Modulation of Deformation by Magmatic Tempo, Coast Mountains Batholith, British Columbia, Canada

Abstract

The tectonomagmatic evolution of the southern Coast Mountains batholith, British Columbia reveals the relation of magmatic tempo, deformation, and relative plate motions which inform models for growth of continental crust and convergent plate dynamics. Magmatism in the southern batholith is episodic, derived primarily from the mantle, and reflects lower-plate dynamics (Cecil et al., 2018, https://doi.org/10.1029/2018gc007874, 2021, https://doi.org/10.1130/ges02361.1). New field, structural, geochronologic, and geobarometric results from the southern batholith document three episodes of deformation, each spatially and temporally coincident with a high magmatic flux event (HFE). Sinistral faulting (<117–103 Ma) and penetrative deformation (110–90 Ma) occurred only within two distinct areas affected by a HFE at 114–102 Ma. Following a magmatic lull, crustal shortening occurred after 90 Ma and before 72 Ma within the region affected by a second HFE (85–70 Ma). After 70 Ma and before 53 Ma, >100 km of dextral slip occurred on the Coast shear zone and minor shortening affected 64–62 Ma plutons, overlapping with the 64–61 Ma HFE. Thus, at the crustal level exposed in the southern batholith, the timing and location of deformation is linked to magmatic tempo and the style of deformation varies through time. These results suggest that magmatic HFE modulate the timing and location of deformation while relative plate motions during HFE influence the style of deformation. Periods of deformation in batholiths may thus record high-flux magmatic events and coeval plate motion but do not necessarily signal changes in plate motions.

Key Points

• The location and timing of deformation and high-flux magmatism coincide across the southern Coast Mountains batholith between ca. 115 and 50 Ma

• High-flux magmatism controls the timing and location of deformation; style of deformation reflects plate motion during high-flux magmatism

• Dextral slip exceeded 100 km on the intra-arc Coast shear zone (<70–53) Ma and caused Salinian-style repetition of the batholith

Plain Language Summary

Understanding how continents grow and the dynamics of plate motions relies on knowing the geologic history of convergent margins. Here we present new data on the evolution of a Late Cretaceous-Paleogene convergent margin preserved in the southern Coast Mountains batholith in British Columbia. Magmatism in the southern batholith is episodic, derived primarily from the mantle, and reflects lower-plate dynamics (Cecil et al., 2018, https://doi.org/10.1029/2018gc007874, 2021, https://doi.org/10.1130/ges02361.1). Our new results show that the southern batholith experienced three episodes of deformation, each one coincident with one of three high-magmatic-flux events. This pattern suggests that the timing and location of deformation here is linked to magmatic tempo and the style of deformation varies independently. Thus, we conclude that magmatic patterns modulate the tempo and location of deformation and that relative plate motions during episodes of high-flux magmatism control the style of deformation. Periods of deformation in batholiths may thus record high-flux magmatic events and coeval plate motion but do not necessarily signal changes in plate motions.

1 Introduction

Cordillera batholiths record the dynamics of ancient convergent continental margins and provide a window into their evolution. The growth of batholiths is commonly linked to first-order tectonic factors such as plate convergence angle and velocity, mechanical coupling, and slab dip, and their magmatic and deformational history underpins models of ancient plate interactions (e.g., DeCelles et al., 2009, 2015; Glazner, 1991; Hamilton, 1988; Holdaway & Mukhopadhyay, 1993; Hollister & Andronicos, 2006; Horton, 2018; Maloney et al., 2013; Monger & Gibson, 2019; Ramos, 2010; Tobisch et al., 1995). Magmatic characteristics of batholiths help define the size and polarity of ancient arcs and track lithospheric removal beneath the arc (e.g., Cecil et al., 2012, 2018; Chapman et al., 2017; Gehrels et al., 2009; Hollister & Andronicos, 2006; Lackey et al., 2012). As the locus of high heat flow, arcs may focus deformation related to plate interactions and the record of ancient plate boundaries is commonly inferred from the deformation of batholiths. Widespread thrust faulting and crustal shortening in batholiths is interpreted as a signal of predominantly orthogonal subduction and terrane accretion, whereas the development of intra-batholithic transcurrent faults is attributed to increased obliquity of convergence (Angen et al., 2014; Busby-Spera & Saleeby, 1990; Chardon et al., 1999; Glazner, 1991; Monger & Gibson, 2019; Monger et al., 1982; Paterson & Miller, 1998; Saleeby & Dunne, 2015; Singleton et al., 2024; Teyssier et al., 1995; Tikoff & de Saint Blanquat, 1997; Tobisch et al., 1995). Normal faulting and periods of crustal extension may signal trench retreat, slab roll back, or weak coupling across the convergent plate boundary (Horton, 2018; Maloney et al., 2013; Ramos, 2010). In detail, however, the linkage of deformation to plate motions may be obscured by uncertainties regarding the relation of deformation to magmatism and crustal heterogeneity within the evolving batholith (e.g., Attia et al., 2022; Burton-Johnson et al., 2022; DeCelles et al., 2015; Seymour et al., 2020).

Advances in understanding magma production in batholiths highlight uncertainties in the interplay of plate dynamics, deformation, and magmatism. Recognition that arc magmatism is episodic with a tempo not directly related to plate motions implies modulation of batholith growth by other processes (Beranek et al., 2017; Burton-Johnson et al., 2022; Chapman et al., 2021; DeCelles et al., 2009, 2015; Ducea & Barton, 2007; Ducea et al., 2009; Paterson & Ducea, 2015). Episodic introduction of upper crustal material into the arc through back-arc thrusting or subduction processes has been proposed to control magmatic tempo (e.g., Cao et al., 2015; DeCelles et al., 2009; Ducea & Barton, 2007; Ducea et al., 2009). This process may be recorded within the evolving arc as a pattern of contractional deformation and crustal thickening coupled with isotopically more evolved magmas during high-flux events (HFE) (Ducea & Barton, 2007; Ducea et al., 2009; Girardi et al., 2012; Pearson et al., 2017). Upper plate deformation, however, does not appear to play a role in HFE's characterized by more primitive magmas derived from mantle or primitive lower crust. Such HFE may be controlled by lower plate dynamics and arc migration and so be largely independent from upper plate deformation (Attia et al., 2020; Cecil et al., 2021; Chapman & Ducea, 2019; Decker et al., 2017; Schwartz et al., 2017). In these situations, deformation may be a response to magmatism and therefore may mimic the distribution and tempo of episodic magmatism within the arc (Cao et al., 2015; Paterson & Miller, 1998; Seymour et al., 2020; Tobisch et al., 1995). Alternatively, deformation may not be correlated with magmatic history (Attia et al., 2022; Maloney et al., 2013). Understanding the complex interplay of deformation and magmatism will help resolve the processes governing growth of continental crust in Cordilleran arc systems.

The Coast Mountains batholith (CMB) of western North America provides a rare opportunity to examine these processes along >1,000 km of arc length and through more than 100 m.y. of batholith growth (Figure 1). As one of the largest calc-alkaline magmatic arcs, the CMB records arc construction from Late Jurassic to Eocene time. Formed in a convergent margin marked by changing plate kinematics, the batholith shows along-strike variations in magmatic sources and tempos, crustal exposure level, and metamorphic-structural history. Evaluation of these changes across and along the batholith helps define the roles of plate kinematics, magmatism, and deformation in batholith growth. Our approach is to establish the tectonomagmatic history of a transect across the southern batholith which we compare to a well-known transect 700 km to the north near Prince Rupert (Figure 1) and recently summarized in Woodsworth et al. (2020). Our results reveal a strong spatiotemporal correlation of deformation and magmatism and show that deformation responded to, rather than influenced, the tempo and location of high-flux magmatism in the middle crust. The response of the evolving batholith to changing plate kinematics is highlighted by discovery of sinistral and dextral shear zones and an enhanced record of crustal shortening. These findings support the view of batholiths as accurate recorders of plate dynamics when magmatic tempo is integrated with deformational style but counter models of magmatic cyclicity modulated by crustal deformation.

