Thermotectonic Evolution of the Wolverine

THERMOTECTONIC EVOLUTION OF THE WOLVERINE METAMORPHIC COMPLEX, BRITISH COLUMBIA: LIMITATIONS ON THE USE OF COMBINED ION EXCHANGE AND NET-TRANSFER REACTION GEOTHERMOBAROMETRY AT UPPER AMPHIBOLITE- FACIES METAMORPHISM

Reid Staples B.Sc. Simon Fraser University, 2007

THESIS SUBMITTED IN PARTIAL FULFILLMENT OF THE REQUIREMENTS FOR THE DEGREE OF

MASTER OF SCIENCE

In the Department of Earth Sciences

SIMON FRASER UNIVERSITY

Spring 2009

APPROVAL

Name: Reid Staples

Degree: Master of Science

Title of Thesis: Thermo-tectonicevolution of the Wolverine metamorphiccomplex, British Columbia:limitations on the useof combinedion exchangeand net-transfer reactiongeothermobarometry at upper amphibolite- faciesmetamorphism

ExaminingCommittee: Chair: Dr. Dirk Kirste AssistantProfessor, Department of EarthSciences

Dr. Dan Marshall SeniorSupervisor AssociateProfessor, Department of EarthSciences

Dr. Derek Thorkelson Supervisor Professor,Department of EarthSciences

Dr. Dan Gibson Supervisor AssistantProfessor, Department of EarthSciences

Dr. Bert Struik E,xternalExaminer GeologicalSurvey of Canada

Date Defended/Approved: March25'h. 2009

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AB STRACT

Peak metamorphism of the Wolverine metamorphic complex ( W M C ) occurred at conditions of 770°-830°C and 7.2-10.4 kbars, and was accompanied by partial melting and the development of tight to isoclinal northeast-vergent folds that are axial planar to a transposition foliation. The WMC is modelled as a diffuse northeast-vergent shear zone that formed beneath a southwest-vergent panel of rocks as the orogenic wedge detached and translated northeastward during Mesozoic contraction. Juxtaposition of Middle Jurassic greenschist-facies upper crustal rocks against upper amphibolite-facies rocks of the WMC, which contain Eocene 40Ar/39Ar cooling ages, suggests that the WMC remained at deep crustal levels until it was rapidly exhumed in the Eocene along the normal

Wolverine fault. Mineral reaction and disequilibrium textures indicate a near- isothermal decompression path from 7.2-10.4 kbars to below 4 kbars, corresponding to a minimum of 11 km of exhumation prior to cooling below

~650°C.

Keywords: Wolverine metamorphic complex; Geothermobarometry; P-T-t path; tectonic evolution; 40Ar/39Ar geochronology

Subject Terms: Metamorphism (geology); Earth temperature; Rock pressure; Rocks, metamorphic

iii

ACKNOWLEDGEMENTS

First, I would like to thank Dan Marshall for giving me the opportunity to work on the high-grade metamorphic rocks of the Wolverine complex. Thank you for all of your support, mentorship, intellectual insight, and perhaps most all giving me the freedom to think for myself. I am very thankful to Dan Gibson and

Derek Thorkelson for taking the time to make suggests that greatly improved this thesis, as well as their helpful and thought provoking discussions. I have also gained considerably from having had the opportunity to meet and learn from someone as knowledgeable as Bert Struik.

I am very grateful to have had such a wonderful lab mate as Karin Fecova, and am forever indebted to her patience, tolerance and willingness to help. I am very thankful to Kevin Cameron for the numerous discussions and for always providing me with a great lab to TA. Gabe, Sarah, Francesca, Liz, Rusty and

Michael have made my time at SFU very enjoyable.

I would like to thank Fil Ferri from the BCGS for providing me with thin- sections and field notes at a time when I had nothing. I am also very appreciative to Mati Raudsepp for his help with the microprobe.

I would like to thank my parents as well as my Nan for their support and for giving me the opportunity to pursue my studies. Finally, my deepest thanks go to Melissa Jackson for all of her love and support.

TABLE OF CONTENTS

Approval ...... ii Abstract ...... iii Acknowledgements ...... iv Table of Contents ...... v List of Figures ...... vi i List of Tables ...... ix 1: Introduction ...... 1 2: Geologic setting ...... 4 3: Lithology ...... 8 3.1 Metamorphic rocks ...... 8 Unit 1a: Biotite-to garnet-grade pelitic schists ...... 9 Unit 1b: Sillimanite- to K-feldspar + sillimanite-grade pelitic schists ...... 9 Unit 2: Calcsilicate gneiss ...... 11 Unit 3: Amphibolite gneiss ...... 11 3.2 Intrusive rocks ...... 12 4: Correlation ...... 16 5: Structure ...... 20 5.1 Pre- to syn-transposition structures ...... 20 5.2 Post-transposition structures ...... 25 6: Metamorphism and microstructure ...... 27 6.1 Petrography – mineralogy and microstructure ...... 27 Sample locations...... 27 Pelitic and semipelitic schist and gneiss ...... 27 Amphibolites ...... 32 6.2 Mineral chemistry ...... 33 Metapelites ...... 34 Amphibolites ...... 35 6.3 Partial melting and P-T constraints ...... 38 Metapelites ...... 38 Amphibolites ...... 40 6.4 Interpretation of observed textures and zoning ...... 45 Partial disequilibrium during diffusional relaxation of growth zoning in garnet ...... 45

Retrograde zoning ...... 50 Qualitative P-T path ...... 53 6.5 Thermobarometry ...... 62 Analytical techniques ...... 62 Near-thermal- and post-thermal-peak thermobarometry ...... 63 6.6 Cooling and decompression as inferred from post-thermal- peak thermobarometry vs. mineral reaction textures ...... 70 7: Geochronology ...... 74 7.1 Previous geochronology ...... 74 7.2 Sampling strategy ...... 75 7.3 40Ar/39Ar results ...... 77 Amphibole ...... 77 Biotite ...... 78 7.4 Cooling History ...... 80 7.5 Discussion ...... 82 8: Orogenic model ...... 86 8.1 Structural divergence ...... 86 8.2 Exhumation ...... 94 9: Conclusions ...... 98 Reference List ...... 101 Appendix 1: Sample locations ...... 112 Appendix 2: Values of diffusion parameters and associated uncertainties ...... 113 Appendix 3: Summary of near-thermal peak thermobarometric estimates and Representative compositions used for thermobarometry ...... 114 Appendix 4: 40A r / 39Ar analytical techniques and sample descriptions ...... 119 Sample RS-07-06B ...... 120 Sample RS-07-54 ...... 120 Sample RS-07-63D ...... 121 Sample RS-07-84C ...... 122

LIST OF FIGURES

Figure 1. Index map showing the location of the WMC within the Omineca Belt of the Canadian Cordillera ...... 6 Figure 2. Geologic map of the WMC; including garnet and sillimanite isograds, as well as 40Ar/39Ar and thermobarometry sample locations...... 7 Figure 3. Lithologic units from the WMC ...... 14 Figure 4. Photo of a single sillimanite crystal 12 cm in length within a garnet-sillimanite-biotite gneiss ...... 15 Figure 5. Physiographic map in the region of the WMC...... 19 Figure 6. Transposition foliation defined by compositional layering and alignment of biotite grains. Boudinaged leucosomes of concordant quartz-feldspar pegmatite ...... 20 Figure 7. Northeast-vergent tight to isoclinal folding within the metapelite and amphibolite gneisses ...... 22 Figure 8. C-S fabric within garnet grade micaceous schist indicating top to the northeast sense of shear ...... 24 Figure 9. Back-scattered electron (BSE) image of monazite from sample RS-07-64B a migmatitic garnet-sillimanite-biotite gneiss ...... 28 Figure 10. Photomicrographs of metapelites from the K-feldspar + sillimanite zone ...... 31 Figure 11. Photomicrographs of plagioclase-amphibole symplectic coronas around garnet from an amphibolite, sample 63C...... 33 Figure 12. Representative compositional zoning profiles of garnet from a metapelite gneiss and amphibolite gneiss of the WMC ...... 37 Figure 13. Petrogenetic grid for pelitic schists displaying important dehydration equilibria and solidus curves ...... 39 Figure 14. Amphibole-bearing pegmatitic leucosome within a migmatitic amphibolite ...... 41

Figure 15. P-T diagram displaying the aH2O-controlled dehydration reaction quartz + muscovite + albite = alkali feldspar + sillimanite + H2O, and the aH2O-controlled solidus curves of the melting

vii

reactions quartz + alkali feldspar + muscovite + H2O = melt, and quartz + alkali feldspar + sillimanite + H2O = melt...... 44 Figure 16. Schematic P-T diagram illustrating possible P-T path trajectories with reference to XGrs isopleths...... 55 Figure 17. Photomicrograph metapelite sample RS-07-92A displaying An- rich contents of plagioclase within textural locations interpreted to have crystallized from the reaction grossular + sillimanite + quartz = anorthite during retrogression ...... 57 Figure 18. Schematic P-T path diagram [modified from Spear (1989) and Spear et al. (1990)] illustrating the relationship between the equilibrium constant (Keq) for the reaction grossular + 2sillimanite + quartz = 3anorthite and isopleths for XGrs and XAn...... 59 Figure 19. P-T diagram illustrating a cooling and decompression path inferred from the absence of retrograde muscovite, and that is consistent with the observed change in composition of late plagioclase, and XGrs zoning at the garnet rim ...... 61 Figure 20. P-T diagram showing calculated near thermal peak and post thermal peak P-T conditions of the metapelites ...... 66 Figure 21. P-T diagram showing calculated near thermal peak temperature conditions of the amphibolites ...... 69 Figure 22. P-T diagram illustrating the hypothetical relationship between closure temperatures for garnet-biotite thermometry and garnet- Al2SiO5-quartz-plagioclase (GASP) barometry ...... 73 Figure 23. Compilation map of geochronological data and mapped isograds from the WMC ...... 76 Figure 24. 40Ar/39Ar apparent age spectra and inverse isochron plots for hornblende...... 79 Figure 25. 40Ar/39Ar apparent age spectra for biotite ...... 80 Figure 26. Diagram of the two principal models for the development of structural divergence within the hinterland of the Canadian Cordillera ...... 89 Figure 27. Conceptual orogenic model for the development and exposure of structural divergence at the latitudes of the WMC ...... 93 Figure 28. Distribution and orientation of Eocene extensional and dextral strike-slip faults in the region of the southern WMC ...... 96 Figure 29. Diagram of events within the Canadian Cordillera that are of interest to the development and exhumation of the WMC ...... 97

viii

LIST OF TABLES

Table 1. Summary of regional geochronology data from chapter 8 discussion...... 83 Table 2. Representative garnet compositions used for thermobarometry...... 115 Table 3. Representative biotite compositions used for thermobarometry...... 116 Table 4. Representative plagioclase compositions used for thermobarometry...... 117 Table 5. Representative amphibole compositions used for thermobarometry...... 118 Table 6. 40Ar/39Ar results for biotite and hornblende...... 123

ix 1: INTRODUCTION

The Wolverine metamorphic complex ( W M C ) is located within the metamorphic and plutonic hinterland of the Foreland fold and thrust belt of the

Canadian Cordillera (Monger et al., 1982). The hinterland is located within the

Omineca Belt of the Canadian Cordillera and is a zone of intense deformation, metamorphism and plutonism that is a consequence of the convergence, accretion and obduction of the Intermontane Superterrane onto the paleocontinental margin of North America (Monger et al., 1982). This zone of tectonic overlap between allochthonous accreted terranes and parautochthonous strata deposited on the western North American cratonal margin, is marked by a regional zone of structural divergence that extends most of the length of the

Omineca Belt (Price, 1986).

Previous mapping within the study area by Ferri and Melville (1994) identified the presence of sillimanite within penetratively, ductily deformed rocks within the WMC, with a northeast sense of vergence that is similar to the low- grade upper crustal levels within the hanging wall. Conspicuously absent within the exposed upper and lower crustal levels at the latitude of the study area, is the southwest-vergent deformation observed within middle crustal levels of the WMC to the north and south.

Two principal orogenic models have been proposed to explain the structural divergence within the Selkirk fan located along strike farther south in

the Selkirk Mountains. A model of “tectonic wedging” was first proposed by Price

(1986) and later modified by Colpron et al. (1998), whereby southwest-vergent structures developed above an eastward propagating wedge of distal North

American strata that was impeded by an inherited crustal ramp. Alternatively,

Brown et al. (1993) suggest that a southwest-vergent prowedge and northeast- vergent retrowedge developed above the point of detachment for an eastward subducting oceanic or marginal basin lithosphere beneath the western North

American craton; analogous to the doubly vergent orogenic model of Willet et al.

(1993).

The tectonic wedging model has been invoked twice at latitudes near the present study area. First, by Bellefontaine (1990), to explain the southwest- vergent structures in the garnet-grade rocks of the central Swannell Ranges to the north, and secondly, by Ferri and Melville (1994) who suggested that the absence of southwest-vergent structures in both upper and lower crustal levels at the latitude of the study area, reflects the area’s location within the indenting wedge of Price (1986).

The primary purpose of this study is to better constrain the P-T and metamorphic reaction history of the WMC, as a means of assessing the role of the WMC during orogenesis. An evaluation of P-T conditions is critical to an understanding of the processes that affected the middle to lower crust during orogenesis. The method of geothermobarometry utilizes the principles of equilibrium thermodynamics, and relies on the assumption that coexisting mineral phases are continuously and simultaneously chemically equilibrating

along the P-T path followed by the rock. Chemical equilibrium must be attained not only between coexisting phases, but also amongst all elements within an individual mineral, and at the same P-T point along the P-T path. Conventional geothermobarometry involves the simultaneous solution of a minimum of two independent end-member phase equilibria, using compositional data taken from coexisting phases presumed to have equilibrated at the same pressure and temperature. This study considers the potential of partial disequilibrium (i.e., disequilibrium between Mg, Fe, Mn and Ca) within garnet at upper amphibolite- facies metamorphism, and the implications this has on combined Fe-Mg exchange thermometry and Ca-end-member barometry.

Finally, linking the deformational history with metamorphic reactions, as well as determining the timing and rate of cooling within the complex, relative to upper crustal levels with a similar sense of structural vergence, are used to develop an orogenic model.

2: GEOLOGIC SETTING

The WMC lies within the Omineca Belt in north-central British Columbia

(Fig. 1), which is a region of tectonic overlap between the North American craton, including rocks deposited on the cratonal margin, and part of a late Paleozoic- early Mesozoic back arc basin, and isotopically juvenile accreted terranes. The complex occurs within the para-autochthonous Cassiar terrane, which consists of

Upper Proterozoic to Permian siliciclastic and carbonate miogeoclinal sediments of the Windermere Supergroup that were deposited along the ancient margin of

North America. The WMC is a fault-bounded, northwesterly-trending culmination of deformed amphibolite-facies rocks (Fig. 2).

To the west of the complex are sedimentary and volcanic rocks of the

Quesnel terrane, an isotopically juvenile terrane, and the Slide Mountain terrane, part of a late Paleozoic- early Mesozoic back arc basin, which together form part of the Intermontane Superterrane. In central British Columbia the Mississippian to Permian Slide Mountain terrane consists of thrust-bound fault slices of chert, phyllite, argillite and basalt assigned to the Nina creek Group, and the associated

Manson Lakes ultramafic rocks, thought to represent the floor of an ocean and/or marginal basin (Ferri and Melville, 1994). The Quesnel terrane rocks lie farther to the west, consisting of a sedimentary and volcanic suite assigned to the Middle

Triassic to Lower Jurassic Takla Group (Ferri and Melville, 1994), and a

sedimentary and volcanic suite of the upper Paleozoic Lay Range assemblage, which is part of the Harper Ranch Subterrane (Ferri and Melville, 1994).

In the region of the study area, the WMC is bounded on its west side by the Wolverine fault, a west side down brittle-ductile normal fault (Ferri and

Melville, 1994), which juxtaposes upper amphibolites-facies rocks within the complex against lower to middle greenschist-facies rocks of the Slide Mountain terrane. The Wolverine fault also marks the boundary between Early Tertiary K-

Ar ages east of the fault within the complex, with Middle Jurassic and older ages west of the fault. Northwest of the study area, displacement along the Wolverine fault apparently decreases, juxtaposing Windermere-type strata of similar metamorphic grade (Ferri and Melville, 1994).

To the east of the complex are Paleozoic miogeoclinal rocks of the

Foreland fold and thrust belt that are separated from the complex by the Northern

Rocky Mountain Trench, a regional dextral transcurrent fault (Gabrielse, 1985;

Gabrielse et al., 2006). The region is cut by northwest- and north-trending dextral strike-slip faults that were active during the Early to Middle Eocene, and the Late Eocene to Early Oligocene, respectively (Struik, 1993).

Figure 1. Index map showing the location of the WMC within the Omineca Belt of the Canadian Cordillera

Figure 2. Geologic map of the WMC; including garnet and sillimanite isograds, as well as 40A r / 39Ar and thermobarometry sample locations.

3: LITHOLOGY

T h e W M C consists of metamorphosed sedimentary and mafic to intermediate intrusive rocks that are intruded by syn- and post-metamorphic felsic sills and dikes. The metamorphic rocks are divided into three primary units based on composition. They separate into 1) metapelite, 2) calsilicate, and 3) amphibolite assemblages. Biotite to garnet grade metapelitic schist occurs at the highest structural and stratigraphic level. With increasing structural and stratigraphic depth, metamorphic grade and penetrative ductile deformation gradually increases and culminates in the migmatitic sillimanite + K-feldspar zone within the centre of the complex. A migmatitic amphibolite gneiss occurs at the lowest structural and stratigraphic level within the complex, with the transition from the metapelite to amphibolite units marked by an intercalation of calcsilicate gneiss with the these two units.

3.1 Metamorphic rocks

The metasedimentary units occur continuously from the west side of the complex at biotite-grade through to garnet-, sillimanite-, and K-feldspar + sillimanite-grade migmatitic schists in the core of the complex (Fig. 2). The continuum of increasing metamorphic grade, ductile deformation, tectonic mineral segregation and degree of partial melting toward the centre of the complex, precludes the separation of metapelitic units on a compositional basis.

Therefore, the metapelitic rocks are broadly subdivided into two end-member units of lowest biotite to garnet grade pelitic schists, and highest sillimanite + K- feldspar grade migmatitic gneiss.

Unit 1a: Biotite-to garnet-grade pelitic schists

The lowest grade metamorphic rocks are biotite- to garnet-grade schists that occur adjacent to the Wolverine fault in the northwest part of the study area.

