Two-stage exhumation of midcrustal arc rocks, Coast Mountains, British Columbia

Transcription

1 TECTONICS, VOL. 24,, doi: /2004tc001750, 2005 Two-stage exhumation of midcrustal arc rocks, Coast Mountains, British Columbia M. E. Rusmore Department of Geology, Occidental College, Los Angeles, California, USA G. J. Woodsworth Geological Survey of Canada, Vancouver, British Columbia, Canada G. E. Gehrels Department of Geosciences, University of Arizona, Tucson, Arizona, USA Received 4 October 2004; revised 11 April 2005; accepted 8 July 2005; published 27 October [1] New geologic data from the Central Gneiss Complex along Douglas Channel help delineate the burial and exhumation of the Late Cretaceous to Eocene Coast Mountains magmatic arc. Arc plutonism was on going between 90 and 60 Ma, replaced by formation of dikes to 52 Ma. Supracrustal rocks were buried to midcrustal levels by 90 Ma, and garnet and kyanite grew from 90 through 70 Ma. Subsequent exhumation was nearly isothermal and took place in two stages. The first occurred from 70 to 59 Ma as the arc contracted obliquely and sillimanite replaced kyanite. Exhumation was slow (0.5 mm/yr), probably accomplished by erosion aided by coaxial crustal thinning. Exhumation rates about doubled in the second stage of exhumation, after 59 Ma and before 52 Ma. Penetrative deformation ended prior to intrusion of the Quottoon pluton at 59 Ma and 6.5 kbar. Production of cordierite rims on garnet at 4.5 kbar and 700 C signaled the end of nearisothermal decompression prior to 52 Ma. Rapid cooling (100 C/10 6 years) of the Central Gneiss Complex followed cordierite growth; temperatures dropped to 250 by 48 Ma. The increase in exhumation rate and the subsequent rapid cooling are attributed to excision of >6 km of crust on a detachment system on the northeastern side of the Central Gneiss Complex. Comparison to other parts of the Coast Mountains arc, the Sierra Nevada, and Fiordland, New Zealand, shows that the amount and tempo of exhumation vary greatly within and between arcs, suggesting that the processes accommodating exhumation vary significantly. Citation: Rusmore, M. E., G. J. Woodsworth, and G. E. Gehrels (2005), Two-stage exhumation of midcrustal arc rocks, Coast Mountains, British Columbia, Tectonics, 24,, doi: /2004tc Copyright 2005 by the American Geophysical Union /05/2004TC Introduction [2] The exhumed midcrustal roots of continental arcs are ideal laboratories in which to study of the processes of arc development and exhumation. The interplay of plate kinematics and crustal behavior are reflected in the magmatism, deformation, and metamorphism of the arc roots, and the processes driving arc formation and exhumation can be tested against the history of these rocks. Models for exhumation have been developed for collisional orogens; it is unclear how these apply to the deep levels of arcs where crustal rheologies, heat flow, and magmatism may differ significantly from collision settings [e.g., Klepeis et al., 2003; Paterson et al., 2004]. Critical questions that revolve around the exhumation of arc roots include whether major exhumation is ultimately driven by gravitational collapse of overthickened crust [e.g., Andronicos et al., 2003; Coney and Harms, 1984; Paterson et al., 2004], shifting plate configurations [e.g., Crawford et al., 1999; Klepeis et al., 2003; Saleeby, 2003], or more complex crustal dynamics and rheologic influences such as root delamination and crustal flow [e.g., Miller and Paterson, 2001; Saleeby et al., 2003; Teyssier and Whitney, 2002; Wernicke and Getty, 1997]. The roles of magmatism [Andronicos et al., 2003; Crawford et al., 1999] and erosion in the exhumation in arcs are also debated [e.g., Farley et al., 2001; Riebe et al., 2001; Small and Anderson, 1995; Spotila et al., 2004]. The timing and pace of exhumation in arcs are a critical factor in differentiating the contributions of these varied mechanisms. [3] Testing of models of arc formation and exhumation is hindered by the relatively few studies of well-preserved exhumed continental arcs; however, recent works in Fiordland, New Zealand, and Cordilleran arcs of North America show intriguing variations in exhumation history and inferred processes. Fiordland exposes deep arc rocks, recording rapid burial to >40 km [Daczko et al., 2001, 2002; Flowers et al., 2005; Hollis et al., 2004]. Exhumation lagged behind deep crustal burial by m.y., during which nearly isobaric cooling may have affected the deep crust for the first m.y. [Flowers et al., 2005]. This late orogenic exhumation is interpreted as the result of plate interactions [Klepeis et al., 2003]. In contrast, much of the 1of25