Figure 1

Map of Coast Mountains batholith, British Columbia and southern Alaska showing Intermontane and Insular composite terranes, major faults, U-Pb ages (Cecil et al., 2018; Gehrels et al., 2009) and study area (Figure 2). MtW, Mount Waddington; NC, North Cascades. Figure adapted from Cecil et al. (2021) and R. B. Miller et al. (2023).

1.1 CMB

The CMB records subduction-related growth from Late Jurassic to early Paleogene time along >1,500 km of continental margin. Plate reconstructions are uncertain but most depict subduction of oceanic plates under North America with varying obliquity of convergence that imposed a shifting transcurrent component on the margin from the Early Cretaceous through early Paleogene (Doubrovine & Tarduno, 2008; Engebretson et al., 1985; Haeussler et al., 2003; Madsen et al., 2006; Stock & Molnar, 1988). West-dipping subduction is proposed in recent reconstructions (Clennett et al., 2020; Hildebrand, 2009; Sigloch & Mihalynuk, 2013, 2017). Further uncertainty surrounds the paleogeography of the CMB during its Cretaceous to Paleogene growth. A component of Early Cretaceous sinistral obliquity has been called upon to translate terranes 600–1,000 km southward (Enkin, 2006; Enkin et al., 2003; Gehrels et al., 2009; Monger et al., 1994; Tochilin et al., 2014). A subsequent dextral component of convergence translated terranes northward after 100 Ma but the amount and manner of this translation is vigorously debated as the “Baja B.C. controversy” (Cowan et al., 1997; Enkin, 2006; Tikoff et al., 2023; Umhoefer, 1987; Wyld et al., 2006). The crux of this debate is that paleomagnetic and some geologic evidence points to ∼1,200–2,700 km of northward translation of the CMB and adjacent terranes yet known fault displacements only sum to about 900 km as summarized in Wyld et al. (2006). Anomalous paleomagnetic inclinations cluster into two groups, one with inferred northward translation of about ∼1,200–1,400 km and ∼2,300–2,700 km for the other (e.g., Tikoff et al., 2023). Sites with the higher inferred displacement occur within and on both sides of the batholith; most of the lower translation sites are east of the batholith (e.g., Tikoff et al., 2023). These sites do not correspond to geologically defined allochthonous “superterranes,” the Intermontane terrane mostly east of the batholith, and the Insular terrane to the west (Colpron & Nelson, 2011; Monger et al., 1982; Wheeler et al., 1991). The location of sites with paleomagnetically inferred anomalous translations shows that CMB is not the boundary between far-traveled terranes and therefore does not constrain the timing of amalgamation of the Insular and Intermontane superterranes terranes. The proposed age of their amalgamation varies from Late Triassic to Late Cretaceous, and the style of amalgamations ranges from collisional to transcurrent (McClelland, Gehrels, & Saleeby, 1992; McClelland, Gehrels, Samson, & Patchett, 1992; Monger et al., 1982; Monger & Gibson, 2019; Rusmore et al., 2013; Tochilin et al., 2014; Umhoefer, 2003; van der Heyden, 1992; White et al., 2016). Resolution of these debates is beyond the scope of this paper but the paleomagnetic and geologic evidence shows that the CMB formed in a dynamic convergent margin with shifting obliquity and potentially large translations through time. The magmatic and deformational history of the batholith reflects this setting.

The magmatic evolution of southern and central segments of the batholith (Figure 1) is well constrained by a robust geochronologic and isotopic data set covering nearly 1,000 km of batholith length (Armstrong, 1988; Cecil et al., 2018, 2021; Cui & Russell, 1995; Friedman & Armstrong, 1990; Gehrels et al., 2009; Mahoney et al., 2009). Pluton age patterns in these segments suggest the CMB includes two distinct Late Jurassic-Early Cretaceous arcs, referred to as the western and eastern arcs (Gehrels et al., 2009; Mahoney et al., 2009). The western arc is distinguished by a pronounced gap in magmatism between 140 and 120 Ma that can be seen in both pluton ages and detrital zircon ages in adjacent strata (Gehrels et al., 2009; Yokelson et al., 2015). Development of a single arc is signaled by a coherent pattern of eastward-younging magmatic ages beginning ca. 120–100 Ma, and magmatism continued until ca. 50 Ma (Cecil et al., 2018; Gehrels et al., 2009). Pluton age patterns and isotopic signatures show that this 120-50 Ma arc formed above an east-dipping subduction zone and, similar to other Cordilleran batholiths, magmatism was episodic (Beranek et al., 2017; Cecil et al., 2018; Chapman et al., 2021; DeCelles et al., 2009; Decker et al., 2017; Ducea et al., 2009; Gehrels et al., 2009; Paterson & Ducea, 2015; Schwartz et al., 2017). Both the central and southern segments of the CMB record a 160–140 Ma HFE, similar to much of the North American Cordillera (Beranek et al., 2017; Paterson & Ducea, 2015). Additionally, both segments share a lull at ca. 75–70 Ma followed by HFE between ca. 60 and 45 Ma. The magmatic tempo, however, differs significantly between ca. 120 and 70 Ma. The central CMB segment is characterized by a protracted HFE from 120 to 78 Ma, while in the southern segment two HFE occur at 114–102 Ma and 85–70 Ma, and are separated by a pronounced lull (Cecil et al., 2018; Gehrels et al., 2009). Crustal thickening and underthrusting of sediments may have contributed to the long Late Cretaceous HFE in the central batholith (Gehrels et al., 2009; Girardi et al., 2012; Pearson et al., 2017). The southern batholith shows a weaker isotopic signal of crustal thickening during HFE magmatism and is more juvenile than the central batholith (Cecil et al., 2021). These characteristics support a mantle-source for magmas and suggest that magmatic tempo is controlled by lower-plate dynamics rather than upper-plate deformation (Cecil et al., 2021).

Deformation during growth of the batholith has been broadly linked to varying obliquity of convergence in Early Cretaceous to early Neogene (e.g., Andronicos et al., 2003; Chardon et al., 1999; Crawford & Crawford, 1991; Hollister & Andronicos, 2006; Journeay & Friedman, 1993; Monger & Gibson, 2019; Rubin et al., 1990; Rusmore & Woodsworth, 1991; Umhoefer & Miller, 1996). Sinistral shear zones on both sides of the batholith are linked to a sinistral component of oblique subduction (Chardon et al., 1999; Gehrels et al., 2009; Monger & Gibson, 2019). In the western batholith, sinistral faults form a northwest-striking system more than 200 km long within the active arc from at least 114 to 101 Ma (Angen et al., 2014; Chardon et al., 1999; Nelson et al., 2011, 2012; Wang et al., 2022). As sinistral faulting waned, widespread crustal shortening became the dominant style of deformation within the Late Cretaceous arc, reflecting a shift toward more orthogonal convergence (e.g., Angen et al., 2014; Crawford & Crawford, 1991; Crawford et al., 1987; Journeay & Friedman, 1993; Monger et al., 1982; Rubin et al., 1990; Rusmore & Woodsworth, 1991, 1994). Southwest-directed thrust faulting began as early as 110 Ma in the western batholith and continued until about 85 Ma (e.g., Cook & Crawford, 1994; Crawford et al., 1987; Journeay & Friedman, 1993; Rubin et al., 1990; Wolf et al., 2010). Younger northeast-directed thrust belts were active along the eastern edge of the batholith between ca. 90 and 68 Ma (Israel et al., 2006; Journeay & Friedman, 1993; Rusmore & Woodsworth, 1991, 1994; Rusmore et al., 2000; van der Heyden, 1992; Woodsworth, 1979; Woodsworth et al., 2000). This contraction was followed by complex but dominantly northeast-side-up slip on the Coast shear zone (CSZ) within the batholith between ca. 65 and 55 Ma (Crawford & Hollister, 1982; Crawford et al., 1987; Gehrels, 2000; Gehrels, McClelland, Samson, Jackson, et al., 1991; Gehrels, McClelland, Samson, Patchett, & Brew, 1991; Ingram & Hutton, 1994; Klepeis et al., 1998; McClelland, Gehrels, & Saleeby, 1992; McClelland, Gehrels, Samson, & Patchett, 1992; Rusmore et al., 2001; Stowell & Hooper, 1990; Wood et al., 1991). A final phase of deformation included local extension on the eastern side of the batholith (Andronicos et al., 2003; Friedman & Armstrong, 1988; Rusmore et al., 2005) and dextral transcurrent faults to the east between ca. 55–35 Ma (R. B. Miller et al., 2023; Monger & Brown, 2016; Umhoefer & Miller, 1996), which collectively translated the batholith northward at least 325–450 km (R. B. Miller et al., 2023; Wyld et al., 2006). The known transcurrent displacements, both sinistral and dextral, are less than inferred from paleomagnetic results throughout the batholith and flanking terranes.