This unit consists of intercalated layers of quartzo-feldspathic and micaceous schists that range between approximately 5 and 50 cm in thickness (Fig. 3A).

The quartzo-feldspathic schists contain an average modal mineral assemblage of

50-60% quartz, 15-20% plagioclase, 20-25% biotite, 0-10% muscovite, and 0-5% garnet. The micaceous schists have an average modal mineral assemblage that consists of 60% biotite, 25% muscovite, 10% quartz, 5% plagioclase, and 0-5% garnet. Garnet in the quartzo-feldspathic and micaceous schists occurs as 1-5 mm porphyroblasts that are often partially or completely replaced by chlorite and quartz.

Unit 1b: Sillimanite- to K-feldspar + sillimanite-grade pelitic schists

Moving from west to east, the sillimanite isograd is crossed before the second sillimanite isograd is crossed in the centre of the complex. The metapelitic schists contain the assemblage quartz, biotite, plagioclase, ± K- feldspar, ± sillimanite, ± garnet. Plagioclase and garnet form porphyroblasts up to 3 cm and 1.5 cm in diameter, respectively. A single sillimanite crystal 12 cm in length was observed in one locale (Fig. 4). The variation in mineral assemblage

within and between outcrops in the K-feldspar + sillimanite zone, such as the presence or absence of sillimanite and garnet, is attributed to compositional variations in the sedimentary protolith.

The addition of K-feldspar and the absence of muscovite is accompanied by a significantly increased volume of mineral segregation into quartz + plagioclase + K-feldspar leucosomes, and melanosomes of biotite + sillimanite + garnet that are concordant with the transposition foliation (Fig. 3B). This texture is interpreted as a migmatitic texture produced by in-situ partial melting by the reactions

muscovite + plagioclase + quartz = sillimanite + K-feldspar + melt, (1)

and

Biotite + sillimanite + quartz + plagioclase = garnet + K-feldspar + melt (2)

(Le Breton and Thompson, 1988). Selvages of biotite and sillimanite surrounding quartz + feldspar leucosomes (Fig. 3C) are interpreted as a restite and/or the products of garnet breakdown and crystallization of melt during cooling.

Following the descriptive terminology of Mehnert (1968) and Ashworth

(1985) for migmatites, the term leucosome is used in a non-genetic sense for light-coloured layers (dominantly feldspar and quartz); melanosome refers to dark-coloured layers (dominantly biotite or hornblende); and restite refers to the remnant portion of the rock from which a substantial amount of the more mobile components have been extracted without being replaced.

In addition to the concordant leucosomes, the increased metamorphic grade and ductile deformation in the centre of the complex, are accompanied by numerous discordant pegmatites and dioritic intrusions.

Unit 2: Calcsilicate gneiss

Calcsilicate gneiss occurs at a structural and stratigraphic level just above the amphibolite gneiss, and is interlayered with migmatitic pelitic schist. The calcsilicate gneiss consists of quartz, biotite, diopside, calcite, ± garnet, ± amphibole, ± epidote. A compositional banding that consists of quartz-, quartz- diopside, biotite-, and biotite-calcite-rich layers parallel to foliation, is likely inherited from compositional variations between layers within the protolith. There are numerous pegmatite and diorite dikes that cut the calcsilicate gneiss, and a diopside-rich rim commonly forms at the contact with the pegmatitic dikes.

Unit 3: Amphibolite gneiss

The amphibolite unit occupies the centre and the structurally and stratigraphically lowest exposed levels of the complex (Fig. 3D). The amphibolite is typically migmatitic and is interlayered with migmatitic pelitic schist, and lesser calcsilicate schist. Leucosomes in the amphibolite are coarse-grained to pegmatitic and consist of quartz, feldspar and amphibole. Melanosomes consist primarily of amphibole ± biotite ± garnet. There is an apparent correlation between increased leucosome (melt volume) with decreasing modal proportions of biotite. The leucosomes are coarser-grained than the surrounding mesosome and occur as mm- to cm-scale bands, pods and lenses that are concordant with

the foliation in the mesosome. Modal amphibole varies and may be up to 80% in places. Numerous discordant quartz and feldspar-bearing pegmatites and diorite dikes cut the amphibolite gneiss.

3.2 Intrusive rocks

The high-grade metasedimentary units contain numerous concordant and discordant pegmatites, and are cut by numerous leucocratic granitoid dikes and stocks. The volume of intrusive rock within the metasedimentary units increases with metamorphic grade, with little to no intrusions at garnet-grade, while comprising up to 50% of the exposed lithology within the K-feldspar + sillimanite zone.

Granitoid stocks and dikes, which are extensive within the core of the complex, range from granodiorite to quartz diorite in composition, with muscovite comprising up to 20% of the modal mineralogy when present. Concordant pegmatites range between several centimetres in width/thickness, whereas discordant pegmatites can be several meters across. The mineralogy of pegmatites is variable, with ubiquitous quartz and plagioclase, and K-feldspar, biotite and muscovite forming a major phase where present. Garnet and sillimanite may form accessory phases. The presence of sillimanite in many discordant pegmatites with a coarse-grained and irregular contact with the surrounding metasedimentary rocks, and amphibolites, suggests the pegmatites intruded into, and crystallized within, a hot country rock.

The chronological, or cross-cutting relationships between the different textural and compositional varieties, in addition to the relative timing of intrusion and crystallization with respect to deformation, is complex and ambiguous. The variation is attributed to a progression of anatectic melting at different crustal levels relative to the thermal peak of metamorphism, and the accumulating regional strain.

Concordant pegmatites and leucosomes are interpreted to be the result of in situ crustal melting. The concordant pegmatites often form the outline of isoclinal folds that are intrafolial with respect to a transposition fabric that developed during an early deformational event. Melt leucosomes surrounded by a restite of biotite and sillimanite outline isoclinal fold closures and indicates that much of the in situ anatectic melting was pre- to syn-kinematic. The concordant pegmatites are commonly highly boudinaged, with some to the point of complete dismemberment. Leucocratic granitoids of a similar composition cut both the metasedimentary and amphibolite units, and range from unfoliated to foliated parallel to the transposition foliation. These leucocratic intrusions are therefore both syn- and post-kinematic.

Figure 3. Lithologic units from the WMC. A) Intercalated quartzo-feldspathic and micaceous schist. Notebook for scale. B) Migmatitic garnet-sillimanite-biotite gneiss. C) Quartz-feldspar leucosomes surrounded by a biotite-sillimanite restite. D) Migmatitic amphibolite gneiss with concordant quartz-feldspar leucosomes cut b y discordant quartz-feldspar pegmatites and granodiorite dikes.

Figure 4. Photo of a single sillimanite crystal 12 cm in length within a garnet-sillimanite- biotite gneiss. Knife for scale.

4: CORRELATION

Structural complexities, such as extensive transposition and isoclinal folding, together with significant recrystallization and mineral segregation during metamorphism and partial melting, prohibit accurate correlation of the high-grade metamorphic units within the WMC. However, in the northwest portion of the study area, a small fault sliver within the Wolverine fault preserves a sequence of low-grade (chlorite to biotite grade), and less complexly deformed metasedimentary lithologies. These low-grade metasediments are similar to

Proterozoic metasediments between the Omineca and Osilinka rivers that have been traced by Ferri and Melville (1994) directly into lithologies underlying the

Swannell Ranges to the north, which are correlated with the Ingenika Group by

Mansy and Gabrielse (1978). The low-grade metasedimentary lithologies are interpreted to overly the higher-grade metasedimentary rocks that are exposed throughout most of the study area, due the position of the low-grade metasedimentary units at a higher structural position within the hanging wall of a splay of the Wolverine fault (Fig. 2).

The Ingenika Group was first proposed by Roots (1954) together with the

Tenakihi Group and the Wolverine Complex. Gabrielse (1975) found the division between Tenakihi and Ingenika Groups to be tenuous and referred to both as the

Ingenika Group. Mansy and Gabrielse (1978) have correlated the Ingenika

Group with the Kaza Group in the northern Cariboo Mountains, and the Miette

Group in the Rocky Mountains. These Groups form part of the Windermere

Supergroup, which extends the length of the Canadian Cordillera (Mansy and

Gabrielse, 1978).

Based on detailed mapping in the northern Omineca Mountains, Mansy and Gabrielse (1978) subdivided the Ingenika Group, from oldest to youngest, into the Swannell, Tsaydiz, Espee, and Stelkuz Formations. Phyllite and limestone preserved in a fault sliver within the Wolverine fault have been assigned to the Tsaydiz and Espee Formations respectively, as proposed by

Ferri and Melville (1994). The Swannell Formation is not exposed within the study area, but has been mapped by Ferri and Melville (1994), approximately 12 km northwest of this study area, where it is juxtaposed against itself across the

Wolverine fault. The biotite- to sillimanite-grade schist of the WMC may represent a metamorphosed equivalent of the sandstone and feldspathic wacke of the Swannell Formation as proposed by Ferri and Melville (1994); however, the calcsilicate and amphibolite gneiss at the structurally and stratigraphically lowest exposed levels of the complex, have no correlatives within the Swannell

Formation.

The stratigraphic sequence in the lowest exposed levels of the WMC is very similar to the stratigraphy exposed within the Sifton Antiform of the Sifton

Ranges, east of Spinel Lake (Evenchick, 1988). The sequence consists of metaquartzite, garnet-sillimanite schist, hornblende schist, paragneiss and marble that is overlain by pelitic schist thought to stratigraphically underlie, or to be equivalent with the Swannell Formation (Evenchick, 1988). The Sifton

Antiform is truncated by the contractional Sifton fault, across which 1.85 Ga crystalline basement is overlain by an extensive unit of metaquartzite and amphibolite about 1000 m thick (Evenchick, 1988). The stratigraphy in the hanging wall of the Sifton fault is generally similar to the Deserters Range

(Evenchick, 1988) where 728 Ma basement is nonconformably overlain by conglomerate, metaquartzite, amphibolite and schist of the Misinchinka Group; a

Windermere Supergroup correlative to the Ingenika group (Mansy and Gabrielse,

1978). The Deserters and Sifton Ranges (Fig. 5) represent two rare localities in the Canadian Cordillera where Windermere Supergroup is in nonconformable contact (Deserters Range) and probable unconformable contact (Sifton Range) with crystalline basement (Evenchick, 1988).

Despite the absence of exposed metaquartzite in the WMC, which may lie stratigraphically below the amphibolite unit in the subsurface, the presence of amphibolite with a similar overlying stratigraphy in the Sifton Ranges implies that the Windermere Supergroup may also directly overly crystalline basement in the

WMC. Evenchick (1988) concluded that the absence of pre-Windermere strata in the region suggests that Middle and possibly some Upper Proterozoic rocks were removed by uplift and erosion in the pre- and early Windermere time.

Figure 5. Physiographic map in the region of the WMC.

5: STRUCTURE

5.1 Pre- t o s y n -transposition structures

A single, morphologically similar foliation is present at all structural levels within the study area, with an average orientation of 140/60SW. Compositional layering, and the alignment and preferred orientation of micas and sillimanite define the foliation (Fig.6).

Figure 6. Transposition foliation defined by compositional layering and alignment of biotite grains. Boudinaged leucosomes of concordant quartz-feldspar pegmatite. The scale of the card is in cm’s.

At the north end of the WMC at Chase Mountain, Parrish (1976) identified a morphologically similar axial planar foliation generated by the transposition of bedding during northeast-vergent isoclinal folding. The extreme flattening of F1 closures was attributed by Parrish (1976) to additional tightening during F2 folding that was approximately coaxial with F1. Slightly farther south of Chase

Mountain in the Blackpine Lake area, where the metamorphic grade is slightly higher (sillimanite-grade), continued tightening, and/or rotation was so great that

Parrish (1976) was often unable to distinguish between F1 and F2 folding. The superposition of F1 and F2 is thought to have produced morphologically similar coplanar elements that define a final composite transposition foliation (ST)

(Williams, 1983; Tobisch and Paterson, 1988). The F1 and F2 folding described by Parrish (1976) cannot be discriminated in the study area; pre- to syn- transposition folding is recorded as rootless, tight to isoclinal folds that are intrafolial with respect to the transposition foliation, and are herein labelled FT (F1 folds of Ferri and Melville, 1994). Interpreted near-thermal peak metamorphic mineral assemblages, such as intergrowths of biotite and sillimanite, are axial planar to FT. In addition, anatectic leucosomes that formed at the thermal peak of metamorphism are aligned with the ST foliation and outline FT (Fig. 7). These features suggest that FT folds formed at conditions near the thermal peak of metamorphism. A small amount of these pre- to syn-transposition isoclinal folds were both observed and measured, but all indicate a northeast-vergence, consistent with measurements of equivalent folds made by Ferri and Melville

(1994) from the same area.

Figure 7. Northeast-vergent tight to isoclinal folding. (A) migmatitic garnet-sillimanite- biotite gneiss. Viewing NNE. (B) migmatitic amphibolite gneiss. Viewing SE. Person for scale.

A poorly understood map-scale feature is the apparent mirrored repetition of lithologies from west to east across the complex, specifically the calc-silicate gneiss horizon at the transition from a dominantly metapelitic to amphibolite-rich lithology (Fig. 2). The presence of a southwest-dipping layer-parallel foliation across the width of the complex may imply that the repetition of lithological units represents the inverted limb of an overturned anticline or syncline. P-T estimates taken along an east-west transect of the complex indicate that the core of this fold represents the deepest exposed portion of the crust, consistent with an anticlinal fold geometry. Alternatively, the few exposed outcrops with a reversal in foliation dip to the east, may indicate that the eastern repetition of lithologies does not represent an overturned limb, but instead reflects at least one upright anticline-syncline pair across the width of the complex that is hidden due to poor exposure.

Accompanying the increase in metamorphic grade with depth in the complex, is a downward increasing strain gradient. Garnet-grade micaceous schists, on the western most side of the complex at the highest structural and stratigraphic level exposed in the complex, exhibit a well developed C-S fabric observable at both outcrop and thin-section scale (Fig. 8) that indicate top-to-the- northeast sense of shear.

Figure 8. C-S fabric within garnet grade micaceous schist indicating top to the northeast sense of shear. (A) Outcrop. (B) Photomicrograph of C-S fabric within a chlorite and sericite-after-garnet – muscovite quartzo-feldspathic schist. Mineral abbreviations consistent with the abbreviations of Kretz (1983).

With increasing structural and stratigraphic depth within the complex, there is a marked increase in: strained and boudinaged pegmatite; northeast-vergent rootless isoclinal folds with a well developed axial-planar transposition foliation; and a well developed compositional banding due in part to increasing anatectic melting (migmatization). Fold asymmetry and sigmoidal shear strain fabrics on feldspars and pegmatitic lenses increase dramatically with depth and indicate top-to-the-northeast shear. It is uncertain whether the main boudinage-forming event occurred due to rotation of compositional layering into the extensional field of the incremental strain ellipsoid during northeast-vergent folding and transposition, or during possible layer parallel extensional unroofing of the WMC.

5.2 Post-transposition structures

The presence of a large open, upright fold is indicated by rare reversals in foliation dips from west to east across the complex. The approximate parallelism of metamorphic isograds with compositional layering and pre- to syn- transpositional structural fabrics indicates that the isograds quenched prior to the development of this large upright antiform. The highest metamorphic grade (K- feldspar + sillimanite zone) and lowest stratigraphic levels occur in the core of the antiform with lower metamorphic grade (garnet and biotite zones) and higher stratigraphic levels situated in its limbs.

The Wolverine fault was first proposed by Armstrong (1949) to explain the discordance of strata between the high-grade rocks within the complex, and the low-grade rocks of the Slide Mountain terrane to the west. It has since been further constrained and described by Ferri and Melville (1994) to explain the

changes in structural trends and metamorphic grade across it, as well as the juxtaposition of early Tertiary K-Ar cooling ages within the high-grade metamorphic rocks of the WMC against Middle Jurassic ages in the much lower- grade rocks to the west. Within the study area, displacement along the

Wolverine fault appears quite significant, having juxtaposed sillimanite-grade rocks within the complex against chlorite-grade rocks of the Slide Mountain terrane.

Mapping within the poorly exposed region along the projected trace of the fault identified only brittle and brittle-plastic deformation features such as slickensided fault surfaces and pseudotachylite. However, slickenlines provide only the most recent motion on the fault in the uppermost brittle part of the crust, and cannot be relied upon for sense of shear during the earlier history of the fault.

6: METAMORPHISM AND MICROSTRUCTURE

6.1 Petrography – mineralogy and microstructure

Sample locations

Eight samples were selected for detailed petrography, microstructural and thermobarometric analysis from a representative regional sampling of lithologies from the WMC. The samples were selected on the basis of preservation of equilibrium mineral assemblages useful for thermobarometry, and lack of pervasive, low-grade retrogression of garnet, biotite and hornblende. Three amphibolite (samples RS-07-63C, RS-07-63D & RS-07-84C) and five metapelites (samples RS-07-02B, RS-07-28, RS-07-64B, RS-07-78 & RS-07-

92A) were chosen from varying structural and stratigraphic levels across the complex (Fig. 2, Appendix 1).

Pelitic and semipelitic schist and gneiss

The pelitic to semipelitic schist and gneiss have been metamorphosed to sillimanite + K-feldspar grade. Most have a mineral assemblage of quartz, plagioclase, K-feldspar, biotite, sillimanite and garnet. Accessory minerals include zircon and monazite (Fig. 9).

Figure 9. Back-scattered electron (BSE) image of monazite from sample RS-07-64B a migmatitic garnet-sillimanite-biotite gneiss. (A) BSE image of monazite within the matrix of biotite, K-feldspar and quartz. (B) A magnified BSE image of the monazite grain. Note at least four domains within the monazite, which likely represents successive growth and resorbtion during metamorphism.

Garnet occurs as large (3-10 mm), subhedral to anhedral porphyroblasts with numerous embayments filled primarily by quartz, and lesser amounts of biotite, sillimanite and rare plagioclase (Fig. 10A,B). The matrix consists of sillimanite, biotite, K-feldspar, plagioclase and quartz. Fibrolitic sillimanite is intergrown with biotite to form melanocratic bands that are deflected around garnet porphyroblasts. Prismatic sillimanite has been microboudinaged, displaying partings perpendicular to the plane of foliation. Both sillimanites occur along the rims and within embayments and pressure shadows of garnet

(Fig.10A,C). Biotite occurs as red-brown euhedral grains with the coarsest variety within the matrix. Finer grained biotite occurs along the rims of garnets

and within embayments of garnet (Fig. 10C). Biotite that has grown within pressure shadows of garnet is euhedral, coarse-grained and in some samples has been retrograded to chlorite along the portion of the grain in contact with garnet. K-feldspar, plagioclase and quartz occur throughout the matrix together with biotite, and as leucocratic bands without biotite.