2 Figure 1. Location of study area. (a) Extent of Coast Belt and major Tertiary faults. (b) Regional geologic map showing location of Figure 2. Modified from Andronicos et al. [2003], Gareau [1991a], Hutchison [1982], Roddick [1970], and Woodsworth et al. [1985]. Sierra Nevada batholith has not been deeply exhumed. Most plutons record emplacement pressures of 1 3 kbar, although an oblique crustal section exposes rocks metamorphosed at 8 9 kbar (summarized by Saleeby et al. [2003]). Exhumation of these deeper rocks is attributed to slab segmentation leading to upper mantle removal and subsequent tectonic underplating of accretionary wedge and/or forearc material shortly after deposition [Grove et al., 2003; Saleeby, 2003]. [4] Deep crustal rocks are also exposed the northern part of the Cordillera arc, the Coast Plutonic Complex (Figure 1). Like the Sierra, the orogen seems to preserve varied levels of the arc [Woodsworth et al., 1991], and in contrast to Fiordland, deeply exhumed sections record nearly isobaric decompression. The southern end of the Coast Plutonic Complex, the North Cascades, exposes metamorphic rocks and plutons recording pressures of as high as kbar [Brown and Walker, 1993; Paterson et al., 2004; Valley et al., 2003; Wernicke and Getty, 1997; Whitney, 1992; Whitney et al., 1999]. Exhumation took place between about 90 and 45 Ma with an average rate of 1.6 mm/yr [Paterson et al., 2004]. Early isothermal decompression was followed by significant cooling with rates as high as 100 C/yr [Paterson et al., 2004; Valley et al., 2003; Wernicke and Getty, 1997; Whitney, 1992; Whitney et al., 1999]. Exhumation began during arc construction in the Late Cretaceous; thrust faulting and erosion are the suggested exhumation mechanisms [Paterson et al., 2004]. Further exhumation in the early Tertiary was coeval with extensional faulting. Exhumation is linked to deep crustal flow [Wernicke and Getty, 1997] or the gravitational collapse of overthickened crust, complicated by variable 2 of 25

3 Figure 2. (a) Geology of the Douglas Channel area, based on the work by Roddick [1970] and our work. F, Foch Lagoon; M, Miskatla Inlet. U-Pb ages from 95MR-07 are from Rusmore et al. [2001] (b) Cross section of Douglas Channel area. Axial planes of F 2 folds shown schematically with dashes. stretching/thickening, erosion, and underthrusting along the arc [Paterson et al., 2004]. [5] A broadly similar history is seen in the Central Gneiss Complex, 700 km northwestward along the arc. Located in the core of the arc, the Central Gneiss Complex exposes 10,000 km 2 of granulite to amphibolite-facies metamorphic rocks which cooled rapidly in the early Tertiary. Work has centered on the northern half of the Central Gneiss Complex, from around the Skeena River north to Portland Inlet (Figure 1). Fundamental results bearing on the formation and exhumation of the arc include documentation of an early phase of crustal thickening and high-pressure metamorphism followed by rapid decompression and cooling of the core in early Tertiary time, exhumation on ductile shear zones parallel and oblique to the orogen, and a complex interplay between magmatism and deformation [Andronicos et al., 2003, 1999; Crawford and Hollister, 1982; Crawford et al., 1987; Hollister, 1982; Klepeis and Crawford, 1999; Klepeis et al., 1998; McClelland et al., 1991]. Most studies reveal a fairly consistent metamorphic history, but the deformation that accompanied metamorphism is highly variable and difficult to integrate into a coherent kinematic picture. Particularly enigmatic is the extent and significance of dextral shearing recognized near the Skeena River [Andronicos et al., 1999; Hollister and Andronicos, 1997; Klepeis and Crawford, 1999]. Although significant exhumation is a hallmark of the Central Gneiss Complex, how this exhumation was accomplished is not clear. Hollister [1982] suggested crustal delamination drove exhumation; crustal extension and channeled return flow of magma was suggested by Andronicos et al. [2003]. Other studies emphasize the role of plutons and crustal anisotropies in localizing shear zones and suggest tilting is responsible for the different crustal levels now exposed [Butler et al., 2001; Cook and Crawford, 1994; Crawford et al., 1999; Klepeis and Crawford, 1999]. [6] Prompted by these uncertainties, we completed an integrated geochronologic, structural, and petrologic study of the little known southern part of the Central Gneiss Complex (Figures 1 and 2). Our goals in this paper are (1) to 3of25