1.2 Study Area: The Bute-Waddington Transect

The geologic framework for our study across the Bute-Waddington transect (Figure 2) was established by mapping projects through the Geological Survey of Canada (Mustard et al., 1994; Roddick & Tipper, 1985; Roddick & Woodsworth, 1977, 2006; Rusmore & Woodsworth, 1993) and complemented by geochronologic surveys (compiled in Rusmore et al., 2013) and topical studies (Bollen et al., 2022; Cecil et al., 2018, 2021; Dafov et al., 2020; Kerrick & Woodsworth, 1989; Nelson, 1979; Rusmore & Woodsworth, 1991, 1994; Rusmore et al., 2013; Woodsworth, 1979). These studies revealed a complex orogen composed of extensive plutons and variably deformed and metamorphosed pendants. The terrane affinity of most pendants is uncertain (Cecil et al., 2021; Dafov et al., 2020; Rusmore et al., 2013; Umhoefer et al., 1994). Pendants belonging to Wrangellia within the Insular terrane occur in the westernmost CMB (Nelson, 1979) but lose their distinctive characteristics within our study area (Figure 2). These pendants are generally small, NW-trending and composed of quartzite, marble, pelite, and metabasalt or fine-grained metavolcanics and metaclastic rocks of greenschist to amphibolite facies (Bollen et al., 2022; Dafov et al., 2020; Roddick & Woodsworth, 2006). Metamorphism of these pendants occurred between ca. 170 – 130 Ma and again between ca. 110–87 Ma and is attributed to heat from magmatism during HFEs (Dafov et al., 2020). Analysis of detrital zircon from the quartzite-bearing metasedimentary assemblages shows the protoliths likely accumulated in the Late Paleozoic and Triassic to Early Jurassic and do not correlate with the Paleozoic margin assemblage typical of the Yukon Tanana terrane. Correlation to the Intermontane superterrane of the plutons which envelop the pendants is supported by plutonic age patterns and paleomagnetic results similar to those typical of the Intermontane terrane (Cecil et al., 2018, 2021; Rusmore et al., 2013). The central and eastern part of the batholith contains extensive pendants of metasedimentary and metaigneous rocks of unknown terrane affinity (Bollen et al., 2022; Roddick & Tipper, 1985; Rusmore & Woodsworth, 1994; Woodsworth, 1979). Triassic through Late Cretaceous strata in the eastern batholith may belong to Stikinia but lack definitive stratigraphic links (Rusmore & Woodsworth, 1994; Umhoefer et al., 1994).

Figure 2

Simplified geologic map of the Bute-Waddington transect area. Samples used in this study are shown; those with new data are labeled. New or modified faults shown in black. A and B = endpoints of cross section in Figure 13. Asz = Alpha Bluff shear zone, Bsz = Blind Creek shear zone, CSZ = Coast shear zone, Cts = Cumsack thrust system, EWTB = eastern Waddington thrust belt, Wsz = Wahkash shear zone. Map and data are from this study and modified from Woodsworth (1979), Roddick and Tipper (1985), Roddick and Woodsworth (2006), Rusmore and Woodsworth (1993), Rusmore et al. (2013), Mustard et al. (1994), Cui et al. (2017), Cecil et al. (2018), and Dafov et al. (2020).

Episodic magmatism constructed the batholith from Late Jurassic through early Neogene time with periods of high apparent magmatic flux at 160–140 Ma, 114–102 Ma, 85–70 Ma, and 61–46 Ma (Cecil et al., 2018). Plutons become younger to the northeast and after ca. 100 Ma, magmatism migrated northeast at about 2–3 km/yr (Cecil et al., 2018). Biotite and hornblende K/Ar and ⁴⁰Ar/³⁹Ar ages generally mimic this pattern, varying from Late Jurassic in the southwest to early Tertiary in the northeast (Rusmore et al., 2013). The eastern batholith was affected by widespread Late Cretaceous northeast-directed thrust faulting within the eastern Waddington thrust belt (Rusmore & Woodsworth, 1991, 1994). Regional metamorphism accompanied and outlasted deformation, with metamorphic grade decreasing to the northeast and ending at 65 Ma (Bollen et al., 2022; Rusmore & Woodsworth, 1994; Woodsworth, 1979; Woodsworth et al., 2000).

2 Methods

The primary goal of our research was to integrate the deformation and metamorphism of the southern batholith with the recently established magmatic framework (Cecil et al., 2018, 2021). Our work covers a subset of the area studied by Cecil et al. (2018, 2021) and was undertaken jointly by the authors of those studies, with multiple senior authors involved in all field work and subsequent coordination of analyses. Our approach was to concentrate on those parts of the batholith that our reconnaissance and regional mapping (Roddick & Tipper, 1985; Roddick & Woodsworth, 2006) suggested were deformed and metamorphosed. The selected areas were studied in detail to establish the structural-metamorphic evolution and its relation to intrusive rocks. Metamorphic results and garnet geochronology are reported in Bollen et al. (2022); here these findings are linked to the tectonomagmatic evolution through integrated field, geochronologic, petrographic, and structural analysis. Particular care was taken to sample intrusive bodies ranging from plutons to small dikes and sills with clear field relations to structures. Large plutons were mapped both internally and along and across their contacts. We integrate our results with geochronologic, paleomagnetic, and geologic studies across this part of the CMB batholith to define the timing and conditions of deformation and its relation to magmatism in the evolving batholith. This record is then considered in the context of the evolution of Cordilleran batholiths.

The area studied crosses the batholith from southwest to northeast for 120 km and is up to 150 km wide (Figures 1 and 2). The region is remote, roadless, and encompasses both heavily forested coastal regions and high alpine mountainous terrain. Exposure is excellent along fjord shorelines and on ridges between glaciers and icefields. Elsewhere, rocks are poorly exposed. Two fjords, Bute and Knight Inlets, provided access by small boat in the western half of the transect; field work was conducted from helicopter-supported basecamps and regional sampling in the mountainous core of the batholith. Field data were collected in ESRI ArcPad on tablets with integrated GPS with 2–3 m accuracy. Digital map compilations (Cui et al., 2017; Roddick & Woodsworth, 2006) were modified based on our mapping. Interpolation of the geology between mapped areas was based on review of Geological Survey of Canada field notes and samples collected in 1967–1972, and analysis of satellite imagery from Google Earth and ERSI. Additional data on dikes and pluton fabrics were retrieved from GSC digital field data obtained from 1970 to 1976. Data were compiled and analyzed using ESRI ArcMap and Stereonet v11.3.0 (Allmendinger et al., 2013). Across the study area, samples are located by distance from the CSZ, allowing direct comparison to data from the central CMB (Gehrels et al., 2009).

2.1 U-Pb Geochronology

Fourteen new igneous zircon and one titanite U-Pb ages were generated from orthogneiss and variably deformed sills, dikes and plutons to constrain the age of deformation and relation to magmatism. Samples were crushed and zircon were extracted using conventional density and magnetic separation techniques. Euhedral and inclusion-free grains were then selected and mounted with reference materials in a 2.5 cm epoxy puck and polished to a 1-μm finish. Zircon mounts were then imaged using backscattered electron (BSE) and cathodoluminescence (CL) detectors coupled to a Hitachi S-3400N scanning electron microscope at the University of Arizona LaserChron Center (ALC). Prior to analysis, common Pb contamination of the mount surface was removed via ultrasonication in a dilute HNO3 + HCl bath. BSE and CL images were used to guide analyses, detect possible inherited or metamorphic domains, and screen for grain defects.

