert et al., 1999).

A positive feedback may occur when tectonic unroof-

ing enhanced by the presence of melt causes more

partial melting during decompression. The key to

discerning the respective roles of H2O-saturated partial

melting (Path B, Fig. 1) and partial melting related to

decompression (Path A, Fig. 1) is to determine where

partial melting occurred on the P–T path.

The Thor-Odin region of the Shuswap metamor-

phic core complex (Fig. 2) is an area in which this

relationship can be explored. The crust in this region

was nearly doubled in thickness to >60 km during

Mesozoic and early Cenozoic accretion of terranes to

Fig. 1. P–T diagram showing the relative positions and slopes of the

H2O-saturated pelite solidus and a typical dehydration-melting reac-

tion for pelitic compositions (after Le Breton and Thompson, 1988).

The arrows are drawn in the direction of melt production. Path A

represents partial melting occurring as a result of crossing of dehy-

dration-melting reactions due to decompression, whereas path B

represents a prograde path where partial melting occurs as the result

of burial and heating.

Fig. 2. (a) Simplified map of the Canadian Cordillera (modified after Wheeler and McFeely, 1991). Shuswap metamorphic core complex (MCC)

is located within the Omineca belt in the hinterland of the Foreland fold and thrust belt. The migmatitic domes of the complex are labeled M—

Malton; FC—Frenchman Cap; TO—Thor-Odin dome (location of study area); V—Valhalla. (b) Geologic map of the Shuswap metamorphic

core complex at the latitude of the Thor-Odin dome (after Vanderhaeghe and Teyssier, 1997). Location of leucogranite sample dated with U–Pb

SHRIMP on zircon (97046) is shown. Cross-section of A–AV is shown in (c). VLF—Victor lake fault. (c) Cross-section across the Thor-Odin

dome showing the relationship of the upper, middle and lower units to the low-angle detachment faults, specifically the association between the

migmatitic lower unit, the network of sills and dikes in the middle unit, and the leucogranite laccoliths (after Vanderhaeghe et al., 1999b).

B.H. Norlander et al. / Lithos 61 (2002) 103–125104

B.H. Norlander et al. / Lithos 61 (2002) 103–125 105

western North America (Coney and Harms, 1984;

Parrish et al., 1988). The rocks that were buried

during collision were fertile paleomargin sediments

and were partially melted, with estimates of total melt

production > 40 vol.% in some areas (Nyman et al.,

1995). Structural and geochronologic data show that

partial melting and extensional deformation were

coeval (Vanderhaeghe et al., 1999b).

In this paper, we present petrologic data from the

deepest exposed structural level in the Thor-Odin

dome. In the study area, migmatitic sillimanite–K-

feldspar gneiss is interlayered with garnet–horn-

blende amphibolites and mafic boudins of orthoam-

phibole-bearing rocks. The mafic boudins contain the

assemblages corundum + spinel + cordierite + ged-

riteF sapphirineF hogbomite and corundum+ spin-

el + cordierite + gedrite + garnet. Reaction textures in

these rocks indicate steep decompression paths at

high temperatures, which probably corresponds to ex-

tensional collapse of the region. Aluminous gneisses

with similar compositions to the mafic boudins are

known in other high-grade regional metamorphic

terrains (e.g., Warren, 1979; Ellis, 1980; Baker et

al., 1987; Johansson and Moeller, 1986; Schumacher

and Robinson, 1987; Droop, 1989; Mohan and Wind-

ley, 1993; Liati and Seidel, 1994; Raith et al., 1997;

Baba, 1999; Moller, 1999), but have not been re-

ported in the Cordillera, with the exception of a lo-

cality in the Okanogan complex in the southernmost

part of the Omineca Belt (Harvey and Hoisch, 1994).

Mafic rocks such as these are of particular interest

because they commonly preserve mineral assem-

blages that indicate high metamorphic temperatures

and contain spectacular reaction textures that record

pressure–temperature path information necessary to

understand the tectonic and thermal evolution of the

region.

The Thor-Odin dome of the Shuswap metamorphic

core complex therefore contains the key elements

needed to study the link between high-temperature

metamorphism, partial melting, and decompression.

The bulk compositions preserve a large portion of the

P–T path, and leucosomes in associated metapelitic

rocks preserve the partial melting history. These data

are combined with existing geochronologic and struc-

tural data from the region to evaluate the relation

between partial melting and various mechanisms for

the rapid exhumation of high-grade rocks.

2. Geologic setting

2.1. Regional geology

The Shuswap complex, the largest of the Cordil-

leran metamorphic core complexes (Crittenden, 1980;

Armstrong, 1982), is located in the southern Omineca

Belt, in the metamorphic and plutonic hinterland of the

RockyMountain Foreland Belt (Fig. 2a). The Omineca

Belt was exhumed subsequent to collision between

accreted terranes to the west and the North American

continent (Monger et al., 1982), and is comprised of

metamorphosed Proterozoic and Paleozoic miogeocli-

nal strata, fragments of accreted terranes, exhumed

Precambrian North American basement, and Paleozoic

to Tertiary plutons. In the southern Omineca Belt, the

Shuswap metamorphic core complex consists of high-

grade metamorphic rocks that were exhumed along

low-angle detachment zones and high-angle normal

faults (Fig. 2b) during Eocene–Oligocene time.

The lowest structural unit of the Shuswap complex

is exposed in domal culminations aligned along the

strike of the belt. From north to south, these are the

Malton, Frenchman’s Cap, Thor-Odin, and Valhalla

domes (Fig. 2a). The cores of these domes are com-

prised of high-grade polymetamorphic rocks, includ-

ing migmatitic para- and orthogneiss and amphibolite

that are interpreted to be part of the Windermere

Supergroup with Proterozoic granodiorite intrusions

(Wanless and Reesor, 1975; Armstrong et al., 1991;

Parkinson, 1991). Unconformably overlying the core

gneisses are metasediments, the metamorphosed

equivalent of a Lower Paleozoic to Lower Mesozoic

platform sequence (Reesor and Moore, 1971; Read

and Brown, 1981; Okulitch, 1984; Scammell and

Brown, 1990; Carr, 1992).

The Thor-Odin dome of the Shuswap complex is

located south of Revelstoke, British Columbia (Fig.

2b). At the latitude of the Thor-Odin dome, a f 15-

km thick crustal section is exposed and is divided into

three superposed crustal units (Fig. 2b,c; Parrish et al.,

1988; Vanderhaeghe and Teyssier, 1997). The upper

unit lies in the hanging walls of shallowly dipping

detachment faults and records Cretaceous cooling

ages (Mathews, 1981; Colpron et al., 1996).

The middle unit, below the detachment faults,

comprises discontinuous layers of metapelitic schist,

calc-silicate, marble, amphibolite, and quartzite. The


dominant fabric is a shallowly dipping foliation and

E–W to NE–SW trending mineral lineation. Thermo-

barometric data from the metasedimentary rocks of

the middle unit constrain the conditions of peak

metamorphism to amphibolite/upper-amphibolite

facies. Amphibolite boudins from Three Valley Gap

(Fig. 2b) contain orthopyroxene and yield P–T esti-

mates of 620–685 jC and 6–7 kbar (Ghent et al.,

1977). Just west of this locality, pressures and temper-

atures of 7.5–9 kbar and 720–820 jC were estimated

for sillimanite and K-feldspar bearing migmatites

(Nyman et al., 1995). Temperatures and pressures of

625–825 jC and 6–7.5 kbar were estimated from

both garnet–hornblende amphibolites and silliman-

ite–K-feldspar bearing metapelitic rocks (Norlander

et al., 1999). Coronas of plagioclase + hornblende +

quartz around garnet occur in garnet–hornblende

amphibolites south of Mt. Odin and suggest decom-

pression (Norlander et al., 1999). However, the occur-

rence of these textures is not widespread in the middle

unit and cordierite has not been found in the middle

unit.