4 integrate our results from the southern Central Gneiss Complex with prior results from farther north to understand the development of the varied structures during arc formation and exhumation and (2) to define the timing, magnitude, and rate of exhumation of the southern area to help distinguish between models of arc exhumation. We show that slow exhumation began during arc construction and subsequently accelerated during extensional faulting. Viewed in total, the heterogeneous deformation of the Central Gneiss Complex most likely reflects local and probably transient strain partitioning within roots of the arc. We infer that plate kinematics likely triggered the faster phase of exhumation, although other factors worked to spatially focus exhumation. Comparison to Fiordland, the Sierra Nevada, and the North Cascades highlights features that may be used to distinguish between exhumation models. 2. Overview of the Central Gneiss Complex [7] The Central Gneiss Complex is composed of complexly deformed amphibolite to granulite facies orthogneiss, gray gneiss, pelitic gneiss, amphibolite, with rare quartzite, marble, and small ultramafic bodies [Douglas, 1986; Hollister and Andronicos, 2000; Hutchison, 1982; Roddick, 1970]. Early Tertiary high-grade metamorphism and cooling are the signature of the Central Gneiss Complex and serve to differentiate it from adjacent rocks. [8] Along its southwestern side, the Central Gneiss Complex is intruded by the Paleocene Quottoon pluton [Roddick, 1970; Thomas and Sinha, 1999; Rusmore et al., 2001; Crawford et al., 1999; Andronicos et al., 1999] (Figure 2). The Quottoon is truncated by the Coast shear zone, a >1200-km-long shear zone dominated by reverse (eastside up) motion between about 65 and 55 Ma [Crawford and Hollister, 1982; Crawford et al., 1987; Gehrels and Boghossian, 2000; Gehrels et al., 1991a, 1991b; Ingram and Hutton, 1994; Klepeis et al., 1998; Rusmore et al., 2001; Stowell and Hooper, 1990; Wood et al., 1991]. West of the Coast shear zone is a welldeveloped mid-cretaceous thrust belt that imbricates rocks ranging from amphibolite facies in the east to subgreenschist strata of Wrangellia on the west [Crawford and Hollister, 1982; Crawford et al., 1987; Gehrels et al., 1992; Hutchison, 1982; McClelland et al., 1992; Rubin and Saleeby, 1992; Rubin et al., 1990]. On Douglas Channel the amphibolite gneiss west of the Coast shear zone is referred to as the Scotia-Quaal gneiss [Friedman et al., 2001; Gareau, 1991]. Sparse K-Ar biotite ages show that the Scotia-Quaal gneiss cooled to C by Ma [Roddick, 1970]. These ages contrast sharply with Ma biotite Ar-Ar ages in the Central Gneiss Complex [Andronicos et al., 2003; Roddick, 1970]. [9] An Eocene detachment system forms the northeastern boundary of the Central Gneiss Complex near Terrace (Figure 1) [Andronicos et al., 2003; Heah, 1990, 1991]. The southern continuation of this detachment is discussed in this paper. Stikinia composes the upper plate of the detachment system, and these weakly metamorphosed Mesozoic strata contrast sharply with the rocks of the Central Gneiss Complex. Stikinia, like the rocks on the western flank to the Coast Mountains, underwent Late Cretaceous thrust faulting. Remnants of this northeast vergent thrust belt are preserved locally east of the Coast Mountains [Evenchick, 1991; Heah, 1990; Journeay and Friedman, 1993; Rusmore and Woodsworth, 1991]. Undeformed Eocene plutons are common in Stikinia adjacent to the Central Gneiss Complex, and early Tertiary strata are only locally present [Gareau et al., 1997a, 1997b]. Unlike the southwestern and northeastern boundaries of the Central Gneiss Complex, the northern and southern boundaries are not well defined, but deeply exhumed Eocene metamorphic rocks are not exposed in the orogen core between Bella Coola and the North Cascades [Woodsworth et al., 1991]. 3. Evolution of the Central Gneiss Complex on Douglas Channel [10] Our study of the southern Central Gneiss Complex focused on the magmatic, structural, and metamorphic history of rocks along Douglas Channel (Figure 2). These results are presented in the next sections and later compared to the rest of the Central Gneiss and results from other deeply exhumed continental arc roots Magmatism [11] Orthogneiss constitutes more than half the Central Gneiss Complex around Douglas Channel and is interpreted as the product of midcrustal magmatism associated with the Coast Mountains arc [e.g., Armstrong, 1988; Gehrels et al., 1991b; Woodsworth et al., 1991]. This record of magmatism provides critical constrains on the environment and age of deformation and metamorphism. Five new U-Pb ages show that magmatism was nearly continuous from 90 to 67 Ma and was followed by synchronous cooling of all samples at 52 Ma. These dates are described here, from oldest to youngest, and integrated with the deformational and metamorphic history of the arc later in this paper. Analytical methods are described in Appendix A; analytical data are presented in Tables 1 and 2 (also see the auxiliary material 1 ). All ages are reported at the 2s level unless otherwise noted. [12] The samples are from sill-shaped orthogneiss bodies that were deformed and metamorphosed with the surrounding gneisses (Figure 2). The two oldest ages are characterized by complex zircon isotopic systems (Figures 3a 3d). To resolve complex discordance in these samples, we conducted analyses along transects along the length of large zircon grains. Ages for each spot, arranged from tip to tip, are reported in Table 2 and shown graphically on Figure 3. [13] In sample 95MR-24, isotope dilution-thermal ionization mass spectrometry (ID-TIMS) analyses overlap concordia but range from 90 to 67 Ma (Figure 3a). Laser analyses across nine grains yield a similar spread in ages, 1 Auxiliary material is available at ftp://ftp.agu.org/apend/tc/ 2004TC of25