The occurrence of selvages adjacent to leucocratic bands, and embayments, pressure shadows and reactions rims around garnet that consist of an assemblage of biotite, sillimanite, quartz and plagioclase, either singularly or as an assemblage, suggests operation of the retrograde net-transfer reaction

(reNTR)

garnet + K-feldpsar +melt = biotite + sillimanite + plagioclase + quartz. (2r)

ReNTR’s result in a change in the modal amounts of the phases involved. In reaction (2r) the modal amount of garnet is reduced, and this is evident from its highly resorbed texture with numerous embayments filled with biotite, sillimanite and quartz (Fig. 3A,B,C). Porphyroblasts of garnet are typically free of prograde, overgrown inclusions, which suggests that temperatures were high enough throughout the growth history such that components either in excess or not required for garnet were able to diffuse away (Passchier and Trouw, 2005; and references therein). Where inclusions are present, it is commonly biotite, quartz or fibrolitic sillimanite. In rare cases, garnet contains an inclusion rich core with the inclusions at random orientations. In other samples the inclusions form an internal foliation (Si) that is planar and at an angle to the external matrix foliation

(Se). However, the relationship of Si to Se is masked by an inclusion-free rim.

Following the reasoning above, the restriction of inclusions to within the garnet core would suggest that earlier (core) garnet growth occurred at lower temperatures along the prograde path, when temperatures were low enough such that diffusion was incapable of removing excess or unrequired components

(inclusions) within the garnet. With continued heating and garnet growth, temperatures became high-enough that diffusion of components away from the growing garnet was efficient enough to produce an inclusion free rim.

A composite transposition foliation (ST) is defined by the preferred orientation of biotite, mineralogical layering of biotite-sillimanite mesosomes and quartz-plagioclase-K-feldspar leucosomes. Peak thermal metamorphic minerals biotite, sillimanite and K-feldspar are aligned in the ST foliation within the matrix, and are deflected around garnet porphyroblasts (Fig. 10D). The deflection of peak metamorphic minerals around garnet indicates that garnet growth was pre-, or syn-kinematic. Intragranular microcracks in garnet perpendicular to the ST foliation (Fig. 10D), and deflection of ST around garnet indicate that compressional deformation and ST formation outlasted garnet growth. The occurrence of some sillimanite, biotite and plagioclase as products of garnet breakdown (see above) that are aligned within ST, suggest that the development of ST not only outlasted garnet growth, but also the thermal peak of metamorphism. The pre- to syn-kinematic nature of metamorphism is consistent with a lack of mineral growth cross-cutting the ST foliation.

Figure 10. Photomicrographs of metapelites from the K-feldspar + sillimanite zone. (a) Sample 92A displays biotite + sillimanite + quartz within embayments in garnet (crossed-polarized light). (b) Sample 02B displays biotite within the pressure shadow and embayments in garnet. Biotite is aligned parallel to ST (crossed-polarized light). (c) Sample 92A displays sillimanite within the pressure shadow of garnet. Sillimanite is aligned parallel to ST (crossed- polarized light). (d) Sample 64B displays fractures in garnet perpendicular to ST (plane-polarized light).

Amphibolites

The common mineral assemblage in the amphibolites is hornblende + plagioclase ± biotite ± quartz ± garnet (Fig. 11). No pyroxene was found in the amphibolites. Amphibole occurs both as subhedral grains (0.5-3 mm) that vary modally from 10 to 70% and as fine subhedral to euhedral grains intergrown with plagioclase to form symplectite coronas around embayed garnet (see sample

63C, Fig. 11). In some cases, plagioclase + amphibole symplectites have completely replaced garnet. Garnet when present, forms anhedral (xenoblastic)

(0.5-4 mm) grains that are typically free of inclusions. When inclusions are present in garnet, no internal foliation or preferred orientation is evident, as such, a relationship with the external foliation cannot be determined. The preferred orientation of hornblende and biotite, together with a compositional layering of melanocratic (amphibole-biotite rich) and leucocratic (plagioclase-rich) layers, define the ST foliation. Plagioclase grains are commonly twinned and vary modally from 10 to 60%. Biotite (0.5-4.0 mm) varies modally between 10 and

40%. Quartz comprises less than 10% of the mode.

Figure 11. Photomicrographs of plagioclase-amphibole symplectic coronas around garnet from an amphibolite, sample 63C. Symplectic plagioclase and amphibole pseudomorphing garnet (A) plane-polarized light, (B), cross-polarized light. Symplectic plagioclase and amphibole corona around garnet (C) plane- polarized light (D) cross-polarized light.

6.2 Mineral chemistry

Within each thin-section, mineral compositions from two garnet domains were quantitatively analyzed using a Cameca SX 50 electron microprobe at the

University of British Columbia. Analyses were run with a 20 kV accelerating voltage and a 10 nA beam focused over a 3 μm beam spot for a dwell time on the order of 10 seconds for most elements and longer dwell times for F and Cl.

Quantitative data for garnet, biotite and plagioclase was obtained for the pelitic gneiss and garnet, biotite, plagioclase and amphibole for the amphibolite gneiss.

Metapelites

Mineral abbreviations are consistent with the abbreviations from Kretz

(1983). Garnet in the metapelites is almandine-rich and symmetrically zoned from core to rim in Mg, Fe, Mn and Ca (Fig. 12A). The mole fraction spessartine,

Xsps, is low, has a flat profile in the core and rises abruptly at the rim. Mole fraction grossular, XGrs, displays a subtle zonation that increasing outwards from the garnet core to a near-rim maximum, before falling abruptly to a low at the garnet rim. The Fe/(Fe+Mg) ratio has a flat, symmetrical profile from the core towards the rim and rises sharply at the rim (Fig. 12A). Fe/(Fe+Mg) values from the five metapelite samples in the K-feldspar + sillimanite zone are as low as

(0.77 – 0.82) in the core of the garnet and up to (0.81 0.90) at the rim.

Fe/(Fe+Mg) zoning at the rim of garnet is not restricted to regions of the garnet adjacent to biotite, but also adjacent to non-ferromagnesian minerals such as sillimanite, plagioclase and quartz. Garnet is also zoned in Fe/(Fe+Mg) around biotite inclusions.

Biotite grains in contact with the garnet rim, either in pressure shadows, reaction rims or embayments in garnet, have the highest Fe/(Fe+Mg) (0.46 –

0.65) and the lowest TiO2 contents. Biotite inclusions in garnet that are isolated from the matrix, have the lowest Fe/(Fe+Mg) values (0.45 – 0.638). Matrix biotites have Fe/(Fe+Mg) values intermediate (0.475 – 0.643) to biotite inclusions

and biotite grains in contact with the garnet rim, and have the highest TiO2 contents.

Matrix plagioclase is nearly uniform in composition, varying < 0.03 wt% within single grains, and has the lowest XAn (varying from 0.31 – 0.69 between samples). Plagioclase grains in pressure shadows, reaction rims and embayments in garnet, have XAn contents 8 – 38% higher than the respective matrix grains.

Amphibolites

Garnet in the amphibolite unit is grossular-rich, relatively spessartine and pyrope poor almandines that are compositionally homogenous from core to rim.

The average composition is Alm0.527Grs0.257Pyr0.126Sps0.090. Garnet in amphibolites are typically unzoned from core to rim (Fig. 12B). This flat chemical profile across the width of garnet grains is interpreted to reflect chemical homogenization at peak metamorphic temperatures, similar to garnet from the metapelites. Garnet in sample RS-07-63C has plagioclase-amphibole symplectic coronas.

Matrix amphibole has no systematic compositional variation from core to rim; however, considerable compositional variation does occur between matrix and symplectic amphibole that forms a corona around garnet. Matrix amphibole is higher in K, Ca, Ti and Al and lower in Mn and Si than symplectic amphibole.

The composition of the matrix amphibole is ferropargasite to hastingsite, whereas symplectic amphibole is ferrotschermakite, following the nomenclature of Leake

(1997). An estimate of Fe3+ content in amphiboles was calculated following the procedure of Schumacher in appendix 2 of Leake (1997). Biotite has no apparent compositional variation between various textural locations, i.e. matrix biotite vs. biotite at the garnet rim, consistent with garnets, which are compositionally homogenous from core to rim, with no Fe/(Fe+Mg) zoning adjacent to biotite at the rim or included within garnet. Similar to the amphibole, the composition of plagioclase varies between matrix and symplectic plagioclase. The XAn for matrix plagioclase has a lower range (0.57 - 0.80) compared to the more anorthitic symplectic plagioclase (0.805 – 0.904).

Figure 12. Representative compositional zoning profiles of garnet from pelitic schists and amphibolite gneiss of the WMC determined from a microprobe traverse. (a) Zoning profile from sample RS-07-64B. A pelitic schist from within the K- feldspar + sillimanite zone. (b) Zoning profile from sample RS-07-63D an amphibolite gneiss from within the K-feldspar + sillimanite zone.

6.3 Partial melting and P-T constraints

Metapelites

The metapelites within the sillimanite + K-feldspar zone of the WMC display evidence of in-situ melting. Thus, peak temperatures are best constrained by dehydration (vapour-absent) melting reactions, which restrict peak metamorphic conditions to 650 - 860 °C and 5.2 – 11.3 kbars (Fig. 13).

Thompson and Connolly (1995) have suggested that the porosity of lower crustal rocks close to the solidus is much less than 1%, and the amount of melt that may be generated at the H2O-saturated solidus from free H2O trapped on grain boundaries is usually less than 3%. The near ubiquitous textural and chemical evidence in samples from metapelitic gneiss for back reaction of the biotite dehydration reaction, and the high volume of leucosome (up to 50%) in some migmatites, suggests that dehydration melting has occurred.

The lack of both a geochemical analysis of the metapelites, as well as a phase equilibrium program containing the thermodynamics for a melt phase, prohibit an isochemical, or “pseudosection”, placement of the dehydration melting reactions unique to the composition of the Wolverine samples. The placement of dehydration melting reactions were taken from experimental and thermodynamically determined reactions of Huang and Wyllie (1973, 1974),

Spear and Parrish (1996) and Spear et al. (1999) with compositions similar to, or appropriate to, the Wolverine metapelites.

Figure 13. Petrogenetic grid for pelitic schists displaying important dehydration equilibria and solidus curves after Huang and Wyllie (1973;1974); Le Breton and Thompson (1988); Spear and Parrish (1996); and Spear et al. (1999). The colored region represents the restricted P-T conditions based on the peak stable mineral assemblage for Wolverine pelitic schists.

The absence of muscovite and the presence K-feldspar and in situ melt in the metapelitic gneiss restricts P-T conditions to the high temperature side of the muscovite dehydration (vapour-absent) melting reaction (1). The presence of in

situ melt, and the textural and chemical evidence for the retrograde breakdown of garnet to form biotite + sillimanite + plagioclase + quartz, indicates that near- thermal peak P-T conditions must lie to the high temperature side of reaction (2).

The absence of orthopyroxene indicates that conditions were on the low temperature side of the vapour-absent melting reaction

biotite + plagioclase + quartz = orthopyroxene + K-feldspar + melt. (3)

The absence of kyanite and the presence of sillimanite put an upper limit on pressure, and the absence of cordierite puts a lower limit on pressure. It is possible that the absence of cordierite within the metapelites is due solely to a bulk composition that is not appropriate for the stabilization of cordierite. From the absence of muscovite, orthopyroxene and cordierite, and the presence of K- feldspar, sillimanite and in situ melt, peak P-T conditions have been restricted to between 650 - 860 °C and 5.2 – 11.3 kbars (Fig. 13).

Amphibolites

Leucosome in the amphibolite gneiss consists of amphibole-bearing pegmatite that form bands and lenses that are gradational at their margins and concordant to the host amphibolite gneiss (Fig. 14). To explain the presence of amphibole in migmatitic leucosomes, Kenah and Hollister (1983) and Viruete

(1999) have invoked vapour-absent biotite incongruent melting of a biotite- plagioclase-quartz assemblage, according to the reaction

biotite + plagioclase + quartz = hornblende + k-feldspar + melt (4)

Figure 14. Amphibole-bearing pegmatitic leucosome (top middle of photo), and amphibole-free pegmatite (bottom left) within a migmatitic amphibolite. Magnet for scale.

However, experimental studies of vapour-absent melting of the assemblage biotite-plagioclase- quartz at moderate to high pressures (Gardien et al., 1995;

Patino Douce and Beard, 1995), record only incongruent dehydration melting reactions that produce the anhydrous mafic minerals ortho- and clinopyroxene.

The relatively few experiments that deal with vapour-absent melting of rocks that contain both amphibole and biotite (Rutter and Wyllie, 1988; Skjerlie and

Johnston, 1993) also record only the production of anhydrous mafic phases through dehydration melting of both biotite and hornblende. The presence of

amphibole in the final product of such melting experiments is therefore due either to relict unreacted amphibole that was present in the starting material (Skjerlie and Johnston, 1993), or because H2O was added to the system (Conrad et al,

1988; Johnston and Wyllie, 1988; Patino Douce and beard, 1995; Gardien et al,

2000). Starting with an assemblage of biotite, plagioclase and quartz, Gardian et al, (2000) have shown that at 10 kbar pressure, 4 wt% total H2O is required to stabilize amphibole between 800-900°C, and at 15kbar <2 wt% H2O is required to stabilize amphibole. Therefore, the presence of amphibole in some leucosomes, and the absence of ortho- and clinopyroxene, suggest melting occurred from vapour-present incongruent melting reactions that involve the breakdown of biotite, and possibly amphibole, and the concomitant crystallization of a new amphibole phase.

biotite + plagioclase + quartz + H2O = hornblende + melt (5)

(Kenah and Hollister, 1983; and Viruete, 1999)

biotite + plagioclase + quartz + hornblende1 + H2O = hornblende2 + melt (6)

(Conrad et al., 1988)

Vapour-present melting that involves the breakdown of biotite is consistent with the observed absence, or decreasing amount of biotite with increasing melt volumes in amphibolite from the WMC, as well as the pegmatitic nature of the leucosomes.

The five migmatitic pelitic schists sampled for thermobarometric analysis have been shown to have undergone dehydration melting above the solidus, wherein H2O is dissolved in the melt, with no free H2O present as a fluid phase.

However, pelitic schists intercalated throughout the amphibolite at varying stratigraphic and structural levels, show varying volumes of leucosomes, with some essentially free of leucosomes. A lower activity of H2O within certain metapelitic layers would result in subsolidus dehydration, and the release of a free fluid phase at P-T conditions identical to units undergoing supersolidus dehydration melting due to a higher H2O-activity. Figure 15 illustrates how dehydration reactions will shift into the subsolidus field with decreasing aH2O, thereby releasing a free H2O-vapour phase. The devolatization, and release of

CO2-rich fluids from calcsilicate units interlayered with the pelitic schist is a potential mechanism for lowering the aH2O within the pelitic schists.

The transition from upper amphibolite facies metabasites to granulite facies conditions is conventionally marked by the replacement of hornblende- plagioclase-quartz ± garnet by orthopyroxene-clinopyroxene, -plagioclase- bearing assemblages at intermediate pressures, and orthopyroxene-free garnet- clinopyroxene-plagioclase assemblages at higher pressures (Fig. 21) (Pattison,

2003). The absence of both ortho- and clinopyroxene from the amphibolites of this study implies that granulite facies conditions were not reached, and according to the appropriate experimentally determined reactions (Ellis and

Thompson, 1986; Conrad et al., 1988; Rushmer, 1991; Winther and Newton,

1991; Wolf and Wyllie, 1994; Patino Douce and Beard, 1995; Nair and Chako,

2000), limits their peak metamorphic temperatures to less than 800-850°C.

Figure 15. P-T diagram displaying the aH2O-controlled dehydration reaction quartz + muscovite + albite = alkali feldspar + sillimanite + H2O (dashed curves), and the aH2O-controlled solidus curves of the melting reactions quartz + alkali feldspar + muscovite + H2O = melt, and quartz + alkali feldspar + sillimanite + H2O = melt (solid curves). Note that at the P-T conditions of the WMC the dehydration reactions lie in the subsolidus field for aH2O < 0.6 (black curves), but lie in the supersolidus field for aH2O ≥ 0.6 (grey curves). Therefore, muscovite dehydration melting should not occur in units with an aH2O < 0.6. Diagram is modified from Johannes and Holtz (1996).

6.4 Interpretation of observed textures and zoning

Partial disequilibrium during diffusional relaxation of growth zoning in garnet

Application of equilibrium thermodynamics to geothermobarometry is based on the assumption that coexisting mineral phases continuously and simultaneously chemically equilibrate at some scale, along the P-T path followed by the rock. Chemical equilibrium must be attained not only between coexisting phases, but also amongst all elements within each individual mineral, and at the same P-T point along the P-T path. The possibility of partial disequilibrium within the garnet core is examined, wherein Mg, Fe and Mn have homogenized at the thermal peak of metamorphism, while Ca preserves the earlier prograde growth zonation.

Several authors (Loomis et al., 1985; Chakraborty and Ganguly, 1991;

Schwandt et al., 1996) have noted that diffusion of Ca in garnet is slower than diffusion of Mg, Fe and Mn. Despite the slower diffusivity of Ca, the coincidence of a relatively flat XGrs profile with a flat XSps and Fe/(Fe+Mg) profile across the garnet core (Fig. 12) may lead to the simple assumption that the Ca component has homogenized into the garnet core at the thermal peak similar to Fe and Mg.

To determine if in fact all of Mg, Fe and Ca diffused into the garnet core I make use of the equation for the characteristic length scale of diffusion for a cation, which is expressed as a function of time:

1 x = (Dt) 2

where D is the diffusion coefficient. Therefore, with knowledge of the diffusion coefficient of the component of interest, it is possible to calculate the length of time required for this component to diffuse into the core of garnet for a given radius.

Stoichiometry dictates that the diffusive flux of an individual species (e.g.