The middle unit is migmatitic and permeated by a

network of granitic sills and dikes originating in the

lower unit and probably connecting to leucogranite

laccoliths (Ladybird suite) within and just below the

detachment zone (Fig. 2c) (Carr, 1992). The structural

relationship between the laccoliths and the lower unit,

together with geochemical data, suggest that the

leucogranites originated from the lower unit (Vander-

haeghe et al., 1999b). A mylonitic fabric in the

leucogranites is developed in the detachment zone.

Garnet-bearing leucosomes are common throughout

the middle unit, but rare leucosomes containing mag-

matic andalusite indicate that partial melting contin-

ued to low pressures.

This study focuses on the lower unit in the area

around Mt. Odin where gedrite-cordierite rocks occur

as boudinaged layers within migmatitic sillimanite +

K-feldspar-bearing metapelitic rocks, garnet–horn-

blende amphibolites, and granitoids (Duncan, 1984;

Bruce et al., 1995) (Fig. 3). The E–W mineral lin-

eation in the migmatitic gneisses is consistent with

the fabric in the middle unit. The foliation wraps

Fig. 3. Detailed geologic map of the study area in the Thor-Odin dome (modified after Vanderhaeghe et al., 1999b). Study area is focused

around Mt. Odin where gedrite–cordierite rocks and garnet amphibolites are exposed as large boudins in the migmatitic gneisses. Locations of

samples used for thermobarometry and locations of zircon samples dated with U–Pb SHRIMP (97013, 97010) are shown.


around the dome and is steeper along the margins of

the dome. In the region of Mt. Odin, Vanderhaeghe

and Teyssier (1997) documented a metatexite–dia-

texite transition where the percentage of partial melt

increases downward in the lower unit. This transition

cross-cuts the stratigraphic sequence in the lower

unit, and thus Vanderhaeghe and Teyssier (1997)

suggest that the diatexites rose ‘‘en masse’’ with

respect to the metatexites. Combined structural study

and U–Pb geochronology using the SHRIMP analy-

tical facility at the Research School of Earth Scien-

ces, Australian National University, indicate that the

pervasive foliation and lineation developed in the

presence of melt during Paleocene time (Vander-

haeghe et al., 1999b).

2.2. Thermochronology

A large thermochronologic dataset for the Shuswap

complex exists, including samples collected from our

study area at Mt. Odin. U–Pb SHRIMP analyses of

zircons by Vanderhaeghe et al. (1999b) were obtained

from two leucosomes at Mt. Odin (see Fig. 3 for

sample localities) and one leucogranite sample located

in the detachment zone at Mabel Lake (Fig. 2b). U–

Pb SHRIMP ages of high-U zircon rims reveal that the

migmatitic rocks of the lower unit crystallized atf 56

Ma, while zircons from syntectonic leucogranite

found in the detachment zone yield ages of f 60

Ma (Fig. 4), consistent with U–Pb zircon ages from

the Ladybird suite of 62–56 Ma (Carr, 1992).

Argon thermochronology (Vanderhaeghe, 1997)

and fission-track analyses (Lorencak et al., 2001)

from the middle and lower units constrain the cooling

history of the terrain. 40Ar/39Ar ages from the middle

and lower units are similar to each other. Hornblende

ages range from 59 to 54 Ma (maximum ages due to

excess argon) and white mica and biotite ages are

tightly clustered between 49.5 and 47 Ma (Fig. 4).

These ages are consistent with K-feldspar 40Ar/39Ar

ages ranging from 50 to 43 Ma, except for two

samples in the immediate footwall of the Columbia

River Fault which are < 30 Ma (Fig. 4). Zircon

fission-track ages from both the middle and lower

units cluster around 50–45 Ma; apatite ages are

typically 50–45 Ma in the middle unit and 45–40

Ma in the lower unit (Fig. 4). In addition, young

apatite ages corroborate Oligocene cooling in the

immediate footwall of the Columbia River fault and

Victor Lake fault (f 30 Ma) (Fig. 4). The U–Pb

SHRIMP zircon ages and the hornblende 40Ar/39Ar

ages suggest that an early, high-temperature cooling

rate may have been as high as 100 jC/Ma, with

slower cooling ( < 50 jC/Ma) occurring at lower tem-

peratures.

3. Petrology and mineral chemistry

In the following sections, we describe the gedrite–

cordierite-bearing boudins and associated metapelitic

rocks and garnet–hornblende amphibolites to demon-

strate the high metamorphic temperatures and pres-

sures experienced by the lower unit and to document

the unusual mineralogy of the gedrite-bearing rocks.

Fig. 4. Synthesis of thermochronologic data from the Thor-Odin

dome. The slopes of selected cooling rates are shown. A apatite

fission-track sample from the upper unit, located in the hanging wall

of the Okanagan Detachment west of Sicamous and north of Salmon

Arm (Fig. 2b), records late Jurassic ages. The thermochronologic

data show fast cooling from f 60 to 45 Ma with most samples

below 100 jC at 40 Ma, with the exception of samples located in

the footwall of high-angle normal faults (Victor Lake Fault and

Columbia River Fault, Fig. 2b). These samples record later cooling

at f 30 Ma.


Mineral compositions and major element distribu-

tion maps were obtained using a JEOL JXA-8900

electron microprobe at the University of Minnesota.

Operating conditions for quantitative analysis (WDS)

were 15 kV accelerating voltage, 20 nA beam current,

and a range of beam diameters (focused for garnet,

defocused to 5 Am for minerals containing volatile

elements). X-ray maps were determined using a beam

current of 100 nA, 50 ms dwelltime, and 5–9 Ambeam diameters. Natural mineral standards and the

ZAF matrix correction routine were used. Represen-

tative analyses are given in Tables 1–4.

3.1. Gedrite–cordierite rock

There are two distinct assemblages in the gedrite–

cordierite rocks: garnet-bearing and sapphirine-bear-

ing. At this time, we have not observed garnet and

sapphirine in the same sample. In both of these rock

types, gedrite is the most abundant mineral in the

matrix (>30–40%), and occurs as coarse sprays

varying in length from 2 mm to 30 cm (Fig. 5a).

Kyanite occurs in both of these rock types; it is

commonly partially to completely replaced by corun-

dum + cordieriteF ilmenite symplectite cores and

spinel + cordieriteF sapphirine symplectite rims (Fig.

5c–g). These symplectite regions are separated from

the matrix gedrite by a band of cordierite (Fig. 5c,d).