Zircons were analyzed using laser ablation inductively-coupled plasma mass spectrometry (LA-ICPMS) at the ALC using methods described in Gehrels et al. (2008) and Gehrels and Pecha (2014). Grains were ablated at a 7 Hz hit rate and with a 30-μm spot size using a Photon Machines Analyte G2 ArF excimer laser system coupled to a Thermo Element2 single-collector magnetic-sector ICP-MS. Resulting ablation pits were approximately 20 μm deep. Between 25 and 35 zircons were analyzed for all samples except 14MR33 and 14MR34, which had low zircon yields. ALC in-house Sri Lanka zircon (SL; 558 Ma) and the Duluth Complex Anorthosite standard FC-1 (1099 Ma; Paces & Miller, 1993) were used as primary reference materials and analyzed after analysis of every 5 unknown grains. Reported ages are given as the weighted mean of all statistically overlapping concordant analyses. Age uncertainties include analytical uncertainty, based on the scatter and precision of individual analyses, and the systematic uncertainty, which includes errors associated with the age of zircon standards, the 238U decay constant, and common Pb composition (used for common Pb correction). In most cases, the combined uncertainty is 1%–2% (2σ).

2.2 ⁴⁰Ar/³⁹Ar Geochronology

Hornblende and biotite from nine intrusive samples were analyzed using ⁴⁰Ar/³⁹Ar methods at Stanford University (McDougall & Harrison, 1999). Samples were irradiated at the Oregon State TRIGA reactor. Irradiation parameters were determined from measurements performed with co-irradiated FCT sanidine, kalsilite glass, and CaF₂. The values of these correction factors are provided below the data tables of individual samples in Data Set S2. Samples encapsulated in a 3 mm diameter × 0.5 mm thick tantalum foil packet. Automated argon release was accomplished using a fiber optic laser microfurnace (908 nm wavelength) controlled by a dedicated microprocessor and Labview software that controlled the all-metal extraction line, getter pumps, and Noblesse mass multicollector spectrometer. Temperature was monitored with a 1.6 nm optical pyrometer that had been calibrated between 250 and 1250°C by the diffusion properties of sanidine glass (Lovera et al., 1997) and the melting point of Aluminum (Oze et al., 2017). Argon isotopic ratios measured on Faraday and ion counter detectors were calibrated against measurements of a ⁴⁰Ar-³⁹Ar-³⁸Ar-³⁶Ar reference gas (Coble et al., 2011).

Because the samples examined experienced natural cooling and/or transient heating related to pluton emplacement within the crust, integrated total gas ages (i.e., K-Ar ages) best represent the age gradients present within them (McDougall & Harrison, 1999). Total gas ⁴⁰Ar/³⁹Ar ages are thus used in all interpretations in this paper. The ⁴⁰Ar/³⁹Ar ages obtained for samples in this study were added to a compilation of K-Ar and ⁴⁰Ar/³⁹Ar ages from Rusmore et al. (2013). This compilation has been adapted to match the present study area, and published ages with >10% error have been omitted.

2.3 Hornblende Thermobarometry

Five samples were selected for hornblende thermobarometry by screening 71 thin sections of intermediate composition plutons. Most samples were discarded either because they lacked the full buffering assemblage Hbl + Pl + Bt + Kfs + Qz + Ttn + Fe-Ti oxide or because they showed partial alteration to low greenschist facies minerals. Samples chosen lack significant alteration and have sharp grain boundaries between hornblende and plagioclase in thin section and SEM imagery. In each sample, 8–9 hornblende and plagioclase crystals were analyzed where they are in contact along euhedral grain boundaries with 1–4 pairs of analyses collected along each boundary. BSE images with spot locations are in Supporting Information S1 and https://doi.org/10.5061/dryad.jsxksn0kt. Hornblende structural formulae were calculated as per Holland and Blundy (1994). Pressures were calculated using hornblende rim compositions in contact with plagioclase, and the Schmidt (1992) experimental calibration as modified for temperature by Anderson and Smith (1995).

3 Results

3.1 U-Pb Geochronology

Weighted mean U-Pb zircon ages range from 161.5 to 73.5 with one outlier at 363 Ma (Table 1). The latter is presumably Wrangellia basement, which is dominated by igneous rocks of Late Devonian to Mississippian age (Ruks, 2015). Most samples yielded simple zircon systematics, with concordant, reproducible ages reflecting magmatic zircon growth (Figures 3a and 3b).

Figure 3

Representative plots of zircon U-Pb data. Concordia (a) and weighted mean (b) plots of sample 14MR27, a dioritic mylonite. Like most samples analyzed as part of this study, 14MR27 yielded reproducible, tightly-clustered ages and grains with U/Th < 10. Concordia (c) and weighted mean (d) plots of sample 14MR46, a granodioritic folded sill. 14MR46 is the most complex sample analyzed as part of this study. It yielded four Jurassic inherited grains and several Cretaceous zircon with high U/Th (see inset plot in (c)). Most of the high U/Th grains were excluded on the basis of being statistical outliers (pink ellipses in (a) and (c) and pink bars in (b) and (d)).

Table 1. U-Pb Geochronology

Sample #	Weighted mean age (Ma) (2σ) zircon (Zr) titanite (TT)	Rock type	Transect position (km NE of CSZ)	Latitude WGS84	Longitude WGS 84	Domain	Geologic context
14MR13	118.6 ± 1.8 (Zr)	Granitic mylonite	−46	50.60	−124.92	Coastal	Mylonitic pluton, Alpha Bluff SZ
14MR27	155.2 ± 1.7 (Zr)	Dioritic mylonite	−20	50.91	−124.83	Coastal	Mylonitic pluton; diorite complex; Cumsack thrust system
14MR33	114 ± 12 (Zr)	Dioritic mylonite	−46	50.60	−124.92	Coastal	Thin (<10 cm) mylonitic sill in 14MR34, Alpha Bluff SZ
14MR34	362.8 ± 5.1 (Zr)	Quartz dioritic mylonite	−46	50.60	−124.92	Coastal	Lens in Alpha Bluff SZ
14MR40	73.9 ± 0.8 (Zr)	Tonalitic sill	10	51.23	−125.06	Interior	Late kinematic sill, (<100 m thick); Waddington TB
14MR46	73.5 ± 1 (Zr)	Granodioritic sill	6	51.2	−125.1	Interior	Folded sill (<15 cm); late kinematic Waddington TB
14MR50	144.1 ± 2.9 (Zr)	Granitoid gneiss	6	51.25	−125.06	Interior	Protolith age; Waddington core
14MR59	74.4 ± 1 (Zr)	Granodioritic sill	4	51.25	−125.14	Interior	Folded sill (<15 cm); late kinematic Waddington thrust belt
15MR01	147.4 ± 2.1 (Zr)	Tonalitic mylonite	−66	50.69	−125.70	Coastal	Mylonitic pluton; Blind Cr sz
15MR02	105.2 ± 1.2 (Zr)	Meta-andesite	−40	50.91	−125.57	Coastal	Dike; cuts 15MR03
							Wahkash sz
15MR03	104.5 ± 1.8 (Zr)	Tonalitic orthogneiss	−40	50.91	−125.57	Coastal	Wahkash sz
15MR05	119.2 ± 1.5 (Zr)	Granitic ultramylonite	−34	51.97	−125.59	Coastal	Wahkash sz
15MR09	128.6 ± 2.1 (Zr)	Tonalite pluton	−44	50.89	−125.63	Coastal	Protomylonite between ASZ and WSZ
15MR21	161.5 ± 2.0 (Zr)	Granodioritic mylonite	−24	50.95	−124.99	Interior	Mylonitic pluton, upper plate of Cumsack thrust system
	78.6 ± 3.2 Ma (TT)

Three samples are more complex, with inherited zircon and/or high U/Th values. Samples 14MR13 and 14MR46 (Table 1) yielded Cretaceous ages but included inherited Jurassic zircon (Figures 3c and 3d). Sample 14MR13, a mylonitic granite, has an interpreted crystallization age of 118.6 ± 1.8 (from 21 concordant analyses; MSWD = 2.8), but also yielded 5 grains with ages between 141.4 and 148.2 Ma. Sample 14MR46, from a folded granodiorite sill, has an interpreted crystallization age of 73.2 ± 1.0 Ma (from 23 concordant analyses; MSWD = 0.95), but also yielded four zircons with ages between 144.6 and 158.3 Ma. It also yielded 12 grains with high U/Th, ranging from 12 to 450. The third sample (14MR59) from another folded sill near 14MR46 also yielded a Late Cretaceous age (74.4 ± 1.0 Ma) and abundant zircon with high U/Th ranging from 13 to 45. Weighted mean and concordia plots for all samples are in the Supporting Information S1 and all individual zircon and titanite U-Pb data are in Data Set S1 and https://doi.org/10.5061/dryad.jsxksn0kt.