5 Table 1. U-Pb Isotopic Data and Ages a Grain Type Grain Weight, mg Pb c, pg U, ppm 206 Pb m / 204 Pb 206 Pb c / 208 Pb 206 Pb*/ 238 U 207 Pb*/ 235 U 206 Pb*/ 238 U 207 Pb*/ 235 U 206 Pb*/ 207 Pb* 96MR-13 Z3A ± ± ± 34 Z1A ± ± ± 71 Z5C ± ± ± 114 Z20E ± ± ± 37 Z10C ± ± ± 32 S5A ± ± ± 240 S8A ± ± ± MR-17 Z5A ± ± ± 27 Z1A ± ± ± 59 Z10D ± ± ± 18 Z30G ± ± ± 23 Z1A ± ± ± 17 Z1A ± ± ± 16 Z1A ± ± ± 16 S10A ± ± ± 270 S10A ± ± ± 360 S4A ± ± ± 190 S6A ± ± ± MR-22 S5A ± ± ± 280 S5A ± ± ± MR-24 Z8A ± ± ± 25 Z1A ± ± ± 22 Z1A ± ± ± 13 Z10C ± ± ± 14 Z30G ± ± ± 14 Z1Ba ± ± ± 13 Z1Ba ± ± ± 12 Z1Ba ± ± ± 12 Z5D ± ± ± 16 Z1Ba ± ± ± 20 Z1Ba ± ± ± 13 Z1Ba ± ± ± 14 Z1Ba ± ± ± 16 Z1Ba ± ± ± 15 Z1Ba ± ± ± 15 Z1Ba ± ± ± 12 Z1Ba ± ± ± 16 Z1Ba ± ± ± 17 Z1Ba ± ± ± 19 S16A ± ± ± 130 S5A ± ± ± 100 S5A ± ± ± 71 95MR-26 Z10A ± ± ± 10 Z1A ± ± ± 16 Z1A ± ± ± 20 Z10C ± ± ± 10 Z30G ± ± ± 20 Z1A ± ± ± 20 Z1A ± ± ± 19 Z3A ± ± ± 12 Z3A ± ± ± 12 S10A ± ± ± 70 S5A ± ± ± 67 S5A ± ± ± 73 97MR-28 Z10D ± ± ± 19 Z8C ± ± ± 20 5of25

6 Table 1. (continued) Grain Type Grain Weight, mg Pb c, pg U, ppm 206 Pb m / 204 Pb 206 Pb c / 208 Pb 206 Pb*/ 238 U 207 Pb*/ 235 U 206 Pb*/ 238 U 207 Pb*/ 235 U 206 Pb*/ 207 Pb* Z1A ± ± ± 38 Z5B ± ± ± 15 Z10D ± ± ± 29 Z5C ± ± ± 62 Z1A ± ± ± 26 Z1A ± ± ± 21 97MR-29 Z10B ± ± ± 11 Z20D ± ± ± 19 Z20D ± ± ± 12 Z20D ± ± ± 10 Z35F ± ± ± 14 Z1Ba ± ± ± 13 Z1Ba ± ± ± 18 Z1Ba ± ± ± 10 Z1Ba ± ± ± 12 97MR-40 Z4A ± ± ± 15 Z4A ± ± ± 7 Z10B ± ± ± 10 Z13C ± ± ± 15 Z30G ± ± ± 10 Z1Ba ± ± ± 19 Z1Ba ± ± ± 11 Z1Ba ± ± ± 15 Z1Ba ± ± ± 10 96MR-47 Z5A ± ± ± 15 Z10C ± ± ± 14 Z10C ± ± ± 23 Z20D ± ± ± 11 Z10D ± ± ± 20 S5A ± ± ± 330 S8A ± ± ± MR-56 Z1A ± ± ± 29 Z1A ± ± ± 23 Z1A ± ± ± 27 Z1A ± ± ± 15 Z1A ± ± ± 20 Z1A ± ± ± 24 Z1A ± ± ± 15 a Thermal ionization mass spectrometry (TIMS). Grain sizes are A = 200 mm, B = 175 mm, C = 150 mm, D = 125 mm, E = 100 mm, F = 80 mm, G = 60 mm; Z, zircon; S, sphene; a, grains abraded in air abrasion device. Number in grain type column indicates the number of crystals analyzed; 206 Pb/ 204 Pb is measured ratio, uncorrected for blank, spike, or fractionation; 206 Pb/ 208 Pb is corrected for blank, spike, and fractionation. All uncertainties are at the 95% confidence level. Uncertainties in isotope ratios are in percent. Uncertainties in ages are in millions of years. Most concentrations have an uncertainty of 25% due to uncertainty in weight of grain. Constants used are 238 U/ 235 U = Decay constant for 235 U is Decay constant for 238 U is Isotope ratios are adjusted as follows: (1) mass-dependent corrections factors of 0.14 ± 0.06%/amu for Pb and 0.04 ± 0.04%/amu for UO2; (2) Pb ratios corrected for ± ng blank with 206 Pb/ 204 Pb = 18.6 ± 0.3, 207 Pb/ 204 Pb = 15.5 ± 0.3,and 208 Pb/ 204 Pb = 38.0 ± 0.8; (3) U has been adjusted for ± ng blank; and (4) initial Pb from Stacey and Kramers [1975], with uncertainties of 1.0 for 206 Pb/ 204 Pb, 0.3 for 207 Pb/ 204 Pb, and 2.0 for 208 Pb/ 204 Pb. All analyses conducted using conventional isotope dilution and thermal ionization mass spectrometry, as described by Gehrels [2000]. although analyses from central portions of the grains cluster at 90 Ma and analyses from grain tips at 55 Ma (Figure 3b). The older grain interiors probably record crystallization, whereas the tips likely record a younger episode of zircon growth. On the basis of higher U/Th in most tip analyses than in core analyses (Figures 4 and 5), this younger growth is probably metamorphic [Rubatto et al., 2001; Williams, 2001]. The weighted means of the two clusters (excluding outliers shown as gray vertical bars on Figure 3b) yield ages of 88.3 ± 3.0 Ma and 56.2 ± 1.9 Ma. We interpret these ages as recording crystallization at 88 Ma, and later renewed growth during metamorphism 6of25