Symplectitic regions lacking a relict phase are inferred

to be pseudomorphs of kyanite, or in some cases

staurolite, based on shape and the presence of the

same symplectitic assemblages with relict kyanite and

staurolite. In the case of kyanite, another phase such

as garnet or gedrite must have been involved in the

pseudomorph-forming reactions. Sillimanite (pris-

matic and fibrous varieties) has partially replaced

some kyanite grains and also occurs as isolated matrix

grains rimmed by cordierite.

3.1.1. Garnet-bearing gedrite–cordierite rock

Large garnets (1–5 cm diameter) contain inclu-

sions of rutile, ilmenite, apatite, and staurolite, and

are rimmed by coronas (1–3 mm) of euhedral ged-

rite (XMg = 0.60–0.67;Table1) + cordieriteF spinel Filmenite (Fig. 5b). Gedrite fills fractures in garnet,

along with anorthite + cordierite, and it occurs as in-

clusions in garnet and as an abundant matrix phase. X-

ray maps show that some gedrite crystals are zoned,

particularly those in the vicinity of garnet, such as

gedrite grains in embayed (resorbed) garnet rim

regions. Zoned grains have moreMg-rich cores relative

to rims; rims are more Fe- and Na-rich than core

regions. Gedrite inclusions in garnet have the same

composition as the core regions of matrix gedrite.

X-ray maps of other phases were also determined to

ascertain zoning patterns. Garnet grains up to 4 cm in

diameter are homogeneous in composition (XPrp = 0.50,

XAlm = 0.48; Table 1) except for a narrow ( < 100 Am)

retrograde rim with increased Fe andMn and decreased

Mg concentrations. Cordierite is abundant in the

matrix, in symplectitic regions, and in garnet coronas,

and is homogeneous in composition (XMg = 0.84)

throughout the analyzed samples.

Large (1 cm long) staurolite inclusions in garnet

are partially replaced by spinel + cordierite + ilmenite

symplectite and are separated from the host garnet by

a band of cordierite (Fig. 5h). The symplectite also

formed in fractures within the staurolite inclusions

(Fig. 5h). Some inclusions are comprised of two

intergrown (twinned) staurolite crystals. The cordier-

ite band and symplectitic region follow the outline of

both crystals in each twin. Staurolite contains rare

inclusions of quartz, the only quartz observed in these

rocks thus far.

Spinel occurs as symplectitic intergrowths with cor-

dierite in pseudomorphs (Fig. 5c–g), and as coarser,

isolated grains in the matrix. A typical composition

for both varieties is XMg = 0.40 (Table 1). Corundum

is found in the matrix as well as in the cores of

pseudomorphs (Fig. 5c,d,g) and in veins containing

large (cm-scale) corundum crystals. Corundum occurs

in several compositional varieties, including red/pink

(colorless in plane light) Cr-bearing corundum and a

dark blue variety with slightly more Ti and Fe than the

Cr-rich corundum. The ruby variety is most common

in pseudomorphs after kyanite and in veins, and the

blue variety occurs in pseudomorphs after staurolite

and as a matrix phase.

Accessory minerals are apatite, monazite, ilmen-

ite, and biotite. Relatively large (0.3–0.5 mm) apa-

tite grains occur throughout the matrix and are

rimmed by spinel in the cordierite + gedrite coronas

around garnet. Ilmenite is relatively abundant in

some samples (8%) and is intimately associated with

spinel both in the matrix and in the symplectites

(Fig. 5d). Biotite is locally abundant in boudins that


have beeninjected by leucocratic material, particu-

larly in regions surrounding pseudomorphs. For

example, in one sample a cordierite + corundum+ -

spinel pseudomorph after a tabular mineral is sur-

rounded by a zone of randomly oriented, coarse

biotite several millimeters wide which has overgrown

the matrix gedrite + cordierite. Some biotite grains

contain relict gedrite, although in other cases biotite

and gedrite are intergrown with each other along

straight grain boundaries.

Fig. 5. (a) Field photograph of gedrite–cordierite rock with gedrite occurring in coarse radiating sprays (garbenscheifer). A quarter is shown for

scale. (b) Field photograph of garnet-bearing gedrite–cordierite rock. Garnet occurs as large porphyroblasts with wide, symplectitic coronas of

gedrite + cordieriteF ilmenite. (c) Scanned image of a thin-section containing a tabular shaped symplectitic region (1.2 cm long) in the gedrite–

cordierite rocks. The core (light colored area) of the region comprises corundum+ cordierite symplectite with a rim of spinel + cordierite

symplectite. A rim of cordierite separates this region from matrix gedrite. (d) Photomicrograph of rim of symplectitic region in gedrite–

cordierite rock (long dimension = 2.5 mm). The core of the region (corundum+ cordierite) is on the right, with a rim of spinel + cordierite

separated from matrix gedrite (on left) by a band of cordierite. (e) Photomicrograph of sapphirine intergrown with symplectitic spinel and

cordierite in sapphirine-bearing gedrite–cordierite rock (long dimension = 3.0 mm). (f) Photomicrograph of hogbomite needles intergrown with

symplectitic spinel and cordierite in sapphirine-bearing gedrite–cordierite rock (long dimension = 1.6 mm). (g) Photomicrograph of sapphirine

intergrown with spinel + cordierite symplectite on the rim of relict kyanite in a sapphirine-bearing gedrite–cordierite rock (long dimension = 1.6

mm). (h) Photomicrograph of garnet-bearing gedrite–cordierite rock with large inclusions of staurolite in garnet (long dimension = 5 mm).

Staurolite is partially replaced at the rim by symplectitic spinel + cordierite and separated from the host garnet by a band of cordierite.

Table 1

Representative mineral compositions for garnet-bearing gedrite–cordierite rocks

Garnet core Garnet rim Gedrite Cordierite Spinel Ilmenite Biotite

SiO2 40.13 38.65 46.49 49.34 – – 38.41

TiO2 < d.l. 0.02 0.13 – 0.03 48.85 1.31

Al2O3 23.12 22.02 14.71 34.50 60.99 0.06 17.93

Cr2O3 – – – – 0.02 < d.l. –

FeO 22.55 31.17 17.73 3.50 27.70 48.49 9.29

MnO 0.26 1.26 0.21 0.03 0.04 0.20 0.01

MgO 13.06 6.74 18.31 10.36 10.51 1.46 20.05

ZnO – – – – 0.26 < d.l. –

CaO 0.57 0.55 0.28 0.01 – – 0.12

Na2O – – 1.11 0.19 – – 0.60

K2O – – < d.l. < d.l. – – 7.62

Total 99.69 100.41 98.96 97.93 99.55 99.06 95.34

Cations per 12 O 12 O 23 O 18 O 4 O 6 O 22 O

Si 3.00 3.01 6.55 4.98 – – 5.75

Al 2.04 2.02 4.11 1.96 0.00

AlIV 1.45 2.25

AlV1 0.99 0.91

Ti 0.00 0.00 0.01 – 0.00 1.89 0.15

Cr – – – – 0.00 0.00 –

Fe2 + 1.41 2.03 2.09 0.30 0.63 2.09 1.16

Mn 0.02 0.08 0.03 0.00 0.00 0.01 0.00

Mg 1.46 0.78 3.84 1.56 0.43 0.11 4.47

Zn – – – – 0.01 0.00 –

Ca 0.05 0.05 0.04 0.00 – – 0.02

Na – – 0.30 0.04 – – 0.18

K – – 0.00 0.00 – – 1.46

XMg – – 0.65 0.84 0.40 0.05 0.79

XAlm 0.48 0.69

XSps 0.01 0.03

XPrp 0.50 0.27

XGrs 0.02 0.02


In some samples, the gedrite–cordierite boudins

have been invaded by leucocratic melt. This quartzo-

feldspathicmaterial typically collects near large garnets

in pressure shadows, and/or occurs in mm-to cm-scale

veins that cross-cut the coarse gedrite crystals. Where

phases such as garnet, spinel, and kyanite have been

incorporated into the quartzofeldspathic veins, they are

surrounded by coronas of columnar cordierite. Gedrite

is commonly partially replaced by biotite near the

leucosomes, and large mats of sillimanite in radiating

sprays are present along some leucosome margins.