3.2 ⁴⁰Ar/³⁹Ar Geochronology

Nine intrusive samples produced nine biotite ages and seven hornblende ages presented in Table 2 and Supporting Information S1, Data Set S2, and https://doi.org/10.5061/dryad.jsxksn0kt.

Table 2. ⁴⁰Ar/³⁹Ar Total Gas Ages

Sample	Biotite (Ma)	Hornblende (Ma)	U-Pb zircon (Ma)^a	Latitude WGS84	Longitude WGS 84	Transect position (km NE of CSZ)	Domain	Geologic context
00MR-15	73.8 ± 1.9	—	148.3 ± 5.3	51.22	−125.54	−8	Coastal	Orthogneiss
00MR-34	49.7 ± 1.3	—	51.0 ± 1.5	51.41	−125.27	18	Interior	Tiedemann pluton
00MR-37	49.1 ± 1.3	51.7 ± 1.4	75.0 ± 1.4	51.47	−125.18	27	Interior	Orthogneiss
14MR61	54.2 ± 1.4	65.1 ± 1.8	72.0 ± 0.9	51.25	−125.15	4	Interior	Postkinematic pluton
14MR68	79.7 ± 2.1	84.5 ± 2.2	84.8 ± 1.1	51.06	−125.13	−16	Coastal	Postkinematic pluton
14MR69	68.8 ± 1.8	89.7 ± 2.4	96.1 ± 1.1	51.14	−124.97	−3	Coastal	Postkinematic pluton
14MR70	79.0 ± 2.1	78.7 ± 2.1	77.9 ± 1	51.00	−124.88	−14	Coastal	Postkinematic pluton
15BCKH13	82.1 ± 2.2	94 ± 2.5	98.4 ± 1.2	50.92	−125.52	−38	Coastal	Cross-cutting dike
2727R08	52.9 ± 1.4	56.2 ± 1.7	60.7 ± 1.2	51.77	−125.21	57	Interior	Postkinematic pluton

• ^a Zircon ages from this paper (Table 1) and Cecil et al. (2018).

3.3 Hornblende Geobarometry

Hornblende pressures from five plutons are presented in Table 3, Data Set S3, and https://doi.org/10.5061/dryad.jsxksn0kt.

Table 3. Hornblende Geothermobarometry

Sample #	Pressure (kb)	Temperature (°C)	Pluton age (Ma)	Rock type	Latitude WGS84	Longitude WGS 84	DistanceNE of CSZ (km)	Domain	Geologic context
14MR61	6.8 ± 0.3	692 ± 12	72.0 ± 0.9	Tonalite	51.25	−125.15	4	Interior	Post-kinematic pluton
14MR68	4.9 ± 0.2	680 ± 11	84.8 ± 1.1	Tonalite	51.06	−125.13	−16	Coastal	Post-kinematic pluton
14MR72	5.6 ± 1.0	701 ± 6	71.9 ± 1.0	Tonalite	51.11	−124.54	11	Interior	Post-kinematic pluton
2727R08	3.9 ± 0.5	645 ± 12	60.4 ± 1.3	Granodiorite	51.77	−125.21	57	Interior	Post-kinematic pluton
89MR65	5.7 ± 0.7	674 ± 23	61.0 ± 0.7^a	Tonalite	51.284	−124.63	21	Interior	Mantle Glacier pluton

• Note. Pluton ages from Cecil et al. (2018).
• ^a Sampled within dated pluton 14MR74.

3.4 Geologic Results

Geologic results are divided into two regional domains, referred to as the Coastal and Interior domains and shown on Figure 2. The domains are distinguished by their geologic histories and are separated by a newly recognized shear zone which we interpret as the southern continuation of the CSZ. The geology of each domain is described prior to discussion of the CSZ. Structural orientation data presented here are in Data Set S4 and https://doi.org/10.5061/dryad.jsxksn0kt.

3.4.1 Coastal Domain

3.4.1.1 Intrusive Rocks in the Coastal Domain

Plutons in the Coastal domain range from gabbroic to granitic in composition, but most are quartz diorite, diorite, and tonalite (Cecil et al., 2018; Roddick & Woodsworth, 2006). Most plutons are moderately to highly elongate northwest to southeast in map view and have near vertical contacts (Cecil et al., 2018; Roddick & Woodsworth, 1977, 2006). Internally, most plutons are unfoliated or have a locally developed magmatic foliation defined by aligned euhedral igneous minerals with well-preserved intrusive textures and compositional bands a few cm wide. About 70% of the pluton outcrops examined have no foliation, and only a few have a prominent magmatic foliation. Although generally weak, the magmatic foliation consistently strikes northwest and dips steeply northeast (Figure 4), similar to the elongate shape of many plutons. Mafic enclaves are common in the plutons, and swarms of closely spaced enclaves compose most of some outcrops. The geometry of these swarms varies from low angle sheets to nearly vertical pipes.

Figure 4

Magmatic textures of plutons in the Coastal domain. Stereonets are equal area. (a) Poles to magmatic foliation. Gray scale shows Kamb contours of standard deviation ranging from 2 (lightest gray) to 18 (black). (b) Poles to dikes. (c) Outcrop photograph of diorite complex with inset of 81 Ma sample 14MR55 (U-Pb zircon age in Cecil et al. (2018)).

A distinctive and widespread unit, the “diorite complex” of Roddick and Woodsworth (1977, 2006) plays a key role in constraining the geologic and tectonic evolution of the southern CMB. The diorite complex, which constitutes about 20% of the area mapped as plutons in the Coastal domain is characterized by compositional and textural heterogeneity within outcrops and individual plutons (Figures 2 and 4). Diorite to quartz diorite is the most common compositions with lesser tonalite and rare gabbro occurring throughout the complex. Individual plutons are a complex mix of diorite containing ill-defined fine-grained mafic to ultramafic (hornblendite) lenses, mafic dikes, and distinctive tonalite pods, sills, and small stocks The boundaries between rock types are irregular and lack layering or a consistent orientation and the plutons are generally unfoliated. Grain size varies greatly within outcrops, ranging from medium to coarse with euhedral hornblende reaching several cm long. Where dated, adjacent tonalite and diorite are coeval (Cecil et al., 2018). U-Pb zircon ages range from 160 to 77 Ma and follow the general trend of decreasing age with distance to the northeast (Cecil et al., 2018). The diorite complex is commonly metamorphosed to lower greenschist facies and cut by networks of irregular fractures filled with epidote, plagioclase, and chlorite.

Within the Coastal domain, many plutons are cut by greenish gray andesitic dikes (Rusmore et al., 2013). Most dikes are planar, <1 m wide, aphanitic to plagioclase- or hornblende-phyric and have a pervasive chloritic alteration. The location and orientation of 48 of these dikes were recorded in the GSC digital records and our prior unpublished mapping. Most of these occur in the western half of the Coastal domain; very few are present in the northeastern reaches of Bute and Knight Inlets. None were mapped in plutons younger than 104 Ma and the only dated dike is 105 Ma (15MR02, Table 1), suggesting the dikes are older than ca. 104 Ma. The dikes show no preferred orientation and most dip steeply; 70% have dips greater than 70° (Figure 4) and half of those with shallower dips occur within ductile shear zones. The planar shape of the dikes and their generally steep dip is consistent with undeformed character of the plutons they intrude, and the pattern of dike orientations suggests that widespread systematic tilting of the plutons has not occurred since the dikes were intruded ca. 104 Ma, consistent with the conclusion of Rusmore et al. (2013).