Mg, Fe, Mn or Ca in garnet), driven by a concentration gradient, depends not only on its own gradient, but also on the concentration gradient of all diffusing species in order to maintain electrical neutrality at each point in the crystal

(Brady, 1975; Anderson and Olimpio, 1977, Lasaga, 1979). For a system that contains n independent components with respect to diffusion, Fick’s first law for

th multicomponent diffusion relates the flux of the i component, Ji, to the

th concentration gradient of the j component, dCj/dx:

n dC j = − J i ∑ j=1 Dij dx

th where Dij is the diffusion coefficient for the i component in response to the gradient in the jth component. The result is (n - 1)2 independent diffusion coefficients. Anderson and Buckley (1973) have concluded that the cross terms in Fick’s equation are small compared to direct terms. Therefore, for a matter of simplification, if we assume that the cross terms can be ignored, the number of flux equations reduces to those that describe the flux of ith component in response to its own gradient. Following the conclusions of Anderson and

Buckley (1973), self-diffusion coefficients, D*, from Carlson (2006) are used with

the assumption that they provide a reasonable approximation, despite the multicomponent character of exchange in a four component garnet.

Carlson (2006) identified a relationship between diffusivity and garnet composition, wherein there is a tendency for each ion to diffusive more rapidly in

Mn-enriched and especially in Ca-enriched crystals. Taking the unit cell dimension (ao , i ) as the average of the unit-cell dimensions of the almandine (alm), pyrope (prp), spessartine (sps), and grossular (grs) end-members, weighted by their mole fractions (Xi):

≈ • + • + • + • ao X alm ao,alm X prp ao, prp X sps ao,sps X grs ao,grs the unit cell dimension may then serve as a proxy for the compositional influence of the host garnet on diffusion rates (Carlson, 2006). End-member unit cell dimensions are ao , a l m = 1.1525 nm; ao , p r p = 1.1456 nm; ao , s p s = 1.1614 nm; ao , g r s =

1.1852 nm (Ganguly et al., 1993; Geiger and Feenstra, 1997). An average garnet composition from the Wolverine metapelites is: Xalm = 0.72, Xpyp = 0.17, Xs p s =

0.05, Xgrs = 0.06, which yields an average unit-cell dimension of a0 ≈ 1.1537.

Using almandine as a reference composition, an expression for the self-diffusion coefficient can be written as a function T, P and ao:

+ −  + ∆  ∗ ∗ Q P V  = + − + ln D ln Do,alm k (ao ao,alm) RT

where lnDo,alm is the frequency factor for diffusion in almandine garnet; Q is an

Arrhenius activation energy; ∆V+ is an activation volume; k describes the sensitivity of the frequency factor to unit-cell dimension; and R is the ideal gas

+ constant (Carlson, 2006). The diffusional parameters lnDo,alm, k, Q, and ∆V for

Fe, Mg, Mn and Ca were extracted by Carlson (2006) by numerical simulation of stranded diffusion profiles in partially resorbed natural garnets of diverse compositions, and are included in appendix 2.

Chakraborty and Ganguly (1991) suggest that the diffusion coefficient of

Fe and Mg are very similar. Therefore, the Mg self-diffusion coefficient calculated from Carlson (2006) is used as a check of Fe-Mg chemical homogenization within garnet cores from the WMC. At the maximum potential temperature of

860°C for the WMC (Fig. 13), the Mg self-diffusion coefficient in garnet, DMg, is

-21 equal to 7.10 × 10 . Therefore, a garnet 5 mm in diameter would require ~28 my at 860°C to chemically homogenize from rim to core. This is not a geologically unreasonable length of time, considering It was noted above that ST was synchronous with and possibly outlasted the thermal peak of metamorphism. A granodiorite stock that intrudes the deepest exposed level of the WMC is apparently unaffected by the ST deformation, and in fact contains a weak foliation oriented at a high angle to ST within the country rock. A 71.0 ± 0.6 Ma U-Pb zircon date and 72.8 ± 0.2 Ma U-Pb monazite date from this granodiorite stock

(Ferri and Melville, 1994), suggests that the thermal peak of metamorphism is pre- Late Cretaceous. It will be shown in a later section on 40Ar/39Ar geochronology that the WMC did not cool below ~500°C until the Eocene (i.e., 49

Ma). Therefore, the garnets experienced temperatures greater than 500°C for a minimum of 25 Ma. It follows, that it is not entirely unreasonable to expect that the garnets experienced temperatures as high as 860°C for ~28 Ma. Given the

considerable uncertainty in diffusion coefficient values, the best evidence for relaxation of Fe and Mg growth zoning, is the flat Fe/(Fe+Mg) profile across the core of all garnets.

From a calculation of the characteristic length scale of diffusion for Ca in garnet at 860°C, similar to as was done for Mg, it is apparent that due to the slower diffusion of Ca in garnet, Ca would require ~94 my to chemically homogenize a garnet 5 mm in diameter. It is highly unlikely that these rocks were exposed to such high temperatures over such an extended period of time, and therefore it is likely that the garnet interior has not re-equilibrated with respect to

Ca at the thermal peak, but rather preserves a prograde growth zonation within its core. On close examination of the XGrs profiles in garnets that do not exhibit considerable resorbtion (embayments) within the core, it is apparent that there is a subtle zonation of consistently increasing XGrs values outwards from the core towards a near-rim maximum (Fig. 12). The fact that Fe and Mg values within the garnet core reflect the conditions at the thermal peak, while Ca values within the core have been preserved from earlier prograde garnet growth, presents a situation of partial disequilibrium within the garnet core, in which the Fe, Mg and

Ca components of garnet are not in equilibrium. It follows, that combined Ca end- member net-transfer reaction and Fe-Mg exchange thermobarometry should be used with caution at upper amphibolite-facies metamorphism. The problem is reduced at lower and higher grades, where diffusion is sufficiently slow and fast, respectively, such that growth zoning is preserved or entirely obliterated with respect to all elements.

Retrograde zoning

On cooling, the distribution coefficients of components between garnet and minerals in the matrix change, creating a new and constantly changing equilibrium composition on the rim of the garnet with cooling. Diffusion in garnet, even during cooling from granulite- to amphibolite-facies conditions for most reasonable geologic cooling rates, is sufficiently slow, such that diffusion ceases first in the core and last on the rim (Spear and Florence, 1992). This diffusional profile is evident along the rim of garnet and around biotite inclusions within garnet. The apparent inability of elements to diffuse to the core of garnet due to changing equilibrium conditions imposed on the rim during cooling implies that the core of garnet has retained the equilibrium composition from the near thermal peak of metamorphism.

The rise in Fe/(Fe+Mg) and XSps towards the garnet rim is symmetrical from core to rim, and not isolated to zones adjacent to ferromagnesian minerals at its rim. Thus, this zoning is interpreted to be due to diffusional modification at the rim caused by the reNTR (2r). Reaction (2r) is consistent with textural evidence of garnet resorbtion, and the presence of quartz, plagioclase, sillimanite and the most Fe-rich biotite within embayments, pressure shadows and reaction rims in garnet. Furthermore, the dissolution of garnet through reaction (2r) releases Mn into the matrix; however, biotite the only other significant Mn-bearing mineral contains negligible amounts of Mn compared to garnet, and Mn is partitioned back into garnet enriching the rim (Kohn and Spear, 2000). Robinson

(1991), Spear (1991), Spear and Florence (1992), and Kohn and Spear (2000)

have discussed how simultaneous operation of Fe-Mg exchange and reNTR’s cause reactant and product mineral compositions to shift in the same direction, i.e. the shift of garnet and biotite to more Fe-rich compositions. A shift in biotite to more Fe-rich compositions has the effect of returning a higher than peak T, using an Fe-Mg exchange thermometer.

Fe/(Fe+Mg) zoning also occurs locally within garnets adjacent to biotite inclusions. This zoning is the result of retrograde Fe-Mg exchange between garnet and biotite through the reaction:

pyrope + annite = almandine + phlogopite. (7)

Reaction (7) is restricted to diffusive Fe-Mg ion exchange between garnet and biotite, with no interaction with matrix phases, and no significant change in mineral modes. In contrast to the net-transfer reactions, the Fe/Mg ratio within garnet adjacent to biotite inclusions increases, while the Fe/Mg ratio of biotite decreases; consistent with biotite inclusions in garnet having the lowest

Fe/(Fe+Mg) values.

An estimate of the temperature at the thermal peak of metamorphism using the garnet-biotite exchange thermometer, requires the preservation and identification of a biotite grain that was in equilibrium with the core of the garnet, and that has not since been modified through either retrograde net-transfer or exchange reactions, (2r) and (7) respectively. In contrast to garnet, diffusion in biotite is sufficiently rapid relative to geological cooling rates, such that the core of biotite quickly equilibrates with the rim, and biotite remains homogenous

(Spear and Florence, 1992). The compositional homogeneity of biotite precludes

use of the core of biotite grains together with the core of garnets to obtain an estimate of thermal peak.

By corollary, if biotite does indeed rapidly homogenize due to the changing equilibrium conditions imposed at its rim, then matrix biotite should be in equilibrium with the garnet rim, and not the garnet core. As no biotite was found included within an inert phase such as plagioclase, this leaves no remaining available biotite grains with a composition preserved from the metamorphic peak.

However, modelling experiments of Spear (1991) have shown that in the situation where the volumetric ratio of garnet to biotite is low (<0.1), as it is for samples from the WMC, the matrix biotite composition at the metamorphic peak is not substantially altered on cooling from the metamorphic peak. The preservation of peak metamorphic compositions within matrix biotite of the WMC is consistent with an Fe/(Fe+Mg) ratio that is intermediate between, the more Fe- rich biotite at the garnet rim that was created by the reNTR (2r), and biotite inclusions whose Fe/(Fe+Mg) ratio has been lowered by retrograde Fe-Mg exchange with the adjacent garnet. If the matrix biotite composition has not been substantially altered since the metamorphic peak, as would be inferred from the conclusions of Spear (1991) and the intermediate Fe/(Fe+Mg) values, then the minimum peak temperature may be at least estimated, although with a higher degree of uncertainty than determined from microprobe error and activity-models alone (Spear and Florence, 1992).

Qualitative P-T path

Prograde path

The XGrs profile reveals a subtle zonation within the garnet core that is characterized by low XGrs that increases outward to a near rim-maxima that is always accompanied by a rimward decrease (Fig. 12). For a given composition and mineral assemblage, the Gibbs method (Spear, 1988) may be used to determine quantitative P-T paths from growth zoning profiles as described by

Spear and Selverstone (1983). From specified values of each independent variable [mineral composition (X) and either T or P], the number of which is equal to the variance of the system, a Jacobian transformation may be performed on the thermodynamic variables P, T, X and mineral abundance (M), such that a mineral composition (X) may be contoured in P-T space. The absolute values are specific to the composition and modal mineral assemblage. As pointed out by Frost and Tracy (1991), difficulty in application of the technique is that it requires knowledge of the exact assemblage present at each stage of garnet growth. Frost and Tracy (1991) point out that in a rock’s path through middle- amphibolite facies, garnet undergoes a complex growth history involving garnet ± chlorite ± staurolite ± aluminosilicate bearing assemblages, and despite knowledge of the final equilibrated mineral assemblage, there is no guarantee that these other phases, for which there is no textural evidence, were not present during the early stages of garnet growth. In the absence of complete textural evidence for the assemblage during garnet growth, i.e., numerous inclusions within garnet, quantitative modeling of the P-T path is not feasible. Despite the

uncertainty regarding the mineral assemblage at the time of garnet growth, it is worth noting that XGrs isopleths for growth of garnet in both a chlorite-bearing assemblage (garnet + biotite + chlorite + plagioclase + muscovite + quartz +

H2O) (see Spear et al., 1990b, Fig. 1) and an aluminosilicate-bearing assemblage (garnet + biotite + sillimanite + plagioclase + quartz ± K-feldspar ± muscovite + H2O) (see Spear, 1989, Fig. 6) in the system NCMnKFMASH have positive slopes that steepen with increasing temperature. Therefore, while a quantitative determination of the P-T path is not possible, a qualitative P-T path trajectory may be estimated with reference to the slopes of the XGrs and mineral abundance (M) isopleths.

Figure 16, modified from Spear (1989), displays XGrs isopleths for the assemblage garnet + biotite + sillimanite + plagioclase + quartz ± K-feldspar ± muscovite + H2O, wherein muscovite is replaced by K-feldspar above the muscovite dehydration reaction. The presence of biotite and, toward the rim, rare sillimanite inclusions within garnet, suggests that garnet growth occurred at least partly within this assemblage. From this diagram, it is apparent that the increase in XGrs outward from a low at the garnet core, requires early garnet growth to have followed a prograde P-T path that is steeper than the XGrs isopleths (Fig. 16, path a).

Figure 16. Schematic P-T diagram illustrating possible P-T path trajectories with reference to XGrs isopleths. Increasing XGrs values outwards from the garnet core restrict the prograde P-T path to a trajectory steeper than the slope of the XGrs isopleths, as illustrated by arrows. Values on the XGrs isopleths are for reference of direction of increasing values, and do not represent the actual quantitative isopleth values. Note that the slopes of the XGrs isopleths steepen significantly to the high temperature side of the reaction muscovite + quartz = sillimanite + K-feldspar + H2O. XGrs isopleths were calculated by Spear (1989) for the assemblage garnet + biotite + sillimanite + plagioclase + quartz ± K- feldpsar ± muscovite + H2O.

Cooling and decompression path

A qualitative estimation of the prograde P-T path trajectory inferred from

XGrs zoning is complicated not only from an assumption of the mineral assemblage present during garnet growth, but also from a lack of Fe and Mg growth zoning that was homogenized at the thermal peak of metamorphism. In contrast to the uncertainties of growth within the garnet core, the identification of

retrograde biotite and sillimanite, together with the absence of retrograde muscovite and the preservation of K-feldspar, indicates that retrograde diffusional zonation at the garnet rim occurred entirely within the assemblage garnet + biotite + sillimanite + plagioclase + quartz + K-feldspar + H2O. Furthermore, the retrograde P-T trajectory may be better constrained due to the presence of both

Ca, as well as Fe and Mg diffusional zoning over roughly similar length scales

(Fig. 12). The shift in Xgrs to decreasing values at the garnet rim may be the result of garnet growth along a prograde P-T path that ranges from an increase of both temperature and pressure along a path shallower than the Xgrs isopleths (Fig. 16, path b), to isobaric heating (Fig. 16, path c), to P-T paths with increasing temperature and decreasing pressure (Fig. 16, path d). However, the abrupt change to decreasing Xgrs values at the garnet rim roughly coincides with the sharp increase in Fe/(Fe+Mg) and Xsps (Fig. 12) attributed to retrograde diffusional ion-exchange of Fe and Mg between garnet and biotite, and operation of reNTR (2r), which requires a path of decreasing temperature. It is therefore likely that the decreasing Xgrs at the rim of the garnet is produced by consumption of grossular by the reaction

grossular + quartz + 2aluminosilicate = 3anorthite (8)

during cooling and decompression (Fig. 16, path e). A decrease in XGrs at the garnet rim due to reaction (8) is consistent with XAn enrichment at the rims of plagioclase grains and anorthite-rich plagioclase in reactions rims and embayments of garnet (Fig. 17).

Figure 17. Photomicrograph metapelite sample RS-07-92A displaying A n -rich contents of plagioclase within textural locations interpreted to have crystallized from the reaction grossular + sillimanite + quartz = anorthite during retrogression. Blue circle outlines embayment within garnet filled with retrograde biotite, quartz and An-rich plagioclase. Matrix plagioclase (green circle) contains an An-rich rim. Photomicrograph taken in plane-polarized light.

A decompression path may be inferred from the relative change of anorthite and grossular, as dictated by equilibrium partitioning from reaction (8), which due to it’s shallow slope, is sensitive to changes in pressure. Spear et al.

(1990b) express that caution must be exercised when using equilibrium constants to infer metamorphic P-T paths from mineral chemical zoning, because the slopes of composition isopleths and lines of constant Keq may not necessarily

be the same. The different slopes of XGrs and lines of constant Keq for reaction

(8) are illustrated in figure 18. Note that not all P-T path trajectories that cross lines of constant Keq, for reaction (8), to lower pressures result in a decrease in

XGrs and an increase in XAn. In fact, decompression with cooling along a slope steeper than Keq, and shallower than the XGrs isopleth, will result in an increase in

XGrs. Therefore, garnet resorbtion with a decrease in XGrs at the garnet rim restricts the cooling and decompression path along a path steeper than both lines of constant Keq for reaction (8) as well as the XGrs isopleths (shaded region in Fig. 18).

Figure 18. Schematic P-T path diagram [modified from Spear (1989) and Spear et al. (1990)] illustrating the relationship between the equilibrium constant (Keq) for the reaction grossular + 2sillimanite + quartz = 3anorthite and isopleths for XGrs and XA n . Note that XGrs may increase along a cooling and decompression path steeper than the line of constant Keq so long as the paths slope is shallower than the XGrs isopleth. Retrograde An-rich plagioclase, as well as a rimward decrease in XGrs coincident with retrograde Fe/(Fe+Mg) zoning at the garnet rim, restricts the cooling and decompression path along a trajectory steeper than the XGrs isopleth. Slopes of XGrs isopleths steepen significantly to the high temperature side of the reaction muscovite + quartz = sillimanite + K- feldspar + H2O. XGrs and XA n isopleths were calculated by Spear (1989) for the assemblage garnet + biotite + sillimanite + plagioclase + quartz ± K-feldpsar ± muscovite + H2O. Note, isopleth and Keq values are for reference of increasing and decreasing direction only, and do not represent the actual values.

The absence of retrograde muscovite indicates a cooling and decompression path that passes below ~650°C and 4kbars, ‘IP’, the point of intersection of the H2O-saturated pelite solidus and the vapour-absent muscovite dehydration reaction (Fig. 19). As cooling begins within the solidus field, H2O is exsolved from the crystallizing melt providing an internal source of H2O and permitting crystallization of the hydrous mineral biotite (e.g., Kohn et al., 1997).

The presence of biotite and sillimanite selvages around leucosomes provides further support for reaction (2r) on cooling. Spear et al. (1999) have discussed that on cooling above the invariant point ‘IP’ H2O is exsolved from the crystallizing melt and retrograde muscovite will crystallize within the leucosome or be distributed throughout the matrix with very little K-feldspar remaining in the final assemblage. With cooling below the invariant point ‘IP’, the melt has fully crystallized upon crossing the H2O-saturated pelite solidus, and an external H2O source is necessary for production of retrograde muscovite during back-reaction of the muscovite dehydration reaction at lower temperatures. The absence of retrograde muscovite in samples from the K-feldspar + sillimanite zone of the

W MC indicates that these rocks decompressed from at least 7.2 kbars to below 4 kbars prior to cooling below approximately 650°C. The trajectory of the cooling and decompression path inferred from the absence of retrograde muscovite, is consistent with the constraints from XGrs isopleths and lines of constant Keq for reaction (8) (Fig. 19).