7 Table 2. U-Pb (Zircon) Geochronologic Analyses by Laser-Ablation Multicollector ICP Mass Spectrometry a Isotopic Ratios Apparent Ages Sample Grain U, ppm 206 Pb/ 204 Pb U Th 207 Pb*/ 235 U % Error, Error, 206 Pb*/ 238 U % Error Correction Error, 206 Pb*/ 238 U m.y. Error, 207 Pb*/ 235 U m.y. Error, 206 Pb*/ 207 Pb* m.y. 96MR MR of25

8 Table 2. (continued) Isotopic Ratios Apparent Ages Sample Grain U, ppm 206 Pb/ 204 Pb U Th 207 Pb*/ 235 U % Error, Error, 206 Pb*/ 238 U % Error Correction Error, 206 Pb*/ 238 U m.y. Error, 207 Pb*/ 235 U m.y. Error, 206 Pb*/ 207 Pb* m.y a All errors are reported at the 1-sigma level and incorporate only uncertainties from measurement of isotopic ratios. U concentration and U/Th have uncertainty of 25%. Decay constants are 235 U = U = , 238 U/ 235 U = Isotope ratios are corrected for Pb/U fractionation by comparison with standard zircon with an age of 564 ± 4 Ma. Initial Pb composition is interpreted from Stacey and Kramers [1975], with uncertainties of 1.0 for 206 Pb/ 204 Pb and 0.3 for 207 Pb/ 204 Pb. 8of25