3.4.1.2 Pendants in the Coastal Domain

Along most of the western side of the CMB, metamorphic rock pendants are low grade in the southwest and higher grade to the northeast. Pendants along the western reaches of Bute and Knight Inlet follow this pattern, with sub-greenschist and low greenschist facies pendants in the southwest, and amphibolite facies pendants to the northeast (Bollen et al., 2022; Dafov et al., 2020). This pattern, however, does not persist across the Coastal domain; instead, metamorphic grade decreases northeast of the amphibolite facies rocks and low-grade volcanogenic strata form an extensive pendant, referred to here as the Cumsack pendant (Figure 2). The pendant is composed of thick (up to 40 m) dacitic and rhyolitic flows and fragmental deposits interbedded with thinner beds of felsic tuff, tuffaceous sandstone, black shale, red siltstone, and minor pebbly mudstone and conglomerate (Figure 5). Clasts are mostly felsic to intermediate volcanics, with rare rounded granitoid clasts. Shale beds locally contain very poorly preserved shelly marine fossil fragments. The age of strata in the pendant is unknown; attempts to retrieve detrital zircon from tuffaceous sandstone were unsuccessful. The pendant lacks the characteristic quartzite + carbonate assemblage of pendants exposed to the southwest and along strike to the northwest (Boghossian & Gehrels, 2000; Dafov et al., 2020). Thus, the Cumsack pendant is anomalous in terms of both lithology and metamorphic grade.

Figure 5

Field photographs of volcanogenic rocks in the Cumsack pendant. (a) Thin bedded tuff capped by intermediate volcanic flow. (b) Weakly deformed volcaniclastic breccia, large clast is a volcanic bomb with chilled margins, (c) Volcaniclastic sandstone.

3.4.1.3 Sinistral Fault System

Three newly recognized sinistral shear zones (Blind Creek, Alpha Bluff, and Wahkash) form an Early Cretaceous sinistral fault system (Figures 2 and 6). The shear zones strike northwest, are 15–20 km apart and more than 50 km long. Mylonitic fabrics are variably developed and preserved along the length of the shear zones, and kinematic information is observed on faces parallel to lineation and perpendicular to foliation (Figure 7); other orientations did not yield consistent kinematics. The relation of shear bands to mylonitic foliation is the most reliable kinematic indicator because crystal-scale features are commonly obscured by annealing.

Figure 6

Coastal domain geologic map and equal area stereonets of structures. Intrusive ages discussed in text are labeled; pluton ages in italics are from Cecil et al. (2018). Unit key shown on Figure 2. Blind Cr shear zone: n = 15 S, 7 L; Alpha Bluff shear zone n = 19 S, 15 L; Wahkash shear zone n = 33 S, 18 L.

Figure 7

Photographs of sinistral shear zones, Coastal domain. Examples of kinematic indicators annotated on surfaces parallel to lineation and perpendicular to foliation. (a) Mylonitic tonalite showing weak elongation lineation on vertical mylonitic foliation surface, Blind Creek sz; (b) Mylonitic tonalite, Blind Creek sz. Panels (a) and (b) are sample 15MR01, 147.4 ± 2.1 Ma, Table 1; (c) Microphotograph of mylonitic tonalite, Blind Creek sz; (d) Outcrop of granitic mylonite (14MR13, 118.6 ± 1.8 Ma, Table 1), Alpha Bluff sz; (e) Cut slab of tonalitic mylonite; hazy appearance caused by annealing, Alpha Bluff sz; (f) Outcrop of tonalitic mylonite, Wahkash sz; (g) Outcrop lineated granitic ultramylonite (15MR05, 119.2 ± 1.5 Ma, Table 1), Wahkash sz; (h) Cut slab of granodioritic mylonite, 126 Ma pluton (Cecil et al., 2018), Wahkash sz.

The southwestern shear zone, the Blind Creek shear zone is best exposed on Knight Inlet, where mylonite and protomylonite derived from quartz diorite to tonalite are intermittently exposed for ∼500 m across strike and form a discrete ∼50 m thick mylonitic core. The shear zone is inferred to continue southeast along strike for >50 km to Bute Inlet based on sparse exposures of mylonitic granite on Loughborough Inlet and a few hundred meters of variably developed protomylonitic and mylonitic fabrics in quartz diorite on Bute Inlet. Mylonitic fabrics are consistent along the length of the shear zone; foliations strike NW, and dip steeply; elongation lineations plunge moderately NW and SE (Figure 6). Mylonitic fabrics show a sinistral to sinistral-reverse sense of shear in outcrops and thin sections (Figure 7).

Mylonitic plutonic samples from the Blind Creek shear zone are 147.4 ± 2.1 Ma (15MR01, Table 1, Figures 7a and 7b) on Knight Inlet and 136.7 ± 1.6 Ma on Bute Inlet (Cecil et al., 2018). An undeformed elongate NW-trending pluton, dated as 104 Ma (Cecil et al., 2018), obscures much of the projected trace of the fault and is inferred to have intruded the Blind Creek shear zone after displacement ceased. Collectively, these ages show the Blind Creek shear zone shear zone was active after 137 Ma and before 104 Ma.

The Alpha Bluff shear zone is best exposed on Bute Inlet where it is marked by ∼1 km of mylonitic intrusive rocks (Figures 7d and 7e). Mylonites and protomylonites are prevalent; ultramylonites are generally a few meters thick and occur near the southwestern edge of the shear zone. Review of GSC field data suggests the shear zone continues along strike through remote country between Bute and Knight Inlets where low-grade strata are faulted against orthogneiss (Figure 6). On Knight Inlet, the Alpha Bluff shear zone is expressed as poorly exposed mylonitic zones within the diorite complex. Mylonitic foliation in the shear zone consistently strikes northwest, is nearly vertical, and contains a well-developed elongation lineation that plunges variably to the northwest and southeast (Figure 6). Outcrop and thin section observations of S-C, S-C’, and σ -type porphyroclasts show a consistent sinistral to sinistral-reverse sense of shear (Figure 7e). Mylonites within the shear zone on Bute Inlet (Figures 2 and 6) have zircon U-Pb ages of 362.8 ± 5.1 Ma from mylonitic quartz diorite (14MR34, Table 1), 114 ± 12 Ma from a mylonitic diorite sill, (14MR33, Table 1), and 118.6 ± 1.8 Ma from a granitic mylonite (14MR13, Table 1, Figure 7d). A distinctive granitic and locally mylonitic 117 Ma pluton (Cecil et al., 2018) occurs on both sides of the shear zone on the eastern shore of Bute Inlet (Figure 6). These ages show the shear zone was active after 119 Ma and limit sinistral displacement after 117 Ma to less than the ~20 km length of the granitic pluton (Roddick & Woodsworth, 2006).

The Wahkash shear zone is largest of the three sinistral shear zones and was partially mapped by Roddick and Woodsworth (2006). It is several kilometers wide, more than 70 km long, and marks a discontinuity in the composition and metamorphic grade of pendants (Figure 6). The widespread and distinctive quartzite-bearing pendants exposed to the west (Dafov et al., 2020) end at the shear zone and metamorphic grade drops abruptly across the shear zone. Amphibolite facies schists southwest of the shear zone are juxtaposed against low-grade volcanogenic rocks of the Cumsack pendant and plutons to the northeast (Bollen et al., 2022; Roddick & Woodsworth, 2006). Truncation of the schist is well exposed on both Bute and Knight Inlets and is visible on satellite imagery between the inlets where it defines a vertical shear zone that strikes 290°–300°. For simplicity, the shear zone is placed at the truncation of the schist, but strain related to the shear zone persists into adjacent rocks for several kilometers on both sides of this discrete contact. To the northeast is 0.5–1.5 km of granitic to tonalitic mylonitic plutons derived in part from a 126 Ma two-mica granite and a 145 Ma leucogranite/fragmental rhyolite stock (Cecil et al., 2018) of the Cumsack pendant. The mylonitic foliation is vertical and strikes NW (average: 312°, 90) a well-developed elongation lineation is gently to moderately plunging on Knight Inlet and vertical on Bute (Figure 6). Mylonite and ultramylonite are more abundant near the contact with the schist and protomylonite is increasingly common to the northeast, suggesting an overall decrease in strain away from the shear zone. Outcrops of mylonites have S-C, S-C′, σ -type porphyroclasts and rare white mica “fish” (Figure 7f–7g). In thin section, however, the mylonitic fabric is thoroughly annealed. Relic compositional layers are composed of quartz, feldspar + quartz, and aligned micas + granular epidote ± chlorite. Quartz is fine to medium grained, equant and polygonal; the cores of plagioclase porphyroclasts remain but tails are composed of mosaics of quartz and feldspar. Clear kinematic information is visible in faces parallel to the lineation and perpendicular to foliation in 9 of 18 locations on Knight Inlet and is consistently sinistral to sinistral reverse (Figure 7). No consistent sense of shear could be discerned in samples on Bute Inlet in any sample orientation. On a larger scale, a sinistral sense of shear is consistent with the more northerly average strike of the mylonitic foliation (312°) compared to the shear zone boundary (290°–300°). Displacement is younger than a U-Pb zircon age of 119.2 ± 1.5 Ma on a granitic ultramylonite within the shear zone (15MR05, Table 1, Figure 7g).