Additional textural support for decompression along a clockwise P-T path comes from plagioclase-hornblende symplectic coronas formed around garnet within amphibolite sample RS-07-63D. These symplectites are interpreted to have formed from the retrograde reaction:

garnet + hornblende1 + quartz = plagioclase + hornblende2 (9) which requires a cooling and decompression path steeper than the shallow positive slope of reaction (9).

Figure 19. P-T diagram illustrating a cooling and decompression path inferred from the absence of retrograde muscovite, and that is consistent with the observed change in composition of late plagioclase, and XGrs zoning at the garnet rim. XGrs isopleths were calculated by Spear (1989) for the assemblage garnet + biotite + sillimanite + plagioclase + quartz ± K-feldpsar ± muscovite + H2O. Note, isopleth and Keq values are for reference of increasing and decreasing direction only, and do not represent the actual values.

6.5 Thermobarometry

Analytical techniques

Pressures and temperatures were calculated using the TWEEQU program

(version 2.32) of Berman (1991), which incorporates the thermodynamic database of Berman (1988) and uses internally consistent end-member thermodynamic data to calculate the position of a set of independent equilibria in

P-T space. The equilibria selected for thermobarometry for the metapelitic gneiss include: the garnet-biotite exchange equilibrium, reaction (7), and the

GASP net-transfer equilibrium (8). The activity models used were that of Berman and Aranovich (1996) for garnet, Berman et al. (2007) for biotite, and Furhman and Lindsley (1988) for plagioclase.

The garnet-biotite exchange thermometer was also applied to the amphibolite gneiss samples, along with the garnet-amphibole-plagioclase-quartz

(GAPQ) barometers:

tschermakite + quartz + pyrope + grossular = anorthite + tremolite

Kohn and Spear (1990) (10)

and

Fe-tschermakite + quartz + almandine + grossular = anorthite + Fe-actinolite

Kohn and Spear (1990) (11)

Amphibole solution properties are absent from version 2.32; therefore, the garnet-biotite and GAPQ equilibria for the amphibolite gneiss were calculated using TWEEQU version 1.02 of Berman (1991), which incorporates the

thermodynamic database of Berman (1988) together with the activity models of

Mader et al. (1994) for amphibole; Berman (1990) for garnet; McMullin et al.

(1991) for biotite; and Furhman and Lindsley (1988) for plagioclase. The older solution models from v. 1.02 for garnet and biotite are applied to the amphibolite gneiss, as opposed to the more recent solution models from v. 2.32, in order to maintain the internal consistency of the thermodynamic data between the equilibria selected for thermobarometry.

In the absence of greater than a single thermometer and barometer, and since the range in P-T estimates due to intracrystalline chemical variation may often far outweigh analytical and thermodynamic data error, the P-T estimates are therefore given as a range using the minimum and maximum composition values for each textural location. Where chemical composition data are unavailable for greater than one textural location, a reasonable estimate of overall uncertainty is around ± 50°C and 1 kbar, as suggested by Berman (1991).

Near-thermal- and post-thermal-peak thermobarometry

Metapelites

The determination that the Ca composition within the garnet core equilibrated at a different P-T point along the P-T path than Fe and Mg, precludes the combined use of Fe-Mg exchange thermometry with the Ca end- member GASP geobarometer for an estimation of near-thermal-peak pressures.

Despite this limitation, garnet core Fe and Mg compositions, interpreted to have equilibrated with the matrix at the thermal peak of metamorphism, may be paired with matrix biotite compositions to determine the placement of the garnet-biotite

equilibria in P-T space at the thermal peak of metamorphism (Fig. 20). The kyanite ↔ sillimanite transition provides an upper pressure limit for the thermal peak of metamorphism. The absence of relict kyanite grains and the presence of sillimanite inclusions within garnet suggests that sillimanite was the stable aluminosilicate polymorph at the thermal peak of metamorphism. The XGrs profile at the garnet rim is interpreted to have resulted from reaction (8) during clockwise decompression and cooling (Figs 18 and 19). Therefore, a lower pressure limit at the thermal peak is constrained from the GASP barometer with compositions taken from the garnet rim and plagioclase within embayments in garnet.

A post-thermal peak P-T point (Fig. 20) is estimated by taking compositions from the garnet rim together with biotite in pressure shadows, reaction rims and embayments in garnet, and plagioclase grains in textural locations indicative of late crystallization, such as embayments within garnet with higher than matrix XGrs contents. Post-thermal-peak P-T estimates are limited to those samples in which biotite grains are identified that are not shielded from the matrix, such that it is in equilibrium with garnet and plagioclase that is affected by reaction (8), and which have minimal to no Fe-enrichment from reNTR (2r).

Figure 20. P-T diagram showing calculated near thermal peak and post thermal peak P-T conditions (shaded regions) of the metapelites. Near-thermal temperatures are calculated from garnet core – matrix biotite pairs, and post-thermal peak temperatures are calculated from garnet rim – reaction rim, embayment filled, or pressure shadow biotite. The kyanite-sillimanite equilibria provides an upper pressure limit. A lower pressure limit is provided from Garnet rim – embayment filled plagioclase pairs. Dehydration melting reactions and solidus curves after Huang and Wyllie (1973;1974); Le Breton and Thompson (1988); Spear and Parrish (1996); and Spear et al. (1999).

Amphibolites

A similar assumption of partial disequilibrium between Fe, Mg and

Ca within the garnet core is made for the amphibolites. Near-thermal-peak temperature estimates are taken from garnet core – matrix compositions (Fig.

21). While upper and lower pressure constraints are not available as they are for the metapelites, the amphibolite samples RS-07-63C, -63D and 84C are interlayered with, and in close proximity [several meters (63C and 63D) to < km

(84C)] to, metapelites in which sillimanite is interpreted as the stable aluminosilicate polymorph at the thermal peak of metamorphism.

Two amphibolite samples RS-07-63C and -63D, and one metapelite RS-

07-64B were sampled within approximately 50 meters from each other, and between which there was no observable, or inferred, faulting or post- peak metamorphic tectonic displacement. Due to the close proximity of these three samples, the lower pressure of metapelite sample RS-07-64B is assumed to also apply to the amphibolite samples RS-07-63C and -63D. The overlap in the range of temperature estimates for each of these three samples, further constrains the

P-T estimates of these samples to 770-830 °C and 4-10.7 kbars.

The upper temperature estimate for sample RS-07-84C lies well above experimental constraints on amphibole-out orthopyroxene-in reactions (Fig. 21).

Therefore, the reaction for the upper stability limit of amphibole in the presence of excess water:

amphibole + quartz + H2O = clinopyroxene + orthopyroxene + melt (12)

(Ellis and Thompson,1986), is taken as the upper temperature limit of near- thermal peak metamorphism.

Figure 21. P-T diagram showing calculated near thermal peak temperature conditions of the amphibolites (shaded region). Temperature constraints are from garnet core – matrix biotite Fe-Mg exchange thermometry. The upper pressure limit is the Kyanite-sillimanite equilibria, considering sillimanite is the stable aluminosilicate polymorph at the thermal peak within the metapelites, which the amphibolites are interlayered with. The lower pressure limit for samples RS-07-63C and -63D is taken from the calculated lower pressure limit for metapelite sample RS-07-64B. The quartz-saturated metabasite solidus is from Percival (1983); Johannes (1978); and Wyllie and Wolf (1993). Experimental constraints on the vapor-absent dehydration reactions: Hbl + Qtz = Opx + Cpx + Pl + L are taken from Ellis and Thompson (1986) (E&T 86); Conrad et al. (1988) (C 88); Beard and Lofgren (1991) (B&L 91); Rushmer (1991) (R 91); Wolf and Wyllie (1994) (W&W 94); Patino Douce and Beard (1995) (P&B 95); Nair and Chako (2000) (N&C 00), and Hbl + Pl + Qtz = Grt + Cpx + L are taken from Winther and Newton (1991) (W&N 91); Rushmer (1993) (R 93); Sen and Dunn (1994) (S&D 94); Patino Douce and Beard (1995). Diagram is modified from Pattison (2003).

Discussion of amphibolite- granulite-facies transition

Pattison (2003) has observed that P-T estimates for a collection of natural granulite-grade othopyroxene-free garnet + clinopyroxene + plagioclase ± hornblende ± quartz assemblages are 50-200 °C lower than implied by experimentally determined dehydration melting reactions with minerals of a similar mineral composition. Pattison has suggested that the lower P-T estimates of natural samples is due to late Fe-Mg resetting in natural mineral samples resulting in a underestimate of peak P-T conditions, and/ or the vapor-absent dehydration melting experiments were overstepped due to sluggish reaction kinetics in the relatively low experimental temperature range of 700–800 °C.

Care has been taken in the present study to identify compositional zoning resulting from Fe-Mg resetting during cooling, with near-peak P-T estimates coming from analyses of the core of minerals that are unaffected by such resetting. The P-T results of amphibolites with the mineral assemblage hornblende + plagioclase + quartz + garnet are consistent with the well- constrained P-T results from the metapelites with which the amphibolite is

interlayered and in close proximity (several m’s). These P-T results restrict the replacement of hornblende + plagioclase + quartz + garnet by clinopyroxene + plagioclase ± garnet ± orthopyroxene-bearing assemblages to a minimum temperature of 830°C at 10 kbars, which coincides with the lower T range of experimentally determined reactions (Fig. 21). Note however that the addition of at least 2-4 wt% H2O to the amphibolites was invoked in the above discussion to explain the presence of amphibole as a stable product during partial melting of the amphibolites. The infiltration of H2O has the effect of stabilizing amphibole to higher temperatures, and may explain, in part, why this upper amphibolite assemblage records higher temperatures than other natural examples with an amphibolite-granulite facies transition assemblage.

With the exception of sample 84C, the P-T estimates of the Wolverine amphibolite lie entirely below the experimentally determined reaction (12) of Ellis and Thompson (1986) for the disappearance of amphibole with excess-H2O.

These findings provide a natural constraint for the minimum T of the amphibolite- granulite facies transition for metabasites at 10 kbars pressures with excess-H2O.

6.6 Cooling and decompression as inferred from post-thermal- peak thermobarometry vs. mineral reaction textures

Post-thermal-peak thermobarometry estimates conflict with mineral textures that suggest a cooling and decompression path that passes below

~650°C and 4kbars. Post-thermal-peak thermobarometry estimates are interpreted to indicate a false shallow retrograde path that is the result of closure of the garnet-biotite Fe-Mg exchange thermometer (7) and the GASP net-transfer

reaction barometer (8) at different temperatures. Several authors (Ghent et al.,

1988, Frost and Chako, 1989, Spear and Florence, 1992) have suggested that reaction (8) should close at a higher temperature (TC) than reaction (7) due to the slower diffusion of Ca and Na in garnet and plagioclase compared to Fe and Mg in garnet and biotite. The GASP reaction does not involve Fe and Mg end- members in the participating phases. Therefore, if we ignore the small changes in the activity coefficient of grossular caused by changes in the Fe/Mg ratio in garnet, continued Fe-Mg exchange between garnet and biotite at temperatures below that at which the GASP reaction has closed, will have the affect of shifting the calculated T and P down a line of constant Keq for the GASP reaction (Fig.

22) (Ghent et al., 1988; Frost and Chako, 1989). Frost and Chako (1989) and

Spear and Florence (1992) have concluded that in the situation where the net- transfer reaction barometer has closed at a higher temperature than the exchange thermometer, a P-T path calculated using the two-point method (i.e., garnet rim – core) with both points taking the composition of the bulk plagioclase, will yield an erroneous near isobaric cooling path. The identification of a small amount of plagioclase within a textural location suggestive of retrograde crystallization, and which has a composition different from the bulk plagioclase will yield a considerably steeper cooling path (Frost and Chako, 1989; Spear and

Florence, 1992). Computer models of Spear and Florence (1992) for the cooling of granulite-facies rocks that incorporate net-transfer and exchange reactions and multi-component diffusion in garnet, illustrate that several possible apparent

P-T paths result, depending on the degree to which garnet and plagioclase react

during cooling. Even in the case in which 100% of the plagioclase reacts with garnet during cooling, they note that due to earlier closure of a net-transfer reaction barometer, a P-T path calculated using a two-point, garnet core - rim, method will never yield a true path of isothermal decompression followed by near isobaric cooling. The modelling results of Spear and Florence (1992) indicate why the calculated post-thermal peak thermobarometric estimates from the

WMC, which incorporate plagioclase interpreted to have reacted to some extent with garnet during cooling, infer a P-T path of cooling with moderate decompression, but that is considerably less steep compared to the path inferred from the absence of retrograde muscovite.

Figure 22. P-T diagram illustrating the hypothetical relationship between closure temperatures for garnet-biotite (GB) thermometry and garnet-A l 2SiO5-quartz- plagioclase (GASP) barometry. P1 represents the point along the P-T path at which the GASP reaction closes; however, continued Fe-Mg exchange between garnet and biotite to a lower temperature provides a false (lower) P-T estimate, P2. A f t e r Ghent et al. (1988).

7: GEOCHRONOLOGY

7.1 Previous geochronology

Previous geochronology in the study area has been carried out by

Wanless et al. (1967, 1970, 1973, 1974, 1979) and Ferri and Melville (1994).

The majority of the geochronology from these studies utilized the K-Ar method on metamorphic biotite and muscovite, with lesser Rb-Sr and a single K-Ar hornblende analysis. K-Ar dates from these studies are consistently Early

Tertiary (33-59 Ma) (Fig. 23). At a localized scale, hornblende dates are consistently older than muscovite, and muscovite older than biotite. However, from one region to another in the complex (separated by 10’s of kilometres), there is an inconsistent record of cooling ages, with biotite in one area returning an older date than muscovite from another. Previous work (Parrish, 1976, 1979) suggested that the WMC had cooled by mid-Cretaceous time and the observed pattern of Eocene K-Ar ages was the result of an Eocene thermal-resetting of

Mesozoic minerals by magmatism and hydrothermal circulation near active faulting. Above the sillimanite isograd, there is an abundance of various generations of cross-cutting, post-transposition pegmatitic and granitic intrusions.

Such an intrusion was dated by Wanless et al. (1971) yielding 47 and 43 Ma K-Ar dates, respectively from a muscovite-biotite pair. Other potential post-tectonic regional heat sources in these areas include the Late Cretaceous Blackpine Lake granitic stock (62±7Ma Rb-Sr whole rock; Parrish, 1976), and the Late

Cretaceous Wolverine Range granodiorite stock (72.8 ± 0.2 U-Pb monazite; Ferri and Melville, 1994); however, these are older than the purported Eocene resetting. Eocene melts are apparently restricted to rocks of sillimanite-grade and higher; however, Eocene K-Ar and Ar-Ar dates are ubiquitous throughout the present study area, with no clear spatial association to areas of increased melt.

7.2 Sampling strategy

Two hornblende and four biotite separates taken from four rock samples were dated by the 40Ar/ 39Ar laser technique at the Pacific Centre for Isotopic and

Geochemical Research, Department of Earth and Ocean Sciences at the

University of British Columbia.

Sample location and rock type are summarized in appendix 1 and figure 2.

The sampling objective, was to have representatives of contrasting structural levels and metamorphic grade across the width of the complex, to determine the variations in ages between different structural levels, and proximity to the

Wolverine fault. Two biotite samples were taken from a pelitic schist. Biotite sample RS-07-06B was taken from the west side of the complex at a high structural level, and in close proximity to the Wolverine fault. The other, sample

RS-07-54 was taken from the opposite side of the complex at a high structural level. Amphibole and biotite samples RS-07-63D and -84C were taken from an amphibolite gneiss from the lowest exposed structural and stratigraphic level, located in the centre of the complex. Amphibole and biotite in both samples 63D and 84C enables a determination of the cooling rate between approximately 500° and 300 °C (Harrison, 1981; Grove and Harrison, 1996).

Figure 23. Compilation map of geochronological data and mapped isograds from the WMC. Black circles identify the sample location for 40A r / 39A r dates from the present study. Grey squares identify the sample location of previous geochronology compiled from Wanless et al. (1979), Parrish (1979), and Ferri and Melville (1994).

7.3 40A r / 39Ar results

The 40Ar/39Ar results are summarized in appendix 5 and plotted as release spectra and inverse isochron plots (Fig. 24 and 35). The consistency of K/Ca within each sample suggests that the hornblendes and biotites are not significantly contaminated by other mineral phases.

Amphibole

The first five low temperature steps of samples RS-07-63D and RS-07-

84C correspond to about 9.4 and 4.4% of the respective total 39Ar, and give relatively high apparent spectra ages (Fig. 24). This is a common observation for metamorphic hornblende (Harrison and McDougall, 1980, 1981) and is interpreted to be the result of incorporation of excess 40Ar into positions that favour enhanced rates of Ar transport out of hornblende, i.e., fluid inclusions and crystallographic vacancies at grain boundaries (Hodges, 1998). The trapped Ar component for samples 63D and 84C are determined from the y-intercept of the inverse isochron plot, and have an 40Ar/36Ar ratio of 1062 ± 240 and 351.7 ± 6.9, respectively. These values are significantly higher than 295.5 ratio of atmospheric argon and indicate the presence of excess 40Ar. Therefore, the correction for non-radiogenic 40Ar in the apparent age spectra, by assuming that any non-radiogenic argon in the specimen has an atmospheric composition, is incorrect. Inverse isochron plots however make no assumption about the trapped argon component and may thus give a better estimate for the age of the sample. The isochron age for samples RS-07-63D and RS-07-84C are 47.0 ±

2.0 Ma and 49.11 ± 0.58 Ma, respectively.

Biotite

The younger ages from the first few step-heating events (Fig. 25) are attributed to minor diffusive loss of 40Ar, possibly during prolonged residence at high temperatures in the deep crust. Each of the four biotite samples define a plateau, with between 65.7 to 94.8% of the total 39Ar released. All four samples have apparent age spectra that agree within uncertainty with the isochron ages.

Inter-sample ages, however, do not agree within uncertainty, and this is attributed to contrasting structural levels between samples. Samples RS-07-

63D and RS-07-84C produce respective apparent-ages of 48.14 ± 0.27 Ma and

48.07 ± 0.26 that are identical within uncertainty, and are both from within the centre of the complex at the deepest structural and stratigraphic levels within the k-feldspar + sillimanite zone. Samples RS-07-06B and RS-07-54 produce apparent ages of 50.66 ± 0.31 Ma and 53.49 ± 0.31, respectively, and they occur on the flanks of the complex, at lower metamorphic grade (garnet grade), and structurally and stratigraphically higher levels. The K-feldspar + sillimanite-grade rocks in the structurally and stratigraphically lowest levels, located in the core of the complex, cooled below 300°C somewhere between 2 to 6 million years later than the garnet grade rocks at structurally and stratigraphically higher levels on the flanks of the complex.