9 Figure 3 9of25

10 Figure 4. Plot of U/Th as a function of 206 Pb*/ 238 U ages from sample 95MR-24. The occurrence of high U/Th in the rims is consistent with the interpretation that the older age (88 Ma) records igneous crystallization, whereas the younger age (56 Ma) records younger metamorphism [e.g., Rubatto et al., 2001; Williams, 2001]. at 56 Ma. Analysis of three fractions of sphene grains yields an age of 51.0 ± 1.5 Ma (Figure 3a). The low precision of this age is a result of the low U and high initial Pb concentration in these grains (Table 1). [14] A similar crystallization age was obtained from 96MR-56. ID-TIMS analyses of individual abraded zircon crystals overlap concordia, but range from 77 to 89 Ma (Figure 3c). Laser analyses along transects across three grains produced 206 Pb*/ 238 U ages ranging from 83 to 96 Ma, with no apparent zonation in age (Figure 3d). The lack of age zonation suggests that the central portions of the grains are undisturbed, and that the discordance seen in the TIMS analyses must be restricted to thin rims that were not entirely removed by abrasion. The weighted mean of the spot analyses (following rejection of one outlier) is 87.2 ± 1.8 Ma, which yields an interpreted crystallization age of 87.2 ± 3.1 Ma when systematic errors are included. [15] Younger and simpler ages were obtained from three other orthogneiss bodies (Figures 3e and 3f). Zircons from sample 95MR-17 (Figure 3e) show six of the seven analyses concordant at 77.5 ± 2.0 Ma. The seventh is slightly older, presumably due to inheritance of older zircon. Three sphene fractions yield a cooling age of 52.8 ± 1.5 Ma, with a fourth analysis that is significantly younger for unknown reasons. A crystallization age of 68 ± 1 Ma was obtained from five concordant zircon fractions from sample 96MR-47 (Figure 3f). Two sphene fractions yield a cooling age of 52.4 ± 1.5 Ma. Similar ages were obtained by TIMS analyses of five zircon fractions from sample 96MR-13. The analyses are of fairly low precision due to the low U concentration of the zircon grains, but the resulting ages are concordant at 67.0 ± 1.5 Ma (Figure 3g). Two sphene fractions yield a cooling age of 52.7 ± 1.5 Ma (Figure 3g). [16] These new crystallization ages show that this part of the Central Gneiss Complex was the locus of fairly continuous magmatism from 90 to 65 Ma. Although similar to those elsewhere in the arc, the ages confirm that magmatism was a significant process during formation of the Central Gneiss Complex along Douglas Channel. As discussed below, intrusion of the orthogneiss took place during prolonged deformation and metamorphism at midcrustal levels and likely provided heat during isothermal decompression of the Central Gneiss. Previous work and results discussed below show that this history is not shared by Paleocene and younger plutons and dikes, limiting the ages of penetrative deformation of the Central Gneiss Complex. The crystallization and cooling ages also constrain the timing and rate of exhumation of the Central Gneiss Complex Deformation [17] Four main structural events (D 1 D 4 ) are recorded in the Central Gneiss Complex on Douglas Channel. These structures are intruded by the 59 Ma Quottoon pluton in the southwest and truncated by a ductile shear zone, referred to here as the eastern boundary detachment, in the northeast. The structures and the Quottoon pluton have been have been reoriented by post-59 Ma vertical axis rotations during formation of a crustal-scale bend in the orogen, referred to as the Hawkesbury Warp (Figure 1b) [Bogue et al., 1999; Roddick, 1970; Symons, 1977] and local small-block tilting [Bogue et al., 1999]. To clarify the Late Cretaceous orientation of structures in the Central Gneiss Complex, stereonets show both the present orientation and one in which the vertical axis rotation has been removed. The small-block tilting is only detectable in the limited areas where paleomagnetic data are available, and so is not removed First Phase (D 1 ) [18] A prominent gneissic foliation (S 1 ) characterizes all rock types in the Central Gneiss Complex around Douglas Channel and is parallel to almost all lithologic contacts within the Central Gneiss Complex. Throughout the Douglas Channel area, S 1 defines a diffuse fold pattern with a shallow NW trending hinge line prior to removal of the vertical rotation of the Hawkesbury Warp (Figure 6). Figure 3. Pb/U concordia diagrams for samples from the Central Gneiss Complex along Douglas Channel. Locations are shown on Figure 2. Error ellipses for the ID-TIMS data are at the 2s level, whereas the ICPMS error bars are shown at the 1s level. (a) ID-TIMS analyses of abraded single zircon grains and of sphere grains in orthogneiss 95MR-24. (b) The 206 Pb*/ 238 U ages of laser-icpms analyses (nine grains) in 95MR-24. (c) ID-TIMS analyses of abraded single zircon grains in orthogneiss 96MR-56. (d) The 206 Pb*/ 238 U ages of laser-icpms analyses (three grains) orthogneiss 96MR-56. (e) Orthogneiss 96MR-17. (f ) Othogneiss 96MR-47. (g) Orthogneiss 96MR-13. (h) Postkinematic dike 95MR-26. (i) Postkinematic dike 97MR-28 in eastern boundary detachment. (j) Sphene age from tonalite 95MR of 25