Southwest of the Wahkash shear zone, the amphibolite-facies schist and orthogneiss record deformation prior to and synchronous with development of the shear zone. Homogeneous quartz diorite to tonalite orthogneiss forms well-foliated bodies several km thick which are undated and locally occurs as mylonitic sills within schist. The schist includes mafic schist, pelite, quartzite and carbonate. Schistosity is defined by amphibole ± biotite, plagioclase, and quartz in mafic schist, and by thin bands of aligned biotite ± sillimanite and quartz + plagioclase in pelites. Schistosity is folded by tight to isoclinal folds with wavelengths of cm to a few meters that likely formed during amphibolite facies metamorphism (Figure 8). White mica lies within and across the schistosity and appears post-kinematic. Synkinematic metamorphic conditions were 4.9–7.4 kbar and 625–685°C in garnet-bearing pelitic schist (Bollen et al., 2022). A 110 Ma diorite pluton (Cecil et al., 2018) crosscuts the schistosity and includes rafts of schist and orthogneiss near its margins (Figure 6). This early phase of metamorphism and deformation is thus older than 110 Ma, compatible with the ca. 170–130 Ma phase of metamorphism suggested by Dafov et al. (2020).

Figure 8

Structures in metamorphic rocks, southwest side of Wahkash shear zone. (a) Outcrop photograph of F1 isocline folded by upright F2 with boudinaged felsic vein in schist; (b) Outcrop photograph of schist with upright F2 folds next to Wahkash shear zone; (c) Equal area stereonets of foliations, lineations, and folds in metamorphic rocks.

This early phase of deformation and metamorphism is overprinted by younger structures within a few km of the Wahkash shear zone. Their proximity and orientation suggest they are linked to shear zone development. Upright non-cylindrical folds (referred to as F2) with wavelengths of cm to a few meters are present in the schist and interlayered orthogneiss sills, where they form type 2 to type 3 fold-interference patterns (Ramsey, 1967) (Figure 8). F2 axial planes dip steeply and strike NNW (ave. plane: 330°, 88°) and hinge lines trend NW-SE with plunge varying from horizontal to vertical. The axial planes strike ca. 30° north of the Wahkash shear zone, compatible with the folds forming as en echelon structures during sinistral shear. Gneissic to mylonitic foliations parallel to the axial planes are common and locally contain NW-SE trending subhorizontal mineral and elongation lineations (Figure 8). In orthogneiss sills, the foliation and lineation are commonly mylonitic and thoroughly recrystallized.

The parallelism of most lineations and foliations obscures their relative ages; they are best distinguished where refolded or directly dated. A km-scale pendant of schists and interlayered orthogneiss is enveloped in a 103 Ma pluton (Dafov et al., 2020), limiting the age of the structures. Further constraints are provided by two new U-Pb ages on an outcrop within this pendant where a dark meta-andesite dike crosscuts mylonitic granodioritic orthogneiss. Large (cm) K-feldspar σ -porphyroclasts in the orthogneiss show a sinistral sense of shear on faces parallel to lineation and perpendicular to foliation in outcrop but post-kinematic metamorphism has obscured the strain fabric in thin section. The crosscutting dike is planar, >6 m long, 15–20 cm thick with well-defined and planar margins. Within the dike, biotite and amphibole phenocrysts are faintly aligned parallel to the foliation in the orthogneiss, and similar to the orthogneiss, microscopic fabrics have been overprinted by post-kinematic metamorphism. These textures suggest the dike intruded very late in the deformation of the orthogneiss with subsequent static metamorphism of both rock types. The orthogneiss yielded a U-Pb zircon age of 104.5 ± 1.8 Ma (15MR03, Table 1). A U-Pb age on zircon from the dike is 105.2 ± 1.2 Ma, which is indistinguishable from the orthogneiss (15MR02, Table 1). These ages show that F2 folding and development of mylonitic fabrics related to the Wahkash shear zone persisted at least locally until 105 Ma and, more regionally, ceased prior to 103 Ma.

Post-kinematic metamorphism is widespread between the Wahkash and Alpha Bluff shear zones and closely followed the cessation of sinistral shear. Dated plutons older than ca. 100 Ma have a static metamorphic texture. The texture is best developed on Knight Inlet where intrusive textures are partially to wholly recrystallized to a medium-grained mosaic of subequant grains. Intrusive textures are only partially recrystallized on Bute Inlet. Unmetamorphosed intrusive textures occur in 99 and 96 Ma plutons, a 98 Ma dike (Cecil et al., 2018), and a ca. 101 Ma leucocratic sill (Dafov et al., 2020). Collectively, these ages limit post-kinematic metamorphism to ca. 105–100 Ma. Over the larger region, chemistry of zircon from pendants southwest of the Wahkash shear zone suggests metamorphism occurred between ca. 110 and 87 Ma, which is attributed to heat from the batholith (Dafov et al., 2020).

The similar orientation, timing, and kinematics suggest the Blind Creek, Alpha Bluff, and Wahkash shear zones are a linked system of sinistral shear zones. A transpressional component is suggested by the variable plunge of the mylonitic lineations. Most displacement took place on the Wahkash shear zone, the largest and only shear zone that disrupts regional geologic trends. The absence of correlative pendants across the shear zone indicates sinistral displacement exceeds its 70 km mapped length. Including minor displacement on the other shear zones, the total likely exceeds 80–90 km. This fault system was active after ca. 115 Ma, ceased by 103 Ma, and was closely followed by static metamorphism before 100 Ma.

3.4.1.4 Cumsack Thrust System

Steeply dipping thrust faults and associated folds cut plutons and strata of the Cumsack pendant near the head of Bute Inlet (Figures 2 and 6). The thrusts are best exposed in the mountains west of the inlet where two NW-striking thrusts were mapped for ∼3 km, and a lower thrust is inferred from Roddick and Woodsworth (2006). The structurally highest of these thrust places a granodiorite pluton over volcanogenic strata of the Cumsack pendant; the lower thrusts lie within and appear to repeat the strata (Figure 9). The faults are marked by a core of several meters of strongly foliated and lineated clastic rocks and mylonites flanked by more weakly deformed rocks that are a few to >100 m thick. Approaching the thrust faults, quartz veins become more common and are increasingly disaggregated. Similarly, granodioritic dikes occur within ∼50 m of the upper thrust and become boudinaged and discontinuous toward the thrust fault. Within the core of the thrusts, coarse-grained sandstones and breccias develop a clast-defined foliation and lineation; finer grained clastic rocks show spaced to slaty cleavage with small chlorite and biotite grains on the cleavage planes. Flinty mylonites derived from felsic tuff have sparse σ-type phenocrysts, shear bands, and a well-developed elongation lineation. The upper thrust is thrust is marked by several meters of granitic mylonite and ultramylonite and >100 m of adjacent protomylonite. An elongation lineation in mylonite and ultramylonite defined by quartz strings, biotite (±chlorite) clots, and feldspar porphyroclasts. Mylonitic fabrics consistently show a northeast-side-up sense of shear (Figure 9). Similar fabrics occur in protomylonitic diorite along the southern continuation of the thrust system near Bute Inlet.