Figure 24. 40A r / 39Ar apparent age spectra and inverse isochron plots for hornblende. Box heights are 2σ on plateau spectra and error ellipses are 2σ on isochron plots. Sample number is labelled within each spectrum.

Figure 25. 40A r / 39Ar apparent age spectra for biotite. Plateau steps are filled, rejected steps are open. Box heights are 2σ. Sample number is labelled within each spectrum.

7.4 Cooling History

Of the two amphibolite samples that contain both hornblende and biotite, one (sample RS-07-63D) yielded hornblende and biotite dates that are concordant within experimental error, while the other (sample RS-07-84C) indicates biotite cooled in as little as 200, 000 years later than hornblende.

These results indicate relatively rapid cooling, ~1000°C/myr, between the respective closure temperatures of 500 ± 50 °C and 300 ± 50 °C for hornblende

and biotite (Harrison, 1981; Grove and Harrison, 1996) and imply rapid exhumation.

An apatite fission-track analysis of Ferri and Melville (1994) suggests cooling below 100 °C at approximately 48.3 ± 3.5 Ma. This sample was taken from a metapelitic schist or gneiss located above the sillimanite isograd and along strike with the biotite sample RS-07-63D, and at a slightly higher structural and stratigraphic position within the complex than biotite sample RS-07-84C.

The Ar-Ar data from this study suggests that garnet-grade rocks at a higher structural and stratigraphic level cooled approximately 2-6 m.y. prior to sample

RS-07-84C. Therefore, taking the apatite fission-track age from Ferri and

Melville (1994) and the biotite cooling age from sample RS-07-84C, an estimate of the cooling rate between the respective closure temperatures of biotite and apatite (300 ± 50°C and 110 ± 10°C, respectively, Grove and Harrison, 1996;

Gleadow et al., 1986) should thus represent a minimum estimate. Agreement of the biotite and apatite fission-track dates within experimental error extends this period of rapid cooling between approximately 500° and 100°C to 5 million years or less, giving a minimum cooling rate of ca. 100°C/myr.

Thermobarometric data indicates that these rocks decompressed by a minimum of 3.2 kbars prior to cooling below ~650°C. This early phase of nearly isothermal decompression followed by rapid cooling in the later stages of unroofing are the two fundamental characteristics of rapid unroofing by tectonic denudation in the thermal models of Ruppel et al., (1988) and Ruppel and

Hodges (1994).

7.5 Discussion

There is an apparent correlation between metamorphic-grade and apparent age (Table 1). For instance, the 166 Ma Rb-Sr muscovite date of

Parrish (1976, 1979) from a muscovite-biotite schist from Chase Mountain at the northwestern end of the WMC, occurs in the lowest grade metamorphic rocks assigned to the WMC (510-530°C at 6 kbars; Parrish, 1976). Slightly further south, in the Blackpine Lake area, rocks metamorphosed to temperatures 100°-

150°C higher than in the Chase Mountain area cooled below the same temperature (500°C; TC for Rb-Sr muscovite; Purdy and Jäger, 1976) at 70 Ma

(Parrish, 1976), ~100 Ma later than at Chase Mountain. At still higher metamorphic grade (730-850 °C and 6.8-12.5 kbars), rocks from the deepest exposed levels of the Complex (present study area) have cooled below 500°C at

47.0 to 49.1 Ma (Ar-Ar hornblende) ~120 Ma later than at Chase Mountain. K-Ar and Ar-Ar biotite ages present a similar pattern, where ages range from 93-84

Ma within lower amphibolite facies rocks in the Chase Mountain area, and decrease to 53-43 Ma in rocks above the sillimanite isograd in the Blackpine

Lakes and Manson River areas.

In addition to progressively younger cooling ages with increasing depth and metamorphic grade, there is also a substantial increase in cooling rate. For example, at the northern end of the WMC at Chase Mountain, Parrish (1976,

1979) obtained muscovite and biotite K-Ar dates from garnet-grade metamorphic rocks that span from 116 to 84 Ma, respectively. A discordance of 21 and 23 Ma between muscovite and biotite K-Ar pairs from the same rock samples and 12 to

82 Table 1. Summary of regional geochronology data from chapter 8 discussion. Isotopic Closure Sample # Geographic location Rock type Mineral dated system temperature (°C)* Metamorphic conditions Age 134 Chase Mountain muscovite-biotite schist Muscovite Rb-Sr 500 ≤ 510°C; 6 kbars (Parrish, 1976) 166 ± 18 (Parrish, 1976;1979) Muscovite K-Ar 350 94.5 ± 3.4 (Parrish, 1976;1979) Biotite K-Ar 300 91.1 ± 3.3 (Parrish, 1976;1979) 77 Chase Mountain biotite-muscovite schist Muscovite Rb-Sr 500 500° - 550°C; 6 kbars (Parrish, 1976) 154 ± 12 (Parrish, 1979) Muscovite K-Ar 350 116 ± 4.1 (Parrish, 1979) Biotite K-Ar 300 93.3 ± 3.3 (Parrish, 1979) 282 Chase Mountain muscovite-biotite schist Muscovite Rb-Sr 500 510° - 530°; 6 kbars (Parrish, 1976) 116 ± 7 (Parrish, 1979) Muscovite K-Ar 350 104 ± 4.1 (Parrish, 1979) Biotite K-Ar 300 83.6 ± 3.0 (Parrish, 1979) 300 Blackpine Lake area schist Muscovite Rb-Sr 500 < 680°C (Parrish, 1976) 70 ± 2 (Parrish, 1976;1979) biotite K-Ar 300 43.7 ± 1.5 (Parrish, 1979) RS-07-06B This study area garnet-biotite schist Biotite Ar-Ar 300 Garnet-grade 50.66 ± 0.31 RS-07-54 This study area garnet-muscovite-biotite schist Biotite Ar-Ar 300 Garnet-grade 53.49 ± 0.31 RS-07-63D This study area amphibolite gneiss Hornblende Ar-Ar 500 755°-865°C; 6.3-11 kbars 47.0 ± 2.0 Biotite Ar-Ar 300 48.14 ± 0.27 RS-07-84C This study area amphibolite gneiss Hornblende Ar-Ar 500 780°-920°C; 6.3-12.6 kbars 49.11 ± 0.58 Biotite Ar-Ar 300 48.07 ± 0.26

* Closure temperatures for Rb-Sr and K-Ar muscovite, as well as K-Ar and Ar-Ar biotite are taken from Purdy and Jäger (1976), and Grove and Harrison (1996), respectively.

83 71 Ma between Rb-Sr and K-Ar muscovite, records a slow thermal evolution for these lower-grade, higher crustal levels, i.e., rocks at Chase Mountain cooled from ~500°C to ~350°C (TC for Rb-Sr muscovite and K-Ar muscovite, respectively; Purdy and Jäger, 1976) in ≤ 71 Ma, which gives a minimum cooling rate of ~2°C/Ma. By contrast, muscovite and biotite K-Ar and Ar-Ar dates from sillimanite- to sillimanite + k-feldspar-grade rocks throughout the complex range from 51-47 Ma and 48-33 Ma, respectively (Wanless et al., 1979; Parrish, 1979;

Ferri and Melville, 1994; present study), and imply a much faster cooling rate for these higher-grade rocks.

The pattern of progressively younger isotopic cooling ages, and faster cooling rate, with increasing metamorphic grade and structural position within the deforming orogen is interpreted to reflect diachronous differential exhumation of different crustal levels. This is supported by the presence of Early Tertiary cooling ages within high-grade rocks of the orogenic core, which according to thermobarometric estimates, require 20-30 km of exhumation. These rocks are juxtaposed across the west-side-down, normal Wolverine fault against significantly lower-grade, upper crustal sections with Jurassic cooling ages.

Due the absence of dates from minerals with closure temperatures greater than 500 °C, we are limited to an interpretation based on cooling temperatures.

The absence of near-thermal peak metamorphic crystallization ages has limited previous estimates of peak metamorphism of the WMC to be synchronous with upper levels which cooled by the Middle Jurassic (Parrish, 1976, 1979; Ferri and

Melville, 1994), with deeper crustal levels either thermally reset (Parrish, 1976,

1979), or held at depth until the Eocene (Ferri and Melville, 1994). The apparent younging of cooling through the 500-300°C temperature interval with increasing metamorphic grade that accompanies the deeper structural and stratigraphic levels is very similar to what is observed in the Selkirk allochthon and Monashee complex in the southeastern Canadian Cordillera. There, Parrish (1995) noted a systematic younging of both peak thermal conditions and subsequent cooling in the 600- 300°C temperature interval with progressively deeper levels. Within the same region, peak metamorphism and the associated deformation are interpreted as occurring diachronously between 175 Ma at high structural levels and 50 Ma at the deepest structural levels (Parrish, 1995; Crowley and Parrish,

1999; Crowley et al., 2000; Gibson et al., 1999, 2005). Due the absence of peak thermal metamorphic ages in the WMC, a similar pattern of decreasing timing of peak metamorphism and deformation cannot be confirmed, nor rejected.

However, as was noted above in the section on metamorphism and microstructure, the thermal peak of metamorphism within the WMC may be constrained to pre- Late Cretaceous.

8: OROGENIC MODEL

Ferri and Melville (1994) provided a detailed tectonic synthesis for the region of the present study area that spans from Proterozoic rifting through to

Tertiary extension and transcurrent faulting. Their tectonic model includes a description of the formation of the WMC as a product of crustal thickening related to the obduction and accretion of Slide Mountain and Quesnellia Terranes from the west onto the paleocontinental margin of North America.

The following discussion, reconciles the Jurassic to Cretaceous(?) crustal thickening, metamorphism, and final Tertiary exhumation of the WMC in terms of recent tectonic models proposed for convergent orogens.

8.1 Structural divergence

In the study area, peak metamorphic mineral assemblages define a composite transposition foliation that is axial planar to northeast-vergent isoclinally folded anatectic leucosomes, suggesting that northeast-vergent deformation was synchronous with, and possibly outlasted, the thermal peak of metamorphism. Conspicuously absent from the study area, is a southwest- vergent phase of deformation that is observed to the north (Bellefontaine. 1989,

1990; Gabrielse, 1971, 1972, 1975; Gabrielse and Campbell, 1991; Mansy,

1972, 1974; Mansy and Dodds, 1976) and south (Murphy, 1987; Struik, 1988;

Rees, 1987; Reid, 2003). Bellefontaine (1990) to the north in the central

Swannell Ranges, and Murphy (1987) and Reid (2003) to the south in the northern Cariboo Mountains, noted an overprint of northeast-vergent by southwest-vergent deformation, and that this reversal in structural vergence was initiated prior to the attainment of peak metamorphic conditions. This conflicts with structures and metamorphic mineral assemblages from the study area, which lack southwest-vergent deformation, and suggest that the thermal peak of metamorphism occurred during northeast-vergent deformation.

A zone of structural divergence trends discontinuously along the length of the Omineca Belt (Price, 1986), with southwest-vergent structures to the west and northeast-vergent structures to the east. Gibson et al. (2005) have shown that D2 southwest-vergent and northeast-vergent deformation on the western and eastern flanks, respectively of the Selkirk fan, were initially developed during the onset of regional metamorphism in the Middle Jurassic. The superposition of

Cretaceous penetrative northeast-vergent deformation superimposed on an earlier transposition fabric is restricted to the higher grade rocks in the eastern flank and does not overprint southwest-vergent structures to the west (Gibson et al., 2005).

Brown et al. (1993) and Gibson et al. (2008) have suggested that the southwest- and northeast-vergent structures of the Selkirk fan originated in a prowedge and retrowedge setting, respectively, analogous to the doubly-vergent structures in the mechanical model of Willet et al. (1993) and Brown et al. (1993)

(Fig. 26A). According to this model, the absence of southwest-verging structures in both middle crustal levels of the WMC and upper crustal levels in the hanging

wall of the Wolverine fault, would suggest a position for this region within the retrowedge, analogous to the eastern flank of the Selkirk fan. However, rectifying the absence of southwest-verging structures by assigning this region to the retrowedge (a position east of all southwest-vergent structures) is problematic as southwest-verging structures occur in mid- to upper-crustal levels, to both the west and east of northeast-vergent structures 40km along strike to the north.

Ferri and Melville (1994) dealt with this problem by placing this area within an indenting wedge according to the tectonic wedging and delamination model of

Price (1986) that has since been modified by Colpron et al. (1998) (Fig. 26B).

According to these models, the accreted oceanic (Slide Mountain) and island arc terranes (Quesnellia), and possibly distal North American strata (Colpron et al.,

1998), form a tectonic wedge between the North American miogeocline and the crystalline basement. The wedge is roofed by a southwest-vergent thrust fault above which rocks are thrust and overturned to the southwest with a southwest- vergent asymmetry, while those within the wedge deform by dominantly northeast-vergent structures. Inherent in this version of the model of tectonic wedging, is that rocks within the tectonic wedge lie at deeper structural levels below both the southwest-vergent roof fault and those rocks above being displaced to the southwest. However, rocks of the Slide Mountain and

Quesnellia terranes, which lack southwest-vergent deformation proximal to the study area, and have therefore been assigned to a location within the tectonic wedge, are at most biotite-grade and commonly chlorite-grade, much lower

grade than southwest-vergent garnet-bearing rocks above the Swannell fault to the north.

Figure 26. Diagram of the two principal models for the development of structural divergence within the hinterland of the Canadian Cordillera. (A) Model presented by Brown et al. (1993) for the development of a southwest-vergent prowedge and northeast-vergent retrowedge above the point of detachment and subduction of the oceanic lithosphere beneath the continental lithosphere. (B) Tectonic wedging model of Colpron et al. (1998) modified after Price (1986).

It is suggested here that the presence of northeast-vergent deformation along strike and within a zone of southwest-vergent structures, may be reconciled by invoking the hybrid model of Gibson et al., (2008; Fig. 8). This model combines the development of a southwest-vergent prowedge and northeast-vergent retrowedge as proposed by Gibson et al. (2008), Brown (1993) and Willet et al. (1993), with the tectonic wedging model proposed by Colpron et al. (1998) and Price (1986). K-Ar data from the literature and Ar-Ar data from this study, as well as Ferri and Melville (1994) and Parrish (1979), are consistent with exhumation of the deepest structurally exposed levels of the WMC in the

Eocene. In the absence of U-Th-Pb dating, it remains possible that the deep crustal levels of WMC were not incorporated into the deforming orogenic wedge until the Cretaceous, similar to some of the northeast vergent high-grade rocks in the southeastern Canadian Cordillera (Crowley, 2000; Brown, 2004; Gibson et al., 2005). By contrast, upper level, biotite-grade rocks of the Proterozoic

Ingenika Group immediately west of the WMC, which lack southwest-vergent deformation were exhumed by 171 Ma. Exhumation of these rocks occurred at the same time, or slightly prior to the onset of > 10km of exhumation that occurred concurrently with development of southwest-verging structures in the southern Omineca Belt (173-167 Ma, Colpron, 1996; Gibson, 2005) when the orogenic wedge was in its early stages of development (Brown, 2004). It is possible that the upper crustal levels of the study area, which are at biotite- to chlorite-grade, were exhumed to sufficiently high crustal levels by Middle

Jurassic (171 Ma; Ferri and Melville, 1994), such that they were not significantly

affected by Middle Jurassic southwest-vergent deformation in the prowedge.

Alternatively, the high-level northeast-vergent Middle Jurassic structures may have been brought down from the east along the Wolverine fault, juxtaposing them along strike with the southwest-vergent structures to the north and south.

According to the tectonic models of Brown (2004) and Gibson et al. (2005,

2008), the Jurassic structures that have been previously exhumed to upper crustal levels remained unaffected as the deforming orogenic wedge detached and migrated eastward, progressively incorporating more inboard, deeper crustal levels into the base of the deforming orogen characterized by a transient northeast-vergent basal shear zone (Brown, 2004). Thermobarometry and microstructure analysis suggests that North American margin rocks of the WMC were buried to depths of 20 to 30 km while reaching temperatures of 730° to

830°C during the development of northeast-vergent structures that were deformed by non-coaxial penetrative flow. Prior to the rotation of structures into a steeper angle during post-transposition broad upright folding, the WMC would have been characterized by a more shallowly dipping transposition foliation with isoclinal recumbent folds. The association of a shallowly dipping transposition foliation and isoclinal recumbent folds coupled with high metamorphic grade, is referred to as a high-grade nappe association (HGNA) by Williams and Jiang

(2005) who suggest that it is common and represents crustal scale (kilometres thick) shear zones.

It is proposed that as the deforming orogenic wedge detached and migrated northeastward, deeper crustal levels (present WMC rocks) became

incorporated into the base of the prowedge within a transient northeast-vergent basal shear zone (Fig. 27A,B). Later extensional exhumation within the prowedge during the Eocene juxtaposed these highly transposed, high-grade northeast-vergent rocks of the WMC against low-grade upper crustal levels that were exhumed in the Jurassic prior to the development of southwest-vergent deformation within the prowedge (Fig. 27C).

Figure 27. Conceptual orogenic model for the development and exposure of structural divergence at the latitudes of the WMC. A diffuse zone of northeast-vergent ductile shear develops at the base of the prowedge, as the orogenic wedge is detached and translated eastward. (A) Late Jurassic after Gibson et al. (2008) (B) Cretaceous after Gibson et al. (2008) (C)Later extensional exhumation along the west-side-down Wolverine fault juxtaposes this zone of high-grade, northeast-vergent ductily deformed rocks of the WMC, against northeast- vergent upper crustal rocks to the west, which were deformed and exhumed prior to the onset of southwest-vergent deformation within the prowedge.

8.2 Exhumation

The southern WMC, exposed approximately 50 km to the south, was described by Struik (1993) to have been exhumed in the Eocene along SW-NE trending normal faults which form extensional crustal pull-aparts between north and northwest trending en echelon dextral strike-slip faults (Fig. 28). While this model seems appropriate to explain the exhumation of high-grade metamorphic rocks in the southern exposures of the WMC between 54 and 55.5 °N, the geometry of inferred normal faulting at the higher latitudes of the present study area strikes northwest, parallel to the regional strike-slip faulting, and an alternative model for exhumation is therefore required.