11 Figure 5. Cathodoluminescence image of grain 4 from sample 95MR-24, showing the position of laser pits across the grain. Each laser pit has a diameter of 25 mm and a depth of 20 mm. Ages shown are 206 Pb*/ 238 U ages with uncertainties at 1s. This hinge line is subparallel with folds formed during both D 2 and D 3, and the overall S 1 pattern reflects these phases of deformation. Undoing these later events would restore S 1 to a roughly horizontal original orientation. [19] Rare intrafolial, rootless isoclinal folds (F 1 ) were found within pelitic to psammitic gneisses. The folds are generally less than half a meter in amplitude and their axial planes are parallel to S 1. Hinge lines trend NE to NW, and locally the folds display type 2 and 3 refolding patterns where overprinted by D 2 folds. Whether the D 1 folds are relics of transposition of an older foliation into S 1 or the product of local perturbations during progressive deformation is unclear. The rarity of the folds and the general continuity of layers over many meters are atypical of transposed fabrics, but the intensity of the later deformation and metamorphism may have reconstituted much of the foliation and obscured earlier transposition fabrics. Regardless of its origin, we refer to the prominent, pervasive metamorphic foliation as S 1 for simplicity and conformity with use elsewhere in the Central Gneiss Complex [Andronicos et al., 1999; Hill, 1984; Hollister and Andronicos, 2000; Selverstone and Hollister, 1980] Second Phase (D 2 ) [20] The second phase of deformation produced widespread folds (F 2 ) with local axial planar foliation (S 2 ), boudins, and a domainal mineral elongation lineation (L 2 ). F 2 folds are typically close to tight folds with a wavelength of 5 10 m and an axial plane subparallel to S 1. Hinge lines plunge gently to the northwest and southeast (Figure 6b). Vergence patterns suggest that the outcrop-scale folds are parasitic to kilometer-scale folds. The axial planar foliation (S 2 ) is present throughout the area mapped, but is only sparsely developed in any one area. S 2 is most easily recognized in pelitic gneisses where it is defined by biotite, fibrolite, and sillimanite crystals up to 1 cm long. Because S 2 is subparallel to S 1, some foliations were probably misidentified in areas where the relation to F 2 is not visible. For clarity, only visibly axial planar S 2 measurements are shown on Figure 6d. Like S 1,S 2 is folded into a large NW trending antiform (F 3 ). [21] The mineral elongation lineation (L 2 ) is developed in all rock types but is only present in local domains. Documentation of the three-dimensional geometry of the domains is limited by exposure and access. Within domains, L 2 is common and parallel to F 2 hinge lines (Figures 6b and 6c). We interpret the lineated domains as zones of higher shear strain, perhaps on limbs of large recumbent folds or in diffuse ductile shear zones. [22] Well-defined shear zones were not observed within the Central Gneiss Complex we mapped. A single outcrop of mylonitic sillimanite gneiss between mafic gneiss may mark a shear zone but was not traceable laterally. Discrete shear bands, many melt filled, occur throughout the gneiss, but the kinematics and timing of these shears is difficult to relate to the penetrative structures. Where the slip direction on several of these shear bands could be determined, we found no consistent sense of shear or relationship to the regional lineation. This lack precluded using the regional lineation direction to interpret the kinematics of the shear bands. The age of the shear bands also varies; some lack melts, others are melt filled, while others crosscut melts, and some clearly postdate D 2 folds and other structures. Giving this range of relative ages, determining the age of any given shear band is difficult. Most likely, the shear bands developed throughout the formation and cooling of the Central Gneiss Complex and record a variety of slip directions. [23] Boudins of mafic layers and felsic veins are ubiquitous in the gneisses and orthogneisses. Most lie in S 1 with their necks perpendicular to L 2. The boudins record shortening perpendicular to S 1 and highly variable amounts of elongation, ranging from 75 to 200% parallel to L 2 where measured in several locations. The boudin orientation suggests they formed during the elongation recorded by D 2 lineations and hinge lines, and were folded with S 1 and S 2 during D 3. Unfolding the large D 3 antiform and vertical axis rotations shows that strain during D 2 included significant NE-SW elongation accompanied by shortening along a nearly vertical strain axis. Figure 6. Equal-area stereonet plots of structures in the Central Gneiss Complex. (left) Present orientation of structures; (right) structures rotated to remove effects of Hawkesbury Warp. (a) Poles to S 1,(b)F 2 hinge lines, (c) L 2 elongation lineations, (d) poles to S 2 and F 2 axial planes, (e) eastern boundary detachment: poles to mylonitic foliation and mylonitic elongation lineation. 11 of 25