3.1.2. Sapphirine-bearing gedrite–cordierite rock

Textural relationships among matrix phases such

as gedrite, cordierite, spinel, kyanite, and ilmenite,

and the occurrence of symplectitic pseudomorphs of

kyanite are similar to the garnet-bearing gedrite–

cordierite rocks. The sapphirine-bearing rocks, how-

ever, lack garnet and associated staurolite and rutile.

Sapphirine (XMg = 0.76; Table 2) occurs as slender

crystals intergrown with spinel and cordierite in the

rims of tabular pseudomorphs, on the rims of relict

kyanite (Fig. 5e,g), and also within matrix cordierite.

Cordierite is Mg-rich (XMg = 0.88) and homogeneous

in composition throughout the samples, despite

occurring in different textural varieties (e.g., within

the pseudomorphs (Fig. 5), intergrown with gedrite).

Spinel is slightly more Mg-rich than average spinel

in the garnet-bearing rock (XMg = 0.52). Very fine-

grained needles of hogbomite occur in some spinel

+ cordierite symplectitic regions (Fig. 5f). The needles

occur in two dominant orientations at f 80j to each

other in thin section. Corundum occurs in two vari-

eties: as colorless, very fine-grained (50 Am) crystals

intergrown with cordierite in pseudomorphs after

kyanite, and as larger (200 Am) dark blue (in plane

light) corundum grains in the groundmass and in some

pseudomorphs. Biotite has partially replaced and is

Table 2

Representative mineral compositions for sapphirine-bearing gedrite–cordierite rocks

Sapphirine Gedrite Cordierite Hogbomite Spinel Ilmenite Biotite

SiO2 11.83 46.84 49.86 0.40 – – 39.43

TiO2 0.27 0.14 – 7.84 < d.l. 49.52 1.54

Al2O3 65.57 17.02 34.58 60.63 63.55 0.09 17.88

Cr2O3 < d.l. < d.l. – < d.l. 0.02 0.03 –

FeO 7.70 12.76 2.53 19.51 22.39 48.06 7.78

MnO 0.21 0.11 0.01 0.27 0.08 0.26 0.02

MgO 14.22 20.63 10.75 10.21 13.39 1.50 20.22

ZnO < d.l. – – < d.l. 0.17 < d.l. –

CaO – 0.19 0.04 – – < d.l. 0.04

Na2O – 0.92 0.14 – – – 0.49

K2O – < d.l. 0.02 – – – 8.46

Total 99.80 98.59 97.93 99.86 99.59 99.46 95.85

Cations per 20 O 23 O 18 O 31 O 4 O 6 O 22 O

Si 0.71 6.46 5.01 0.08 – – 5.84

Al 4.09 14.62 1.98 0.01

AlIV 2.30 1.54 2.16

AlV1 2.31 1.23 0.96

Ti 0.01 0.01 – 1.21 0.00 1.91 0.17

Cr 0.00 0.00 – 0.00 0.00 0.00 –

Fe2 + 0.38 1.47 0.21 3.34 0.50 2.06 0.96

Mn 0.01 0.01 0.00 0.05 0.00 0.01 0.00

Mg 1.26 4.24 1.61 3.11 0.53 0.12 4.46

Zn 0.00 – – 0.00 0.00 0.00 –

Ca – 0.03 0.00 – – 0.00 0.01

Na – 0.25 0.03 – – – 0.14

K – 0.00 0.00 – – – 1.60

XMg 0.76 0.74 0.88 0.48 0.52 0.05 0.82


intergrown with gedrite in some samples. Fine-

grained ilmenite occurs both within the matrix and

intergrown with spinel in the pseudomorphs, and

more rarely as inclusions in gedrite.

3.2. Sillimanite–garnet–K-feldspar gneiss

Metapelitic rocks that host the boudins are migma-

titic and contain granitic leucosomes bounded by

Fig. 6. (a) Photomicrograph of garnet porphyroblast in sillimanite–garnet–K-feldspar gneiss (long dimension = 6 mm). Garnet contains

inclusions of K-feldspar (also present in the matrix) and is partially replaced at the rim by cordierite. (b) Photomicrograph of sillimanite

pseudomorph after kyanite, partially replaced by spinel + cordierite symplectite (long dimension = 3.0 mm). Cordierite displays columnar grain

shape with grain boundaries orthogonal to sillimanite. (c) Field photograph of sillmanite pseudomorphs of kyanite defining an E–W lineation.

(d) Field photograph of garnet–hornblende amphibolite with large garnet porphyroblasts. (e) Photomicrograph of rim of garnet porphyroblast

partially replaced by symplectitic plagioclase + hornblende + quartz in garnet–hornblende amphibolite (long dimension = 3.0 mm).


biotite-rich selvages. The typical mineral assemblage

in the gneiss is biotite + garnet + quartz + plagioclase +

cordierite + fibrolite (F prismatic sillimanite) +K-feld-

spar + chlorite (retrograde) +muscovite (retrograde)Fkyanite (relict)F rutileF ilmeniteF apatiteF zircon.

Garnet occurs as large (1.2–1.5 cm) porphyroblasts

with abundant inclusions of all groundmass phases

(including sillimanite) except rutile (Fig. 6a). Garnet

is Fe-rich with XAlm = 0.68, and is slightly zoned

with higher concentrations of Ca and Fe and lower

concentrations of Mg in the core (Table 3). X-ray

maps show that some garnets have a retrograde, high-

Mn rim ( < 150 Am). Cordierite hosts fibrolite mats and

has partially replaced garnet rims in some samples; it is

also found as polygonal grains with inclusions of

fibrolite and rutile and rimmed by muscovite. Plagio-

clase is homogeneous in composition (XAn = 0.32), and

commonly contains blebs of K-feldspar. Quartz is

subhedral and has substantial subgrain development

and serrated grain boundaries indicating dynamic

recrystallization. Kyanite occurs as relict grains in

the matrix and has been partially replaced by silliman-

ite (Fig. 6b,c). In addition, sillimanite (and more rarely

relict kyanite) is found rimmed by spinel + cordierite

symplectite and separated from the matrix by a rim of

cordierite (Fig. 6b). The cordierite in these rims ex-

hibits a columnar shape with grain boundaries per-

pendicular to the margin of the symplectitic region

(Fig. 6b).