Peak P-T estimates place the WMC in the hinterland of a thick-skinned orogen. The growth and destruction of an entire orogen is modelled by Willet et al (1993), and is shown to be consistent with brittle-frictional critical Couloumb wedge theory (e.g., Davis et al., 1983; Dahlen, 1984), extended by Williams et al.

(1994) to include the effects of both brittle-frictional and ductile behavior. The theory holds that the wedge deforms internally, steepening its surface slope until it reaches a critical taper (Davis et al., 1983; Dahlen, 1984; Williams et al., 1994).

Under horizontal compression, a critically tapered wedge is on the verge of failure, with all the strain accommodated along a basal decollement as the wedge slides, and continues to grow at a constant taper as material is accreted to the toe and rear of the wedge (Davis et al., 1983; Dahlen, 1984). The critical taper of a horizontally compressive wedge is dependent on a balance of forces, including among others, the internal strength of the wedge, as well as the frictional

resistance to sliding, or basal traction force, which in turn is a function of the horizontal compressive force (Davis et al., 1983; Dahlen, 1984; Williams et al.,

1994). This metamorphic study has revealed that the WMC, once located within the base of the orogenic wedge, has experienced considerable devolatization and anatectic melting. Rutter and Brodie (1995), and Van der Molen and

Paterson (1979) and Rutter (1997), respectively, have shown that metamorphic devolatization reactions, and the presence of melt can substantially reduce rock strength (see also: Holtzman et al., 2005; Rosenberg and Handy, 2005; Harris,

2007). As the base of the orogenic wedge weakens, the taper angle must decrease. If the base is weakened at a rate faster than the taper can adjust through accretionary deformation, then the orogen will deform internally, entering a state of extension causing it to collapse in order to achieve the required taper angle (Willet et al., 1993). Extension by movement on normal faults results in rapid exhumation through tectonic denudation of overlying rocks. This process of extension driven by melt-enhanced weakening of the base of the orogenic wedge would have been accentuated by a reduction in the far-field compressive stresses caused by a decrease in the rate of convergence between the Farallon and Kula plates with the North American plate between 56 and 50 Ma (Stock and

Molnar, 1988), contemporaneous with Ar-Ar cooling (uplift) dates.

Figure 29 is a summary of the timing Cordilleran tectonic events related to the development and exhumation of the WMC. For a more detailed geological synthesis of the WMC and adjacent terranes spanning from the Proterozoic to

Tertiary, see Ferri and Melville (1994).

Figure 28. Distribution and orientation of Eocene extensional and dextral strike-slip faults in the region of the southern WMC. Note that to the north, the extensional Wolverine fault is at a high angle to the extensional faults to the south, and parallels the Eocene dextral strike-slip faults. (Modified from Struik, 1993)

Figure 29. Diagram of events within the Canadian Cordillera that are of interest to the development and exhumation of the WMC. Timing of events complied from Ferri and Melville (1994); Murphy et al. (1995); Monger et al. (1982) Gibson et al. (2005; 2008); Crowley et al. (2000); and references therein.

97 9: CONCLUSIONS

The accretion and obduction of Slide Mountain and Quesnellia onto parautochthonous rocks deposited onto the paleocontinental margin of North

America, led to significant crustal shortening and thickening, with middle to lower crustal levels of the WMC exposed to upper amphibolite-facies metamorphic conditions of 770-830°C and 7.2-12.5 kbar . A calculation of the characteristic length scale of diffusion for Fe, Mg and Ca within garnet is consistent with garnet zoning profiles, suggesting partial disequilibrium within the garnet core, wherein

Fe and Mg homogenized at the thermal peak of metamorphism, while Ca retains values reflecting prograde garnet growth. Disequilibrium between Fe, Mg and Ca within the garnet core prohibits the combination of Ca end-member barometry and Fe-Mg exchange thermometry to provide a unique estimate of pressure at the thermal peak of metamorphism. This problem is unique to the conditions of upper amphibolite-facies metamorphism, wherein the diffusivity of Fe and Mg is sufficiently rapid to homogenize into the core of a typically sized garnet (5 mm), while temperatures remain too low for Ca to diffuse this distance over a geologically reasonable length of time.

Within the study area, peak metamorphism was contemporaneous with the development of northeast-vergent, tight to isoclinal folds that are intrafolial to an axial planar transposition foliation. The absence of southwest-vergent deformation in the study area, as is observed to both the north and south, can be

explained by modelling the Wolverine complex as a diffuse northeast-vergent basal shear zone that formed beneath a southwest-vergent prowedge as the orogenic wedge translated eastward. Extensional exhumation of the Wolverine complex juxtaposed the northeast-vergent rocks of the WMC against the low- grade upper crustal levels of the Slide Mountain Terrane, which was exhumed in the Middle Jurassic prior to the development of a southwest-vergent prowedge.

Significant anatectic melting at the base of the orogenic wedge would have facilitated the decoupling of this basal shear zone from the overlying southwest-vergent rocks, but would have also significantly reduced the internal strength at the base of the orogen leading to a reduction in the orogen’s taper through extension. A shallow cooling and decompression path inferred from post-thermal peak thermobarometry conflicts with textural evidence that is consistent with a near isothermal decompression path. The erroneously shallow cooling path, as calculated by post-thermal peak thermobarometry, is attributed to closure of the GASP net-transfer reaction barometer at a higher temperature than the garnet-biotite Fe-Mg exchange thermometer (Ghent et al., 1988; Frost and Chako, 1989; Spear and Florence, 1992). The identification of a small amount of retrograde An-rich plagioclase within embayments in garnet yields a steeper cooling path than would otherwise be observed in its absence (Frost and

Chako, 1989; Selverstone and Chamberlain, 1990; Spear and Florence, 1992).

Despite the significant reaction of plagioclase with garnet during cooling and decompression, the P-T path determined from a two-point garnet core – rim pair calculated by combined ion-exchange and net-transfer reaction

thermobarometry, still underestimates the amount of isothermal decompression as suggested from mineral reaction and disequilibrium textures.

The calculated near isothermal decompression path indicates that a minimum of

11 km of exhumation occurred prior to cooling below 650°C. Hornblende and biotite 40Ar/39Ar data, as well as a previously published apatite fission track date, indicates that this early phase of near isothermal decompression was followed by a period of rapid cooling between 500° and 100 °C in the Eocene. A near isothermal decompression path followed by a stage of rapid cooling are the two diagnostic characteristics of the thermal models of Ruppel et al., (1988) and

Ruppel and Hodges (1994) for rapid unroofing by means of tectonic denudation.

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APPENDIX 1: SAMPLE LOCATIONS

NAD 27 S a m p l e Purpose Zone Northing Easting RS-07-02B Thermobarometry 10U 6155677 432963 RS-07-28 Thermobarometry 10U 6162972 434780 RS-07-63C Thermobarometry 10U 6156615 434392 RS-07-63D Thermobarometry 10U 6156615 434392 RS-07-64B Thermobarometry 10U 6156045 433900 RS-07-78 Thermobarometry 10U 6165624 434611 RS-07-92A Thermobarometry 10U 6156119 433312 RS-07-84C Thermobarometry 10U 6162662 435027 RS-07-06B 40Ar/ 39Ar geochronology 10U 6171183 413049 RS-07-54 40Ar/ 39Ar geochronology 10U 6170063 438943 RS-07-63D 40Ar/ 39Ar geochronology 10U 6156615 434392 RS-07-84C 40Ar/ 39Ar geochronology 10U 6162662 435027 RS-07-15 Geochemistry 10U 6153755 436550 RS-07-26A Geochemistry 10U 6162972 434780 RS-07-26C Geochemistry 10U 6162972 434780 RS-07-40C Geochemistry 10U 6162137 435825 RS-07-63A Geochemistry 10U 6156615 434392 RS-07-84B Geochemistry 10U 6162662 435027 RS-07-84C Geochemistry 10U 6162662 435027

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APPENDIX 2: VALUES OF DIFFUSION PARAMETERS AND ASSOCIATED UNCERTAINTIES

+ lnD*o,alm D* Q ∆V in m2/s k 1/nm k J / m o l cm3/ m o l Fe 17.638 ± 0.776 465.7 ± 15.4 264.55 ± 4.63 13.568 ± 0.799 Mg 19.947 ± 0.919 419.9 ± 41.7 244.21 ± 5.32 8.565 ± 0.647 Mn 17.619 ± 0.876 496.9 ± 51.5 264.44 ± 5.99 9.631 ± 1.111 Ca 22.572 ± 0.716 511.1 ± 16.6 230.56 ± 7.15 9.795 ± 1.139

All values calculated by and taken from Carlson (2006)

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APPENDIX 3: SUMMARY OF NEAR -THERMAL PEAK THERMOBAROMETRIC ESTIMATES AND REPRESENTATIVE COMPOSITIONS USED FOR THERMOBAROMETRY

Near-thermal peak T Near-thermal peak P Post-thermal peak T Post-thermal peak P Sample Rock type estimate (± 50°C) estimates (± 1 kbar) estimate (± 50°C) estimates (± 1 kbar) RS-07-02B metapelite 730-800 6.3-10.3 597-622 3.7-4.6 RS-07-28 metapelite 755-820 7.1-10.4 705 6.2 RS-07-63C amphibolite 770-915 4.0-12.5 –– –– RS-07-63D amphibolite 755-865 3.8-11.4 –– –– RS-07-64B metapelite 791-834 8.3-10.6 –– –– RS-07-78 metapelite 755-845 5.0-11.0 600 2.9 RS-07-84C amphibolite 780-920 max 12.6 –– –– RS-07-92A metapelite 750-785°C 7.2-9.7 660-680 4.8-6.1

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Table 2. Representative garnet compositions used for thermobarometry. Sample 02B 02B 28 28 63C 63C 63D 64B 78 78 84C 92A 92A Point 44 55 46 28 8 23 6 14 8 22 11 17 10 Textural position Core Rim Core Rim Core Rim Core Core Core Rim Core Core Rim wt% oxides NaO 0.02 0.03 0.01 0.00 0.00 0.00 0.02 0.01 0.04 0.04 0.02 0.00 0.00 MgO 5.19 3.28 5.19 4.53 3.84 2.73 2.52 4.08 3.94 1.83 3.18 5.19 4.74 Al2O3 21.78 21.43 22.04 22.12 21.88 21.66 21.71 21.61 21.71 21.41 21.46 21.65 21.65

SiO2 37.49 37.01 37.55 37.36 37.51 37.10 36.55 36.79 36.83 36.49 36.40 36.46 37.78 CaO 0.93 0.91 1.86 1.60 7.82 8.03 10.61 2.06 1.95 1.88 10.84 3.51 2.14 TiO2 0.02 0.05 0.00 0.00 0.03 0.02 0.07 0.01 0.00 0.00 0.08 0.02 0.02

Cr2O3 0.04 0.02 0.00 0.01 0.00 0.07 0.00 0.02 0.00 0.01 0.02 0.00 0.02 MnO 0.88 2.65 1.37 1.99 2.91 4.54 4.54 1.77 3.66 10.37 5.75 0.69 0.84 FeO* 32.83 34.29 32.37 33.07 26.00 26.05 23.76 32.82 31.73 28.13 21.69 31.36 33.37 Total 99.19 99.67 100.38 100.68 100.00 100.20 99.77 99.17 99.85 100.16 99.43 98.88 100.56

* = Total Iron (FeO + Fe2O3)

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Table 3. Representative biotite compositions used for thermobarometry. Sample 02B 02B 28 28 63C 63C 63D 64B 78 78 84C 92A 92A Point 19 13 20 27 6 25 6 17 16 36 6 22 11 Textural position Matrix psgra Matrix egb Matix scc Matrix Matrix Matrix Garnet rim Garnet rim Matrix psgra wt% oxides F 0.16 0.14 0.24 0.18 0.02 0.21 0.31 0.31 0.63 0.57 0.43 0.25 0.36 NaO 0.18 0.25 0.20 0.30 0.13 0.19 0.10 0.19 0.11 0.14 0.17 0.11 0.15 MgO 9.56 8.04 8.92 10.12 11.47 11.21 8.29 7.57 7.10 7.37 10.23 9.99 10.79 Al2O3 20.03 22.30 18.36 18.69 16.05 15.01 13.99 18.74 18.03 17.91 15.06 19.04 19.76

SiO2 35.48 35.86 35.14 35.85 33.45 36.48 35.45 34.32 35.01 35.67 36.21 35.20 35.49 Cl 0.00 0.01 0.03 0.00 0.01 0.00 0.30 K2O 9.79 9.74 9.63 9.85 6.13 9.71 9.22 9.66 9.68 9.44 8.48 9.26 8.97 CaO 0.02 0.02 0.02 0.00 0.10 0.02 0.03 0.01 0.00 0.05 0.10 0.01 0.05 TiO2 2.63 1.90 4.01 3.84 2.09 3.18 5.05 3.32 2.90 1.58 1.81 2.44 0.64

Cr2O3 0.06 0.07 0.12 0.05 0.06 0.09 0.00 0.10 0.01 0.01 0.04 0.11 0.00 MnO 0.09 0.02 0.06 0.07 0.15 0.12 0.51 0.20 0.37 0.55 0.45 0.00 0.06 FeO* 18.60 16.83 19.86 18.22 23.68 20.12 22.31 22.04 21.40 20.45 21.20 18.77 18.38 Total 96.60 95.19 96.59 97.17 93.34 96.34 95.26 96.75 95.24 93.75 94.16 95.19 94.65

* = Total Iron (FeO + Fe2O3) a = Pressure shadow at garnet rim b = Embayment in garnet c = Symplectic corona

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Table 4. Representative plagioclase compositions used for thermobarometry. Sample 02B 02B 28 28 63C 63C 63D 64B 78 78 84C 92A 92A Point 12 19 20 11 9 14 3 4 1 12 7 5 2 Textural position Matrix psd Matrix egb Matrix scc Matrix Matrix Matrix psd Garnet rim Matrix egb wt% oxides NaO 9.03 8.62 7.48 7.47 5.12 2.93 6.98 8.07 7.56 6.18 3.66 5.38 2.09 MgO 0.00 0.00 0.01 0.01 0.00 0.00 0.00 0.00 0.00 0.01 0.00 0.00 0.00 Al2O3 22.80 23.50 24.55 25.08 28.47 32.25 25.59 24.12 25.03 27.30 30.95 28.08 33.55

SiO2 64.16 62.59 61.48 60.57 55.35 50.13 57.70 62.10 58.00 56.58 50.97 54.51 47.47

K2O 0.29 0.24 0.58 0.16 0.13 0.03 0.40 0.24 0.24 0.23 0.10 0.33 0.08 CaO 3.85 4.58 6.01 6.67 10.52 14.78 7.54 5.57 6.69 9.16 13.82 10.45 16.51 MnO 0.03 0.01 0.02 0.00 0.00 0.00 0.00 0.00 0.01 0.00 0.00 0.00 0.00 FeO* 0.07 0.02 0.06 0.04 0.19 0.21 0.17 0.06 0.16 0.13 0.06 0.06 0.10 Total 100.22 99.56 100.18 100.00 99.78 100.34 98.38 100.16 97.69 99.59 99.56 98.81 99.80

* = Total Iron (FeO + Fe2O3) d = Garnet pressure shadow b = Embayment in garnet c = Symplectic corona

117

Table 5. Representative amphibole compositions used for thermobarometry. Sample 63C 63C 63D 84C Point 8 15 9 3 Textural position Matrix scc Garnet rim Matrix

wt% oxides F 0.134 0.167 0.173 0.024 NaO 1.245 1.178 1.242 1.272 MgO 8.576 8.661 6.718 8.581 Al2O3 13.317 11.921 12.314 12.204

SiO2 40.655 42.252 39.471 42.283 Cl 0.008 0.006 K2O 1.259 0.668 1.831 1.042 CaO 10.936 10.634 11.165 10.950 TiO2 1.567 0.734 1.572 0.770

Cr2O3 0.035 0.101 0.010 0.033 MnO 0.236 0.558 0.659 0.782 FeO* 19.224 20.273 20.731 17.835 Total 97.192 97.155 95.884 95.774

* = Total Iron (FeO + Fe2O3) c = Symplectic corona

118 APPENDIX 4: 40AR/ 39AR ANALYTICAL TECHNI QUES AND SAMPLE DESCRIPTIONS

Two hornblende and four biotite separates taken from four rock samples were dated by the 40Ar/ 39Ar laser technique at the Pacific Centre for Isotopic and

Geochemical Research, Department of Earth and Ocean Sciences at the

University of British Columbia.

Conversion of 39K to 39Ar is done by irradiation of the sample with fast neutrons in the McMaster University reactor. Samples are individually wrapped in Al foil, and stacked in a small sample can. The efficiency of the 39K → 39Ar reaction is determined using a flux monitor (biotite or sanidine from the Fish

Canyon tuff in Colorado) with an inferred age of 28.02 Ma (Renne et al., 1998), interspersed with every 3 to 5 unknowns. Finally, argon isotope analyses are performed with a Micromass VG5400 noble gas mass spectrometer.

Each sample was step heated with between 9 to 15 steps, and the results displayed on release spectra and inverse isochron plots (Figs 26 & 27). In this study, a “plateau” age is a mean age weighted relative to the fraction of 39Ar released from a sequence of adjoining steps with ages that lie within 2σ uncertainty of the error-weighted mean of the group, and together represent

>50% of the total 39Ar released.

The calculated “plateau” ages on age spectra is based in part on the assumption that the trapped Ar component has the same 40Ar/36Ar ratio as

119

modern atmosphere (295.5). However, samples often contain a trapped component with 40Ar/36Ar > 295.5. This higher ratio is attributed to excess 40Ar.

Therefore, the ages from samples that contain excess 40Ar, are determined from inverse isochron plots, which make no assumption about the composition of the trapped argon component, and permit the independent characterization of both trapped and radiogenic Ar components. The apparent ages of spectra become older with excess 40Ar.

Sample RS-07-06B

Sample RS-07-06B is a garnet-biotite quartzo-feldspathic schist. Garnet occurs as poikiloblasts up to 5mm in length. Garnet is highly resorbed, containing abundant inclusions of quartz and lesser biotite. Biotite defines the schistosity, and has a reddish-brown to near colorless pleochroism. Some biotite grains have been altered and completely pseudomorphed by chlorite. Other biotite grains show slight replacement by chlorite along the rim and cleavage planes, while other grains show no alteration. A separate of biotite grains yielded a date of 50.66 ± 0.31 Ma based on five contiguous steps containing 77.9% of the 39Ar, and is interpreted as a metamorphic cooling age through 300 ± 50°C.