12 Figure 6 12 of 25

13 [24] Collectively, the structures formed during D 2 suggest it represents a period of intense, nearly recumbent folding accompanied significant vertical flattening and NE-SW elongation. Folding, rather than development of large discrete shear zones characterized the deformation. Major flattening and elongation that occurred during the deformation probably resulted from crustal thickening produced by the folding, and perhaps thrusting present outside the area [Andronicos et al., 2003, 1999; Crawford et al., 1987, 1999; Douglas, 1986; Hollister and Crawford, 1986]. D 2 is the final phase of deformation that produced penetrative structures visible at the outcrop scale Third Phase (D 3 ) [25] The third phase of deformation produced a crustalscale upright antiform that affected the entire Central Gneiss Complex along Douglas Channel. Smaller structures associated with D 3 were not observed. The antiform is best defined by systematic deflection of S 2 and F 2 axial planes across the Central Gneiss Complex, with generally gentle dips in the central area defining the hinge region (Figure 6d). An exception to this pattern occurs around Miskatla Inlet where S 2 dips steeply northeast; this is probably the northeast limb of an incompletely mapped minor fold. The F 3 hinge line trends northwest and is horizontal (Figure 6d); because S 1 and S 2 are subparallel, a similar pattern is visible in equal-area plots of S 1 (Figure 6a). The antiform is truncated by the Quottoon pluton; therefore, like the Quottoon pluton, as well as D 1 and D 2 structures, the antiform has been rotated during formation of the Hawkesbury Warp. Removing the vertical axis rotation of the warp shows the antiform formed as a northeast (015 ) trending, upright horizontal fold. On the basis of simple line length calculations, the antiform accommodated about 20% shortening along an ESE-WSW trending horizontal axis. [26] The presence of a large antiform on Douglas Channel was first recognized by Roddick [1970], who called it the Foch Antiform. The southwest limb of the Foch Antiform as shown by Roddick [1970] is composed of rocks now recognized as mylonites within the Coast shear zone [Rusmore et al., 2001], whereas the fold we mapped is defined entirely by rocks of the Central Gneiss Complex. Despite this difference, we apply the name Foch Antiform to the large D 3 antiform as it generally coincides with that of Roddick [1970]. Other than minor folding near the hinge, no other structures appear to have formed in the Central Gneiss Complex during D Fourth Phase (D 4 ) [27] Scattered mesoscale folding marks a fourth phase of deformation in the Central Gneiss Complex. Although widespread, the folds are not common in any area. These small, upright, open folds have wavelengths of about a meter and can be seen folding L 2 in outcrop. Their axial planes are north striking and nearly vertical; hinge lines trend north and are horizontal. Though the data are sparse, the axial planes do not appear to be folded, suggesting the folds are younger than the Foch Antiform. Because they were not observed near the Quottoon pluton, their relation to it is not known and there is no younger limit on the age of their formation. Likewise, it is not known whether they formed before or after the Hawkesbury Warp, so their original orientation is uncertain. Given these limitations, D 4 structures are not discussed further in this paper Eastern Boundary Detachment [28] The eastern boundary of the Central Gneiss Complex is a ductile shear zone we call the eastern boundary detachment (Figures 1 and 2). Where examined on Douglas Channel the shear zone is 2 3 km wide, west striking and north dipping. The Central Gneiss Complex forms the lower plate with 67 Ma orthogneiss adjacent to the shear zone (Figure 2). The upper plate is mafic volcanic rocks and Carboniferous diorite (see auxiliary material) of Stikinia. In contrast to the Central Gneiss Complex, this diorite is weakly metamorphosed to subgreenschist facies. Away from the shear zone, primary intrusive textures are well preserved and the diorite is generally unfoliated. Mafic enclaves and xenoliths of mafic volcanic rocks are common. [29] The shear zone is marked by a progression from unfoliated diorite with near-equant mafic inclusions in the upper plate to strongly mylonitic rocks flanking the contact between the diorite and the Central Gneiss Complex. Approaching the shear zone from the north, the first sign of deformation is a gradual increase in flattening and elongation of the mafic inclusions. A penetrative foliation and lineation defined by crystal shapes appear 1 2 km north of the shear zone, along with localized mylonite zones a few meters wide. Nowhere is the contact itself exposed. It is covered by a 2-km-wide delta on the southeast side of the channel; mylonitic diorite is exposed to the north and orthogneiss crops out south of the delta. Undeformed granitoid plutons and dikes intrude and obscure the shear zone on the west shore of Douglas Channel, although mylonites derived from both the diorite and the lower plate orthogneiss are locally exposed. [30] The mylonitic foliation strikes west and dips north (Figure 6e). The lineation plunges to the northwest, slightly oblique to dip. Removing the rotation of the Hawkesbury Warp would produce a NW strike and a NE plunging lineation. Well-developed shear bands, and s-type porphyroclasts consistently show down lineation motion, indicating the shear zone had predominantly normal motion. The age of the shear zone is limited by the 67 Ma intrusive age of the orthogneiss and the age of an undeformed, white granite dike that crosses the shear zone. U-Pd dating of this dike (97MR-28) produced a crystallization age of 52 ± 3 Ma (Figure 3i) On the basis of these ages, the eastern boundary detachment was active between 67 and 52 Ma Metamorphism [31] Intense metamorphism accompanied the penetrative deformation of the Central Gneiss Complex. Pelitic gneisses provide the most useful assemblages. The oldest assemblage, preserved in garnet cores and present as relics in the matrix, is kyanite + biotite + garnet + quartz + plagioclase. Later than, and partially replacing, this assemblage is sillimanite + biotite + garnet + quartz + plagioclase ± K-feldspar. Cordierite + biotite + sillimanite + garnet ± K-feldspar as rims on garnets is the youngest (nonretrog- 13 of 25

14 Table 3a. Microprobe Analyses of Representative Minerals From Pelitic Rock for Garnet (12 Oxygen) a rade) assemblage observed. Primary muscovite + quartz have not been found. Unlike areas of the Central Gneiss Complex farther north [Crawford et al., 1987; Hollister, 1975; Sisson, 1985], staurolite has not been found in garnet cores or in matrix assemblages, and andalusite is absent. Fe-Ti oxides and accessory minerals are present in all rocks. Sample 95MR-19d 95MR-36 95MR-35 1 core 2 rim 3 core 4 rim 5 core 6 rim 7 core 8 rim 9 core 10 rim 11 core 12 rim SiO Al 2 O TiO Cr 2 O FeO MnO MgO CaO Total Si Al Ti Cr Fe Mn Mg Ca X Almandine X Pyrope X Grossular X Spessartine a Sample locations are shown on Figure 2 and in the auxiliary material. Values are in wt %. Analyses were done by wavelength dispersion methods on a Jeol Superprobe 733 electron microprobe at California Institute of Technology. Operating conditions included an accelerating potential of 15 kv, sample current of 10 na, and counting times of 40 s. Well-characterized natural and synthetic minerals were used as standards. Kakanui hornblende was monitored as a secondary standard. Conventional ZAF data reduction was done with the program CITZAF 3.03, written by J. T. Armstrong. [32] Garnet porphyroblasts reach 4 cm across and are subhedral to strongly embayed. The largest are commonly oblate in shape. Many garnet porphyroblasts have inclusionrich cores and nearly inclusion-free rims. The garnet cores contain kyanite inclusions, whereas scattered sillimanite inclusions occur in the garnet rims. Other inclusions in both Table 3b. Microprobe Analyses of Representative Minerals From Pelitic Rock for Biotite (22 Oxygen, Anhydrous) a 95MR-19d 95MR-36 95MR-35 95MR SiO Al 2 O TiO Cr 2 O FeO MnO MgO K 2 O Na 2 O Total Si Al Ti Cr Fe Mn Mg K Na a See Table 3a footnotes. 14 of 25