3.3. Garnet–hornblende amphibolites

The typical mineral assemblage of these rocks is

hornblende + garnet + biotite + plagioclase + quartz +

Table 3

Representative mineral compositions for metapelitic rocks

Garnet core Garnet rim Biotite Plagioclase K-feldspar Cordierite Muscovite Ilmenite

SiO2 38.31 38.49 33.94 59.60 64.35 48.72 45.04 –

TiO2 0.06 0.02 3.43 – – – 0.15 51.07

Al2O3 21.68 22.02 19.57 25.53 19.21 34.33 36.57 0.06

FeO 31.58 29.06 21.31 0.06 < d.l. 6.53 0.71 45.35

MnO 0.38 0.18 0.14 – – 0.24 < d.l. 2.09

MgO 5.00 6.15 7.59 – – 8.71 0.47 0.07

CaO 4.14 3.55 0.08 6.62 0.09 0.01 0.01 < d.l.

Na2O – – 0.19 7.64 2.00 0.11 0.46 –

K2O – – 9.65 0.13 14.36 0.01 10.85 –

Total 101.15 99.49 95.91 99.58 100.01 98.66 94.26 98.63

Cations per 12 O 12 O 22 O 8 O 8 O 18 O 22 O 6 O

Si 2.99 3.01 5.21 2.66 2.96 4.96 6.06 –

Al 2.00 2.03 1.35 1.04 4.12 0.00

AlIV 2.79 1.95

AlV1 0.75 3.85

Ti 0.00 0.00 0.40 – – – 0.02 1.97

Fe2 + 2.06 1.90 2.74 0.00 0.00 0.56 0.08 1.95

Mn 0.03 0.01 0.02 – – 0.02 0.00 0.09

Mg 0.58 0.72 1.74 – – 1.32 0.10 0.01

Ca 0.35 0.30 0.01 0.32 0.01 0.00 0.00 0.00

Na – – 0.06 0.66 0.18 0.02 0.12 –

K – – 1.89 0.01 0.84 0.00 1.86 –

XMg – – 0.39 – – 0.70 0.54 0.00

XAn – – – 0.32 0.01 – – –

XAb – – – 0.67 0.17 – – –

XOr – – – 0.01 0.82 – – –

XAlm 0.68 0.65

XSps 0.01 0.00

XPrp 0.19 0.25

XGrs 0.12 0.10


rutile + ilmeniteF titaniteF apatiteFmonazite. Cli-

nopyroxene was observed in one sample, but no

orthopyroxene has been found in the Thor-Odin am-

phibolites. Hornblende is tschermakitic in composition

(Table 4). It occurs both in the matrix, ranging in size

from 0.2 to 0.7 mm, and as smaller grains in symplec-

titic regions around garnet along with plagioclase and

quartz (Fig. 6e; Table 4). Garnet porphyroblasts range

in size from 3 to 8 mm in diameter and are homoge-

neous in composition and Fe-rich (Fig. 6d; Table 4).

They are surrounded by coronas of symplectitic plag-

ioclase, quartz, hornblende (Fig. 6e). Plagioclase, bio-

tite, quartz, ilmenite and rutile inclusions are

distributed throughout garnet; ilmenite and rutile are

typically associated with titanite. Plagioclase in the

matrix averages 0.7 mm in size, is commonly twinned,

and exhibits reverse zoning. Plagioclase occurring in

the symplectitic regions around the garnet is signifi-

cantly more Ca-rich, with XAn = 0.75–0.80, compared

to matrix plagioclase (XAn = 0.43–0.48) (Table 4).

Quartz is subhedral and commonly contains subgrains.

4. Pressure–temperature conditions

Pressures and temperatures of metamorphism are

constrained by thermobarometry and inferences from

petrogenetic grids. These data are combined with

interpretation of reaction textures to reconstruct seg-

ments of the P–T paths. Mineral equilibria were

calculated using the internally consistent thermody-

namic database of Berman (1991, updated 1997;

Table 4

Representative mineral compositions for garnet–hornblende rocks

Garnet Hornblende Plagioclase (matrix) Plagioclase (symplectite) Biotite Ilmenite

SiO2 38.65 42.82 56.57 49.56 36.75 –

TiO2 0.06 1.03 – – 4.12 51.98

Al2O3 22.01 13.56 28.03 32.27 14.94 0.06

FeO 25.19 19.05 0.14 0.40 20.55 44.58

MnO 1.56 0.39 – – 0.13 0.91

MgO 3.72 8.90 – – 11.22 1.37

CaO 9.85 11.29 9.50 14.63 < d.l. 0.07

Na2O – 1.41 6.07 3.04 0.27 –

K2O – 0.87 0.17 0.05 9.42 –

Total 101.04 99.31 100.48 99.96 97.39 98.97

Cations per 12 O 23 O 8 O 8 O 22 O 6 O

Si 3.00 6.37 2.53 2.26 5.52 –

Al 2.01 1.48 1.74 0.00

AlIV 1.63 2.48

AlV1 0.74 0.17

Ti 0.00 0.12 – – 0.47 1.98

Fe2 + 1.63 2.34 0.01 0.02 2.58 1.89

Mn 0.10 0.05 – – 0.01 0.04

Mg 0.43 1.97 – – 2.51 0.10

Ca 0.82 1.80 0.46 0.72 0.00 0.00

Na – 0.41 0.53 0.27 0.08 –

K – 0.16 0.01 0.00 1.81 –

XMg – 0.45 – – 0.49 0.05

XAn – – 0.46 0.73 – –

XAb – – 0.53 0.27 – –

XOr – – 0.01 0.00 – –

XAlm 0.55

XSps 0.03

XPrp 0.14

XGrs 0.27


TWQ version 2.02) and Thor-Odin mineral composi-

tions, with other thermobarometers used for amphib-

ole-bearing assemblages as described below.


In the garnet-bearing gedrite–cordierite rocks, tem-

peratures were estimated using the garnet–cordierite

Fe–Mg exchange thermometer. Maximum calculated

temperatures using garnet compositions from just

inside the retrograde rim paired with adjacent cordier-

ite (Table 1) are 725–800 jC (Fig. 7a). Quartz only

occurs as inclusions in staurolite and it is unlikely

that it was in equilibrium with the other phases. There-

fore, calculations of pressure were not possible due to

the lack of a barometer in the assemblage; however, a

higher pressure history is suggested by the existence

of the texturally early assemblage containing kyan-

ite + garnet + rutile which constrains pressure to >9

kbar at f 750 jC. If the protolith of this rock was

mafic, the occurrence of staurolite as inclusions in

garnet suggests pressure >5.5 kbar (Selverstone et al.,

1984; Arnold et al., 2000).


Temperatures of 725–850 jC were estimated for

the metapelitic rocks (Fig. 7a) using the garnet–

biotite Fe–Mg exchange thermometer with rim garnet

compositions paired with adjacent biotite composi-

tions. The absence of orthopyroxene in the Thor-Odin

rocks suggests that the higher calculated temperatures

(>800 jC) are not geologically real. Pressures of 8–

10 kbar were determined using rim garnet and adja-

cent plagioclase compositions and garnet–plagio-

clase–sillimanite–quartz barometry.