Sample RS-07-54

Sample RS-07-54 is a garnet-muscovite-biotite schist. Garnet and muscovite form poikiloblasts and porphyroblasts up to 3mm in length that are set in a quartz-rich matrix. Garnet is highly resorbed with numerous inclusions of quartz. A crystallographically controlled orientation of muscovite and biotite

120

defines the foliation. Biotite has a reddish-brown to faint reddish-brown, almost colorless pleochroism. Chlorite alteration of biotite is rare, and where present is restricted to a narrow zone at the rim or along fractures and cleavages. A separate of biotite grains yielded a date of 53.49 ± 0.30 Ma based on nine contiguous steps containing 80.1% of the 39Ar, and is interpreted as a metamorphic cooling age through 300 ± 50°C.

Sample RS-07-63D

Sample RS-07-63D is a migmatitic garnet-biotite-amphibole gneiss, or amphibolite. The sample represents the mesosome which contains approximately 50% hornblende and 30% biotite. Biotite has a very dark brown to light brown pleochroism and is very fresh with no alteration. Amphibole has a dark green to very light green pleochroism, and is also fresh with no alteration.

Biotite is relatively free of inclusions, while amphibole contains a minor amount quartz and biotite inclusions. A separate of biotite grains yielded a date of 48.14

± 0.27 Ma based on six contiguous steps containing 65.7% of the 39Ar, and is interpreted as a metamorphic cooling age through 300 ± 50°C. A separate of hornblende grains fails to define a plateau. The relatively high apparent spectra ages from the first few low temperature steps indicates the presence of excess

40Ar. An 40Ar/36Ar ratio of 1062 ± 240 is significantly higher than 295.5 (the composition of atmospheric argon). An inverse isochron plot, which makes no assumption about the trapped argon component and may therefore provide a better estimate for the age of the sample, yields an age of 47.0 ± 2.0 Ma.

121

Sample RS-07-84C

Sample RS-07-84C is a migmatitic garnet-biotite-amphibole gneiss, or amphibolite taken from part of the mesosome. Amphibole and biotite constitute

50 and 20% of the modal mineralogy, respectively. Biotite and amphibole contain numerous inclusions of zircon and/or monazite. Both biotite and amphibole show no evidence of alteration. A separate of biotite grains yielded a date of 48.07 ± 0.26 Ma based on eleven contiguous steps containing 94.8% of the 39Ar, and is interpreted as a metamorphic cooling age through 300 ± 50°C. A separate of hornblende grains fails to define a plateau. The relatively high apparent spectra ages from the first few low temperature steps indicates the presence of excess 40Ar. An 40Ar/36Ar ratio of 351.7 ± 6.9 is significantly higher the assumption for the composition of atmospheric argon when computing a plateau age. An inverse isochron plot, which makes no assumption about the trapped argon component yields an age of 49.11 ± 0.58 Ma, and is interpreted as a metamorphic cooling age through 500 ± 50°C.

122 Table 6. 40A r / 39Ar results for biotite and hornblende. RS-07-06B Biotite J = 0.005039±0.000010 Laser P o w e r ( % ) 40Ar/39Ar 38Ar/39Ar 37Ar/39Ar 36Ar/39Ar Ca/K Cl/K %40Ar atm f 39Ar 40Ar*/39ArK Ag e 2 21.4619±0.0173 0.0406±0.1266 0.0257±0.1953 0.0626±0.0461 0.061 0.004 86.38 0.9 2.684±0.814 24.24±7.30 2.2 7.9964 0.0063 0.0262 0.0593 0.0187 0.0821 0.0091 0.0434 0.043 0.002 32.59 7.18 5.228 0.122 46.91 1.08 2.4 6.4764 0.0079 0.0233 0.0432 0.0182 0.0385 0.0028 0.0558 0.042 0.002 11.35 12.88 5.613 0.064 50.32 0.57 2.6 6.3185 0.0082 0.0236 0.0676 0.0208 0.0493 0.0020 0.0796 0.048 0.002 7.92 11.15 5.671 0.068 50.83 0.60 2.8 6.0861 0.0051 0.0234 0.0337 0.0180 0.0421 0.0012 0.0422 0.042 0.002 5.2 26.55 5.687 0.034 50.97 0.30 3 5.9625 0.0053 0.0236 0.0208 0.0162 0.0520 0.0010 0.0627 0.037 0.002 3.68 18.39 5.637 0.035 50.53 0.31 3.1 6.0955 0.0061 0.0231 0.0505 0.0162 0.0673 0.0014 0.0627 0.038 0.002 4.74 8.94 5.627 0.045 50.44 0.39 3.3 5.9778 0.0082 0.0220 0.0827 0.0155 0.1010 0.0011 0.1423 0.036 0.002 3.34 8.1 5.581 0.068 50.03 0.60 3.6 6.0450 0.0151 0.0218 0.0743 0.0284 0.0680 0.0016 0.1414 0.066 0.002 4.93 5.9 5.492 0.112 49.25 0.99

RS-07-54 Biotite J = 0.005032±0.000008 Laser P o w e r ( % ) 40Ar/39Ar 38Ar/39Ar 37Ar/39Ar 36Ar/39Ar Ca/K Cl/K %40Ar atm f 39Ar 40Ar*/39ArK Ag e 2 24.0633±0.0110 0.0380±0.2364 0.0121±0.4524 0.0719±0.0519 0.022 0.003 86.44 0.29 3.035±1.090 27.34±9.74 2.2 11.2322 0.0095 0.0159 0.0908 0.0079 0.1925 0.0195 0.0325 0.019 0 49.24 1.35 5.502 0.189 49.27 1.67 2.4 7.9829 0.0050 0.0156 0.0339 0.0036 0.1353 0.0074 0.0326 0.009 0 26.55 6.14 5.776 0.077 51.69 0.68 2.6 6.9339 0.0044 0.0141 0.0432 0.0028 0.0757 0.0033 0.0313 0.007 0 13.67 12.11 5.920 0.041 52.96 0.36 2.8 6.3061 0.0061 0.0137 0.0513 0.0022 0.0620 0.0011 0.0572 0.006 0 4.8 22.74 5.950 0.042 53.22 0.37 2.9 6.2202 0.0084 0.0142 0.0406 0.0014 0.2602 0.0008 0.0796 0.003 0 2.8 9.2 5.964 0.054 53.35 0.48 3 6.2742 0.0069 0.0139 0.0355 0.0015 0.1489 0.0009 0.0668 0.004 0 3.25 8.18 5.983 0.046 53.51 0.41 3.1 6.3007 0.0059 0.0138 0.0303 0.0009 0.3570 0.0009 0.0881 0.002 0 3.16 7.15 6.006 0.044 53.72 0.39 3.2 6.2361 0.0045 0.0143 0.0222 0.0009 0.3431 0.0008 0.1055 0.002 0 2.89 8.47 5.970 0.038 53.40 0.33 3.3 6.3315 0.0053 0.0141 0.0413 0.0009 0.2826 0.0010 0.1313 0.002 0 3.34 5.22 6.003 0.052 53.69 0.46 3.5 6.2391 0.0051 0.0140 0.0547 0.0015 0.1599 0.0008 0.0810 0.004 0 2.85 8.72 5.977 0.037 53.46 0.32 3.7 6.2299 0.0047 0.0136 0.0471 0.0023 0.0683 0.0007 0.1141 0.006 0 2.41 8.29 5.993 0.038 53.60 0.33 4 6.3484 0.0069 0.0156 0.0611 0.0049 0.2198 0.0014 0.3151 0.012 0 2.48 2.13 5.954 0.134 53.26 1.18

123

RS-O7-63D Biotite J = 0.005040±0.000010 Laser P o w e r ( % ) 40Ar/39Ar 38Ar/39Ar 37Ar/39Ar 36Ar/39Ar Ca/K Cl/K %40Ar atm f 39Ar 40Ar*/39ArK Ag e 22.2 45.7016±0.0189 0.0848±0.1637 0.0493±0.2419 0.1554±0.0754 0.112 0.01 99.42 0.09 0.218±3.402 1.98±30.90 2.4 23.1609 0.0151 0.0737 0.1250 0.0273 0.2548 0.0647 0.0578 0.065 0.011 80.54 0.22 4.229 1.078 38.04 9.60 2.6 12.1006 0.0058 0.0647 0.0715 0.0102 0.1493 0.0226 0.0303 0.026 0.011 53.99 1.16 5.422 0.206 48.64 1.82 2.8 7.5025 0.0059 0.0531 0.0168 0.0049 0.1216 0.0068 0.0173 0.013 0.009 26.29 6.25 5.463 0.049 49.00 0.43 3 6.2809 0.0075 0.0466 0.0220 0.0039 0.1045 0.0021 0.0472 0.01 0.007 9.13 6.66 5.634 0.054 50.51 0.48 3.2 5.9034 0.0068 0.0488 0.0254 0.0037 0.0676 0.0011 0.0833 0.01 0.008 5.03 12.02 5.549 0.048 49.76 0.42 3.4 5.7085 0.0060 0.0467 0.0148 0.0046 0.0728 0.0006 0.0967 0.012 0.008 2.52 7.91 5.495 0.038 49.29 0.34 3.6 5.5489 0.0055 0.0486 0.0117 0.0156 0.0336 0.0005 0.0699 0.042 0.008 2.11 15.88 5.380 0.032 48.26 0.28 3.8 5.5708 0.0050 0.0458 0.0210 0.0155 0.0205 0.0006 0.0738 0.041 0.007 2.71 15.47 5.368 0.030 48.16 0.27 3.9 5.5078 0.0046 0.0440 0.0158 0.0021 0.0350 0.0005 0.0610 0.006 0.007 2.2 22.07 5.340 0.026 47.91 0.23 4 5.6485 0.0096 0.0455 0.0391 0.0035 0.0893 0.0008 0.1192 0.009 0.007 3.19 5.39 5.383 0.060 48.29 0.53 4.2 5.6554 0.0063 0.0503 0.0538 0.0089 0.0850 0.0009 0.1101 0.023 0.008 2.87 3.52 5.380 0.045 48.27 0.40 4.4 5.6724 0.0065 0.0461 0.0155 0.0061 0.1015 0.0009 0.1453 0.016 0.007 2.91 3.37 5.391 0.053 48.37 0.47

RS-07-63D Hornblende J = 0.005035±0.000008 Laser P o w e r ( % ) 40Ar/39Ar 38Ar/39Ar 37Ar/39Ar 36Ar/39Ar Ca/K Cl/K %40Ar atm f 39Ar 40Ar*/39ArK Ag e 2 725.9559±0.0830 0.9344±0.1680 -0.9818± 5.4360 1.5172±0.0891 0 0.16 59.91 0.02 296.859±32.3471649.23±117.78 2.3 333.7845 0.0137 0.4734 0.0544 1.4485 0.6803 0.7077 0.0262 3.987 0.077 61.99 0.12 126.464 5.155 888.89 28.61 2.6 29.3922 0.0077 0.0873 0.0638 0.0516 2.3164 0.0357 0.0327 0.138 0.015 34.52 1.1 18.890 0.372 163.91 3.08 2.9 20.1803 0.0070 0.0696 0.0895 ERR 0.0237 0.0369 0 0.012 33.36 1.65 13.191 0.274 116.01 2.34 3.1 11.7689 0.0048 0.1138 0.0240 0.8277 0.0208 0.0079 0.0315 2.274 0.023 18.29 6.55 9.518 0.088 84.45 0.76 3.2 7.2930 0.0061 0.1580 0.0137 1.9453 0.0224 0.0029 0.0955 5.351 0.033 6.98 6.78 6.686 0.091 59.73 0.80 3.3 6.4505 0.0080 0.1419 0.0177 1.8431 0.0163 0.0020 0.0305 5.069 0.029 4.29 7.86 6.083 0.054 54.42 0.47 3.4 7.2467 0.0069 0.1327 0.0274 1.5532 0.0283 0.0030 0.0839 4.271 0.027 7.91 5.04 6.547 0.088 58.51 0.77 3.5 6.4753 0.0045 0.1492 0.0156 1.8274 0.0143 0.0020 0.0589 5.026 0.031 5.07 13.51 6.085 0.046 54.44 0.40 3.6 6.4150 0.0051 0.1319 0.0192 1.7890 0.0142 0.0019 0.0704 5.018 0.027 4.18 11.64 6.078 0.051 54.38 0.45 3.7 6.2555 0.0048 0.1403 0.0130 1.9563 0.0133 0.0018 0.0480 5.488 0.029 4.31 27.94 5.945 0.039 53.21 0.35 3.8 5.8541 0.0058 0.1406 0.0164 2.0607 0.0188 0.0016 0.1061 5.782 0.029 2.49 8.96 5.626 0.061 50.39 0.54 3.9 6.3034 0.0073 0.1424 0.0293 1.9288 0.0886 0.0035 0.1703 5.413 0.03 4.15 1.21 5.578 0.183 49.97 1.62 4.5 6.0271 0.0065 0.1386 0.0247 1.9276 0.0197 0.0018 0.0545 5.408 0.029 3.53 7.62 5.721 0.049 51.23 0.43

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RS-07-84C Biotite J = 0.005034±0.000008 Laser P o w e r ( % ) 40Ar/39Ar 38Ar/39Ar 37Ar/39Ar 36Ar/39Ar Ca/K Cl/K %40Ar atm f 39Ar 40Ar*/39ArK Ag e 2 59.1754±0.0214 0.0857±0.1150 0.0576±0.2537 0.1926±0.0471 0.14 0.009 95.55 0.11 2.518±2.411 22.72±21.62 2.2 37.2090 0.0100 0.0651 0.1291 0.0412 0.1920 0.1141 0.0225 0.103 0.007 90.02 0.22 3.583 0.700 32.25 6.25 2.4 18.7189 0.0103 0.0532 0.0518 0.0279 0.0673 0.0470 0.0326 0.073 0.007 73.62 0.71 4.818 0.438 43.23 3.88 2.6 8.9157 0.0051 0.0462 0.0343 0.0047 0.0801 0.0127 0.0219 0.012 0.007 41.55 4.18 5.145 0.087 46.13 0.77 2.8 6.7390 0.0047 0.0441 0.0174 0.0023 0.1288 0.0047 0.0344 0.006 0.007 20.22 8.55 5.320 0.055 47.68 0.48 3 5.7569 0.0045 0.0468 0.0200 0.0020 0.0741 0.0013 0.0442 0.005 0.008 6.02 10.08 5.354 0.030 47.98 0.26 3.2 5.6253 0.0042 0.0462 0.0157 0.0024 0.0589 0.0008 0.0403 0.006 0.007 3.85 15.48 5.360 0.025 48.03 0.22 3.3 5.5782 0.0049 0.0461 0.0219 0.0024 0.0889 0.0006 0.1029 0.006 0.007 2.55 8.28 5.374 0.033 48.16 0.29 3.4 5.5862 0.0049 0.0443 0.0172 0.0030 0.0714 0.0006 0.0725 0.008 0.007 2.64 7.81 5.375 0.030 48.17 0.27 3.5 5.5971 0.0045 0.0451 0.0231 0.0038 0.0370 0.0007 0.0822 0.01 0.007 2.96 7.34 5.366 0.030 48.08 0.27 3.6 5.5726 0.0045 0.0464 0.0211 0.0054 0.0729 0.0007 0.0943 0.014 0.007 2.98 9.94 5.349 0.031 47.93 0.27 3.7 5.5908 0.0125 0.0456 0.0255 0.0078 0.0502 0.0007 0.0932 0.021 0.007 2.99 9.8 5.366 0.070 48.09 0.62 3.8 5.6306 0.0050 0.0501 0.0165 0.0485 0.0147 0.0008 0.0676 0.13 0.008 3.38 8.27 5.378 0.032 48.20 0.28 4 5.5804 0.0042 0.0531 0.0198 0.1250 0.0190 0.0006 0.0728 0.343 0.009 2.4 6.81 5.379 0.027 48.20 0.24 4.2 5.6446 0.0068 0.0538 0.0286 0.0915 0.0222 0.0010 0.2066 0.25 0.009 2.91 2.42 5.352 0.071 47.96 0.62

RS-07-84C Hornblende J = 0.005037±0.000010 Laser P o w e r ( % ) 40Ar/39Ar 38Ar/39Ar 37Ar/39Ar 36Ar/39Ar Ca/K Cl/K %40Ar atm f 39Ar 40Ar*/39ArK Ag e 2 394.6646±0.0292 0.3476±0.0919 1.6473±3.5264 0.9341±0.0424 4.55 0.039 69.01 0.12 122.011±9.652 864.28±54.33 2.2 465.9757 0.0154 0.3230 0.1360 0.3236 4.9397 1.2383 0.0272 0.864 0.019 78.34 0.22 100.859 8.545 741.12 51.49 2.5 364.0489 0.0094 0.2723 0.0465 0.5325 2.6914 1.0764 0.0200 1.454 0.013 87.41 0.48 45.733 5.765 373.99 42.58 2.7 144.7338 0.0262 0.1359 0.2605 0.5674 2.8837 0.4245 0.0314 1.552 0.01 86.53 0.66 19.329 2.292 167.61 18.98 2.9 27.8047 0.0076 0.1347 0.0316 2.5577 0.0583 0.0650 0.0207 7.042 0.025 67.48 2.9 8.927 0.385 79.35 3.35 3.1 12.2435 0.0055 0.1262 0.0123 2.3933 0.0139 0.0203 0.0216 6.586 0.025 46.5 20.52 6.507 0.134 58.18 1.18 3.3 7.9000 0.0061 0.1027 0.0190 2.1316 0.0186 0.0078 0.0222 5.864 0.02 24.92 14.1 5.858 0.064 52.46 0.56 3.4 9.6192 0.0105 0.1033 0.0236 2.1042 0.0362 0.0132 0.0377 5.789 0.02 36.56 6.16 5.982 0.160 53.56 1.41 3.6 11.3908 0.0050 0.1111 0.0171 2.3395 0.0224 0.0174 0.0246 6.438 0.022 42.05 11.83 6.535 0.133 58.43 1.17 3.8 7.3317 0.0062 0.1041 0.0174 2.3004 0.0141 0.0062 0.0368 6.33 0.021 20.63 27.29 5.770 0.077 51.69 0.68 4 6.1613 0.0068 0.0991 0.0343 2.1827 0.0267 0.0029 0.0470 6.006 0.02 7.68 8.74 5.564 0.057 49.87 0.51 4.5 6.3510 0.0078 0.0902 0.0233 1.9921 0.0301 0.0034 0.0901 5.48 0.017 9.58 6.98 5.594 0.102 50.13 0.90

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