Temperatures in garnet–hornblende amphibolites

were estimated using exchange thermometers: gar-

net–hornblende (Graham and Powell, 1984), horn-

Fig. 7. (a) P–T diagram for rocks in the lower unit of the Thor-Odin

dome. Recorded pressures and temperatures for the lower unit are

725–850 jC and 8–10 kbar. Calculated equilibria for the observed

textures are also shown. (b) P–T diagram showing the locations of

selected dehydration-melting reactions.


blende–plagioclase (Holland and Blundy, 1994), and

garnet–biotite. Garnet rim compositions paired with

adjacent mineral compositions for all thermometers

yield similar results. Pressure estimates were deter-

mined using the garnet–plagioclase–hornblende–

quartz (Kohn and Spear, 1990; Fe-tschermakite

model) and garnet–plagioclase– rutile – ilmenite–

quartz (GRIPS) barometers. The estimated pressures

vary depending on the textural relationship among the

phases. The box in Fig. 7a for sample 99-144 (P= 8–

10 kbar) represents data from garnet rim compositions

that were paired with adjacent hornblende and plagio-

clase where the symplectitic corona was not well

developed. However, in the same sample, garnet rim

compositions paired with adjacent plagioclase and

hornblende in the symplectitic corona yield lower

pressures (reaction 6, Fig. 7a); these results are dis-

cussed in the next section. In sample 97-14, where

rutile, ilmenite, plagioclase and quartz are found as

inclusions in garnet, pressures determined using the

GRIPS barometer with inclusion compositions paired

with adjacent garnet compositions yield much higher

pressures (10 kbar at 750 jC, Fig. 7a) than those

estimated using garnet–plagioclase–hornblende–

quartz barometry (97-14 box, Fig. 7a).

5. Reactions and P–T paths

The above section reports pressure–temperature

conditions recorded by various mineral equilibria in

the rocks from the Thor-Odin dome. Each of these

calculated equilibria determines a point on the P–T

path experienced by these rocks. In this section,

additional equilibria were calculated, and constraints

on the shape of the P–T path were determined by the

interpretation of the sequence of equilibria necessary

to explain the observed textures. Equilibria were

calculated using mineral compositions and TWQ

software (Berman, 1991, updated 1997; TWQ version

2.02), except for the hornblende-bearing equilibrium

(6), which uses the calibration of Kohn and Spear

(1990). We also used THERMOCALC software

(Powell and Holland, 1988; version 3.1, with the data

set of Holland and Powell, 1998) to calculate equi-

libria in the MASH system to characterize the role of

gedrite and garnet in producing the pseudomorph

textures.


We have calculated stable equilibria for the break-

down of Al2SiO5 to produce the phases observed in

the symplectitic regions

garnetþ Al2SiO5 þ H2O ¼ cordieriteþ corundum

ð1Þgarnetþ corundumþ H2O ¼ cordieriteþ spinel

ð2ÞThe cordierite phase in all of the calculated equi-

libria is hydrous, as the presence of a volatile phase in

cordierite is indicated by microprobe analyses.

According to the calculated petrogenetic grid, equili-

bria (1) and (2) would be encountered sequentially

during decompression, after the rocks have moved

from the kyanite to the sillimanite stability field (Fig.

7a). This sequence of reaction implies that the original

mineral (kyanite or other tabular aluminous mineral)

was replaced from the core outward in the pseudo-

morphed grain shown in Fig. 5c,d. Petrographic obser-

vation of partially replaced phases, however, shows

that replacement is highly irregular, and therefore the

reaction sequence in relation to core vs. rim may have

depended on the presence and location of fractures.

The equilibria plotted in Fig. 7a were calculated

using TWQ, and do not include equilibria involving

gedrite because the activity–composition relations of

amphiboles are not well known. The equilibrium

kyaniteþ gedriteþ H2O ¼ cordieriteþ corundum

(not plotted) could also account for the growth of

cordierite and corundum at the expense of kyanite,

and would be encountered at similar P–T conditions

as equilibrium (1).

The presence or absence of sapphirine is deter-

mined by local variations in bulk composition, with

sapphirine occurring in slightly more Mg-rich rocks.

In the more Mg-rich rocks, sapphirine may have been

produced by the equilibrium

gedriteþ Al2SiO5 þ spinel ¼ sapphirine þ H2O

Alternatively, sapphirine may have been produced by

the equilibrium

garnetþsillimaniteþH2O¼ sapphirineþcordierite

ð3Þ


if garnet was originally present in the sapphirine-

bearing rocks (Fig. 7a). This equilibrium was deter-

mined using garnet compositions from the garnet-

bearing gedrite–cordierite rock as an approximation,

along with cordierite compositions from the sapphir-

ine-bearing rock. The equilibrium, as calculated, is

metastable. In order to calculate this equilibrium, we

added the phase sapphirine using thermodynamic data

from experiments (Chatterjee and Schreyer, 1972;

Hensen, 1972; Doroshev and Malinovskiy, 1974;

Seifert, 1974; Ackermand et al., 1975). Therefore,

the cause of this apparent metastability is likely the

lack of internally consistent thermodynamic data for

sapphirine in the Berman (1991) database.


The occurrence of cordierite in these rocks can be

explained by the equilibria

garnetþ sillimaniteþ H2O

¼ cordieriteþ spinelþ quartz ð4Þ

garnetþ sillimaniteþ quartz þ H2O ¼ cordierite

ð5Þ

These equilibria were calculated using garnet com-

positions associated with cordierite in the sillimanite–

garnet–K-feldspar gneiss. The flat slope of these

equilibria in P–T space suggests that they were

crossed during decompression (Fig. 7a).


Hornblende and plagioclase compositions from the

symplectitic coronas around garnet together with

garnet rim compositions were used to calculate the

equilibrium

garnetþH2O¼hornblendeþplagioclaseþquartz

ð6Þ

plotted in Fig. 7a. As discussed in the previous section,

mineral compositions within the symplectitic regions

record slightly lower pressures but similar temper-

atures. The higher pressures estimated by inclusion

thermobarometry, along with the flat slope of this

equilibrium, suggest nearly isothermal decompression.

5.4. Partial melting

The recorded P–T conditions are at or above dehy-

dration-melting reactions for muscovite and/or biotite-

bearing metapelitic rocks

muscoviteþ plagioclaseþ quartz

¼ K � feldspar þ sillimaniteþ biotiteþ L

ðMelt1Þ

biotiteþ albiteþ sillimaniteþ quartz

¼ garnetþ K � feldspar þ L ðMelt2Þ

calculated from experimental data (Fig. 7b) (Melt 1:

Patino Douce and Harris, 1998; Melt 2: Vielzeuf and

Holloway, 1988; Stevens et al., 1997). The dehydra-

tion-melting reaction

biotiteþ plagioclaseþ quartz

¼ orthopyroxene þ K � feldspar þ L ðMelt3Þ

puts an upper limit on the temperature reached (Fig.

7b) (Carrington and Harley, 1995; Spear et al., 1999),

as the Thor-Odin dome rocks lack orthopyroxene.

Some amount of decompression from the maximum

recorded P–T conditions in the metapelitic gneisses is

necessary for the dehydration-melting reaction to

proceed

biotiteþ sillimanite ¼ garnet þ cordieriteþ L

ðMelt4Þ

(Fig. 7b) (Carrington and Harley, 1995; Spear et al.,

1999). This melting reaction may explain the occur-

rence of regions of polygonal cordierite grains in the

sillimanite–garnet–K-feldspar gneiss.

6. Discussion

6.1. Metamorphism and decompression

The lower unit of the Thor-Odin region was

metamorphosed at high temperatures (700–800 jC)and experienced nearly isothermal decompression

from f 10 to 4–5 kbar. Evidence for decompression

is abundant in the lower unit. Symplectitic textures


are ubiquitous and texturally late cordierite is found in

all of the aluminous rocks. In the gedrite–cordierite

rocks, decompression is indicated by the occurrence

of a relict high-pressure assemblage (garnet + kyani-

te + rutile) which has reacted at high temperatures to

form a cordierite + spinel + corundumF sapphirine +

sillimanite assemblage (Fig. 5). Cordierite reaction

rims around mafic boudin phases occur where these

minerals have come into contact with quartzofeld-

spathic material. Leucocratic melt also accumulated

within the boudin necks. These data and observations

suggest that the gedrite–cordierite boudins in the

Thor-Odin dome interacted with the leucocratic melts

during decompression at high temperature.

Evidence for decompression is also found in both

the sillimanite–garnet–K-feldspar gneiss and garnet–

hornblende amphibolite. In the sillimanite–garnet–K-

feldspar gneiss, cordierite occurs around garnet, with

spinel in symplectitic regions around sillimanite and

relict kyanite, and in the matrix (Fig. 6). Finally, the

garnet–hornblende amphibolites contain garnets with

symplectitic coronas of plagioclase + hornblende +

quartz (Fig. 6). Previous workers have suggested that

fine-grained reaction products, coronas, partial re-

placement along the rim of grains, and symplectitic

textures similar to the ones that we observe in these

rocks are indicative of reactions failing to go to com-

pletion (e.g., Droop and Bucher-Nurminen, 1984).

It is likely that these reactions proceeded during

rapid changes in pressure– temperature conditions.

The P–T results, abundance of cordierite in these

rocks, and reaction textures are evidence for the

sequential crossing of a series of relatively pressure-

sensitive equilibria. These observations suggest that

the Thor-Odin dome underwent nearly isothermal

decompression.

Similar mineral assemblages and reaction textures

with this P–T path are found in high-temperature

metamorphic terrains around the world (e.g., Mohan

and Windley, 1993; Brown and Raith, 1996; Ouze-

gane et al., 1996; Davidson et al., 1997). The high

temperatures reached in the Thor-Odin dome are

consistent with reported temperatures and pressures

(820F 30 jC, 8F 1 kbar) for metapelitic rocks in the

Valhalla region in the southern Shuswap complex

(Fig. 2a) (Spear and Parrish, 1996). The Valhalla

rocks also record a fast cooling path of f 25 jC/Ma. However, metapelitic Valhalla rocks do not con-

tain cordierite and therefore probably record a differ-

ent P–T path.

6.2. Timing

The timing of crustal anatexis and the significance

of the ductile fabric in the lower unit of the Thor-Odin

dome are debated. Earlier workers have interpreted U–

Pb Precambrian ages in the lower unit as evidence that

the dominant deformation and melting event was pre-

Cordilleran and that the lower structural levels have

not been significantly affected by Cordilleran meta-

morphism and deformation (Armstrong et al., 1991;

Parkinson, 1991; Parrish, 1995; Crowley et al., 2001).

It is therefore possible that some of the rocks discussed

in the present study experienced Precambrian meta-

morphism.

There is evidence, however, that a significant

regional metamorphic episode occurred in the Terti-

ary, and that this is the timing of major partial melting

and decompression. The dominant ductile, extensional

fabric in the lower unit is consistent with the fabric in

the overlying Paleozoic cover found in the middle

unit. If the observed symplectitic textures developed

during the Precambrian, it is unlikely that they would

be preserved through a Cordilleran period of burial

and reheating. In addition, reaction textures indicate

that leucocratic melt interacted with minerals in the

mafic boudins during decompression. The outer rims

of zircon in leucosomes from the study area as well as

zircons from leucogranites from higher structural

levels yield early Tertiary U–Pb SHRIMP ages, sug-

gesting that the dominant preserved melting and

deformation event is Cordilleran (Vanderhaeghe et

al., 1999b) (Fig. 4). It is unlikely that decompression

could have occurred much prior to the extraction and

crystallization of these melts because the crust would

cool rapidly during decompression.

6.3. Partial melting in the crust

The cluster of U–Pb SHRIMP zircon ages around

55–60 Ma from the Ladybird leucogranite suite and

leucosomes from the lower unit migmatites suggests

that this represents a time when a significant volume

(perhaps 15–20%) of the crust was partially molten.

The existence of a partially molten mid-crust in a

thickened crust setting is supported by geophysical


observations, which average a large section of crust

from active orogens such as the Central Andes and the

Tibetan Plateau. In each of these cases, seismic and

conductivity studies reveal the existence of low seis-

mic velocity zones linked to high conductivity zones

which are interpreted to represent abundant partial

melt in the mid-crust (Nelson et al., 1996; Schilling

and Partzsch, 2001).

Partial melting of the Thor-Odin rocks may have

occurred both during burial and heating and during

decompression. Partial melting on the prograde path

likely occurred with the breakdown of muscovite and

biotite (Fig. 7b). However, the extent of melting

would be limited by the amount of water available

to drive water-saturated partial melting. The abun-

dance of leucogranitic material throughout the Thor-

Odin dome region requires the production of a large

amount of partial melt, with estimates of total melt

production >40 vol.% in some locations (Nyman

et al., 1995). The continuation of partial melting by

dehydration-melting reactions such as those shown in

Fig. 7b may explain the large volume of partial melt

that is observed. Furthermore, the occurrence of

magmatic andalusite in leucosomes in a middle unit

migmatite suggests that partial melting continued to

low pressures.

The occurrence of a large percentage of melt would

lead to a reduction of strength of the partially melted

layer (Arzi, 1978; Van der Molen and Paterson, 1979;

Vanderhaeghe and Teyssier, 2001). Structural and

thermochronologic data from the Thor-Odin dome

shows that a temporal link exists between partial

melting and exhumation of the migmatites (Vander-

haeghe and Teyssier, 1997; Vanderhaeghe et al.,

1999b). Therefore, it has been suggested that the

instability in the crust created by a partially molten

layer could be the driving force for rapid unroofing in

the Thor-Odin dome. Using petrologic, structural, and

thermochronologic data, we evaluate several pro-

cesses of unroofing for the Thor-Odin dome region.

6.4. Processes of unroofing

In the Thor-Odin dome, we have documented a

decompression path that indicates exhumation of at

least 15 km while the rocks remained hot, suggesting

that unroofing of these rocks was rapid. This is

consistent with the fast cooling rates determined from

thermochronologic studies (Vanderhaeghe, 1997;

Vanderhaeghe et al., 1999b; Lorencak et al., 2001).

The removal of upper crust may occur by erosion and

tectonic processes.

Although the exhumation of high-grade rocks in

metamorphic core complexes has not been attributed

traditionally to erosion, recent studies suggest that

focused erosion can account for rapid unroofing (e.g.,

Zeitler et al., 2001). If erosion were the dominant

process for removing the upper crust in the Thor-Odin

dome, it would necessarily be efficient and rapid to

explain the existing data. Estimates for the amount of

material from the Thor-Odin region deposited to the

east and west during the Cenozoic account for only a

few percent of that necessary to explain unroofing of

at least 15 km (Vanderhaeghe et al., 1999a). There-

fore, is it likely that unroofing was mostly accommo-

dated by tectonic processes, with erosion playing a

secondary role.

The temporal link between partial melting and

decompression suggests that tectonic unroofing of

the Thor-Odin dome may have been facilitated by

an instability created by partially molten crust. Un-

roofing of metamorphic core complexes has often

been explained by progressive movement along low-

angle detachment zones (Davis, 1980; Spencer and

Reynolds, 1989; Foster and Fanning, 1997). In the

Shuswap complex, the Ladybird leucogranite suite is

present at the level of the detachments. Hollister

(1993) suggested that strain localization in the crust

will occur when melt is present, and thus the initiation

of these detachment zones could be facilitated by the

presence of the leucogranitic melt at this level. Fur-

thermore, displacement along these high-strain zones

might be magnified by the presence of melt (Hollister

and Crawford, 1986), leading to more rapid unroofing

of high-grade rocks below the detachments.

Another process that results in decompression of

high-grade, partially molten rocks is diapirism. The

domal structure of the complex raises the question of

the role diapiric ascent may have played in the

observed decompression. The buoyant rise of diapirs

has been used to explain the formation of gneiss

domes in the French Massif Central (Schuilling,

1960; Faure and Cottereau, 1988), New England

(Thompson et al., 1968), North American Cordillera

(Howard, 1980), and Alaska (Calvert et al., 1999). In

this case, buoyant rise of the diapir accommodated by


return flow causes progressive and potentially rapid

decompression of partially molten rocks within the

diapir.

6.5. Model for unroofing of the Thor-Odin dome

In the lower unit rocks of the Thor-Odin dome, we

have documented P–T paths that indicate maximum

pressures of at least 8–10 kbar and high-temperature

decompression to 4–5 kbar. In contrast, pressure

estimates for the middle unit rocks in the Thor-Odin

region are slightly lower (6–8 kbar) and these rocks

do not contain widespread evidence for significant

high-temperature decompression. These data suggest

that rapid unroofing of at least 15 km of crust was not

homogeneous in the region, but rather localized in the

area of the Thor-Odin dome, which is located on the

eastern side of the complex. However, thermochrono-

logic data from the area suggest that the cooling

histories of the middle and lower units were similar

from f 550–125 jC, as documented by the similar-

ity of argon cooling ages and zircon fission-track ages

between the two units (Vanderhaeghe, 1997; Lorencak

et al., 2001). The difference in apatite fission-track

ages between the middle and lower units (45–50 and

40–45 Ma, respectively) suggests that there may have

been some differential unroofing at this time, possibly

accommodated by high-angle normal faults such as

the Victor Lake Fault (Fig. 2b).

The localized decompression and asymmetric loca-

tion of the dome might be explained in a detachment

model for unroofing by a greater amount of motion on

the eastern detachment, with denudation on the west-

ern side being more distributed and accommodated by

a series of high-angle normal faults. In this case,

progressive unroofing of rocks below the eastern

detachment might lead to a steep age gradient on that

side, with the youngest rocks occurring directly below

the exposed detachment. This expected age gradient is

not consistent with the existing thermochronologic

data.

A diapiric model for the lower unit might explain

the localized decompression in the dome. In this

model, the middle unit rocks remain relatively sta-

tionary in the crust as the lower unit rocks rise as a

diapir. However, this model does not explain the

homogeneous cooling history below the detach-

ments. In order to explain the petrologic and ther-

mochronologic data, a three-stage model for the

progressive unroofing of the Thor-Odin region is

proposed (Fig. 8).

In this three-stage model, the crust is nearly

doubled in thickness during accretion of terranes to

western North America (Coney and Harms, 1984;

Parrish et al., 1988) (Fig. 8a). Partial melting of the

fertile crust occurs and, with the accumulation of

sufficient melt, a lower to mid-crustal diapir devel-

ops, leading to rapid decompression of the lower unit

rocks (Fig. 8b). As pressure decreases, dehydration-

melting reactions are crossed, enhancing partial melt-

ing of the rocks within the diapir. The second stage of

unroofing occurs with the initiation of low-angle

detachment zones above the rising diapir (Fig. 8c).

The emplacement of leucogranitic melts derived from

the partially molten crust enhances movement along

the detachments. The continued production of melt

might occur as more portions of the crust cross

dehydration-melting reactions with decreasing pres-

sure. As the rocks below the detachments are

unroofed, mylonitic fabrics in the leucogranites are

developed, and rapid cooling occurs, with fluids

possibly playing a critical role in the rapid removal

of heat (Morrison and Anderson, 1998). Finally, a

third stage of unroofing of the lower unit rocks

occurs with movement along high-angle normal

faults on either side of the dome. The low-temper-

ature differential cooling of the lower unit in this

stage is recorded in the youngest apatite fission-track

ages.

6.6. Conclusion

The high-grade rocks in the Thor-Odin dome

record a significant history of high-temperature

decompression, followed by rapid cooling. Mineral

assemblages and reaction textures from rocks in the

deepest structural level of the Thor-Odin dome indi-

cate decompression from the kyanite zone (P>8–10

kbar) to the sillimanite–cordierite zone (P < 5 kbar) at

Tf 750 jC. Structural, geochronologic, and petro-

logic data show a link, both temporally and rheolog-

ically, between unroofing of these high-grade rocks

and the occurrence of partially molten crust. In a

thickened-crust setting, such as the Thor-Odin dome,

the driving force for rapid unroofing of the orogenic

core might be the instability that is created by the


Fig. 8. Conceptual model for tectonic unroofing of the Thor-Odin dome showing snapshots of the crustal section over time. Representative P–T

paths for the middle and lower unit rocks are shown for each stage of the model. (a) Crustal thickening results in the burial of fertile paleomargin

sediments and thickening of the crust. This portion of the P–T path for the middle unit and lower unit rocks is shown by the dashed arrows on

the P–T diagram. Heating during burial results in partial melting in the mid-crust. Crystallization of leucogranitic melt located at shallower

depths is recorded by U–Pb zircon ages of f 60–55 Ma. The respective positions of the upper unit (UU), middle unit (MU), and lower unit

(LU) are shown. (b) Accumulation of melt in the crust results in the diapiric rise of material in the mid- to lower crust and decompression of

migmatitic rocks in the lower unit. As migmatites rise towards the surface, crystallization of leucosomes occurs, recorded by U–Pb SHRIMP

ages of zircon rims at f 56 Ma. (c) Unroofing occurs by movement along low-angle detachment zones at the level of ponded leucogranite melt

and along high-angle normal faults. Thermochronologic data show that, at 49 Ma, the rocks that are at the surface today were below 300 jC. (d)In the final stage, movement along high-angle normal faults accommodates further unroofing of lower unit rocks.


positive feedback between decompression and partial

melting.

Acknowledgements

We thank B. Evans, J. Stout, C. Kopf, A. Fayon,

and J.-P. Burg for helpful discussions. This research is

supported by NSF grant EAR-9814669 to CT and

DLW. Constructive reviews by J. Cheney, E. Ghent,

and S. Paterson helped substantially to improve this

manuscript.

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partial melting and decompression of the thor-odin dome, shuswap metamorphic core complex, canadian...

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