Figure 9

Photographs and equal area stereonets, Cumsack thrust system. All photographs with kinematic indicators are cut parallel to lineation and perpendicular to foliation. (a) Field photograph of thrust faults. (b) Outcrop of mylonitic granodiorite in upper plate near sample 15MR21 (161.45 ± 2.0 Ma, Table 1), examples of kinematic indicators highlighted; (c) granitic mylonite in Cumsack thrust; 2 km NW of (b), asymmetry of mylonitic fabric suggests NE-side-up sense of shear; (d) Microphotograph of σ -type porphyroclast from sample shown in (c); (e) Equal area stereonets of bedding, folds, foliations and lineations. Mean mylonitic foliation and cleavage is 305, 64NE (red).

Throughout the pendant, bedding in volcanogenic strata is generally parallel to the thrusts, varying locally where folded or truncated at low angles along fault segments (Figure 9). Folds were recognized only in the well-exposed middle plate where the largest has a wavelength of ∼20 m and smaller folds may be parasitic to it. The folds are open to tight, inclined plunging, and the largest fold verges southwest. Axial planes strike northwest and dip steeply northeast, subparallel to mylonitic foliations and hinge lines plunge gently to the northwest (Figure 9). The geometry of bedding, folds, and mylonitic fabrics is consistent with development within a southwest-directed thrust system which has steepened during or after formation.

A structurally higher part of the Cumsack thrust system may occur to the northeast, where elongate NW-striking lenses of varied rock types were mapped (Roddick & Woodsworth, 1977, 2006) (Figure 6). Our reconnaissance indicates most rocks are unfoliated to schistose diorite complex with minor foliated dacite-rhyolite resembling strata of Cumsack pendant and mylonitic granodiorite sills, one of which is 182 Ma (Cecil et al., 2018). Foliations are near vertical and NW-striking and lineations are down dip and locally more prominent than foliation. Regional mapping (Roddick & Tipper, 1985; Roddick & Woodsworth, 1977, 2006) shows the shear zone crossed by 79–78 Ma plutons (Cecil et al., 2018) but difficult access prevented us from confirming this relation.

The age of thrust faulting is not well constrained. New U-Pb zircon ages show mylonitic plutons in the thrusts were intruded in Late Jurassic; mylonitic diorite near Bute Inlet is 155.2 ± 1.7 Ma (14MR27, Table 1) and mylonitic granodiorite thrust over the Cumsack pendant is 161.45 ± 2 Ma (15MR21, Table 1, Figures 6 and 9). Titanite from this sample is 78.6 ± 3.2 Ma (Table 1). The large gap between the zircon and titanite U-Pb ages is unusual in the Coast Mountains (Cecil et al., 2018) and may record recrystallization and related processes during formation of the mylonite (Moser et al., 2022, 2023). If the regional mapping is correct (Roddick & Tipper, 1985; Roddick & Woodsworth, 2006), cross-cutting 79-78 Ma plutons place an upper limit on the age of the thrust system.

3.4.2 Interior Domain

East of the Cumsack thrust system, poor exposure obscures a significant geologic break which we conclude is the southern continuation of the CSZ (Figures 1 and 2). Evidence of the shear zone and its history are discussed after the geology of the Interior domain (Figure 10).

Figure 10

Geologic map of Interior domain with equal area stereonets. Plutons fabrics (a–b), structures in Waddington core (c–e), and Mt Raleigh pendant (f–g). Gray-scale contour interval = 2 standard deviations. Units match Figure 2; EWTB = eastern Waddington thrust belt, Od = Determination orthogneiss, Op = Pagoda orthogneiss, Os = Sisyphus orthogneiss. Intrusive ages labeled; italics show pluton ages of Cecil et al., 2018. Green dots: hornblende geobarometry and pluton age (Figure 2 and Table 3). Map based on data from this study; Woodsworth (1979), Roddick and Tipper (1985), Rusmore and Woodsworth (1993), Mustard et al. (1994), Cui et al. (2017), and Cecil et al. (2018).

3.4.2.1 Intrusive and Metamorphic Rocks in the Interior Domain

Intermediate to felsic plutons compose much of the Interior domain. Granodiorite forms more than 60% of the mapped pluton area with granite, tonalite, and quartz diorite each composing about 10% of the pluton area. The distinctive diorite complex present in the Coastal domain is absent and mafic compositions are rare (Figures 2 and 10). Plutons are generally large, nearly half are >100 km², and sub-equant in map view with irregular but steep contacts. Most plutons are unfoliated, compositionally homogeneous, and have few mafic enclaves. Sparse magmatic foliations generally strike NW and dip steeply (Figure 10a). Overall, the plutons are larger, more felsic, and more homogenous than in the Coastal domain.

Two large early Paleogene plutons are elongate, in contrast to the more equant plutons typical of the Interior domain (Figure 10). The 53 Ma Tiedemann pluton, the largest in this domain, strikes northwest for ∼80 km and is 5–15 km wide (Cecil et al., 2018; Roddick & Tipper, 1985). Along its length, pluton contacts are near vertical, sharp, with arrays of dikes and abundant stoped blocks within 0.5–1 km of the contact. Internally the pluton is generally unfoliated granodiorite with minor granite, lacks significant compositional layering or mafic inclusions, and appears undeformed in outcrop and thin section. Cutting obliquely across structures and metamorphic fabrics in rocks on both sides and along its length, the Tiedemann pluton limits the age of significant deformation in much of the Interior domain to pre-53 Ma.

On its eastern side, the Tiedemann pluton intrudes a mildly deformed northwest-striking pluton (Roddick & Tipper, 1985) which we refer to as the Mantle Glacier pluton (Figure 10). Published ages (Cecil et al., 2018) show it is 63–61 Ma, negating prior correlation to an older orthogneiss body (Rusmore & Woodsworth, 1994). The pluton is elongate (>15 km long and ∼5 km wide) with steep contacts where observed on its eastern side and as inferred from map patterns (Roddick & Tipper, 1985). This shape is mimicked by the internal compositional layering and foliations. The composition varies across the pluton, from foliated tonalite in the northeast to a tonalite-granodiorite sill complex in the southwest. The sills are 0.5 m to a few meters thick, strike northwest, and are vertical. A pervasive magmatic foliation defined by crystal shapes also strikes NW and is nearly vertical (Figure 10b). A subsolidus foliation is also parallel to magmatic foliation and sills (Figure 10b). In thin section, minor subsolidus recrystallization is shown by nearly euhedral crystals with irregular grain boundaries typical of grain boundary migration, fringes of neoblastic quartz and plagioclase and rarely, hornblende. A weak down-dip lineation marked by aligned crystals is present locally and compositional layers form symmetric boudins locally. Rare thin (cm) mylonitic shear zones parallel the foliation. Sparse kinematic indicators show both reverse and normal slip with a few conjugate pairs suggesting flattening strain (Rusmore & Woodsworth, 1994). Although not strongly deformed, the Mantle Glacier pluton accommodated subhorizontal, SW-NE directed shortening and is the only record of widespread ductile deformation between 63 and 53 Ma.

Metamorphic rocks in the Interior domain are grouped into three regions relative to the Tiedemann pluton: the Waddington core southwest of the pluton, the eastern Waddington thrust belt to its northeast, and the Mt. Raleigh pendant south of the end of the Tiedemann (Figure 10). Here we present new data on the deformation in these regions to develop an integrated structural-magmatic-metamorphic evolution of the Interior domain.

3.4.2.2 Waddington Core

Individual pendants in the Waddington core are ∼1–10 km thick and collectively form a northwest striking structural domain which is heavily intruded by Late Cretaceous plutons and significantly covered by ice. Rocks in all pendants are intensely deformed and metamorphosed to amphibolite facies. Most pendants are gray banded orthogneiss, granitic orthogneiss, and quartzofeldspathic schist to gneiss with lesser pelite (Figure 11). Rare metaconglomerate preserves relict clasts, including a few granitic clasts. Amphibolite is rare and quartzite and marble are absent. Intrusive protoliths are inferred for the gray banded gneiss and granitic orthogneiss; U-Pb zircon ages on these protoliths are 144.1 ± 2.9 on granitic orthogneiss (15MR50, Table 1) and 162.5 ± 2.1 (Cecil et al., 2018) on gray banded gneiss shown in Figure 11. Attempts to date granitic clasts or detrital zircon from the metaclastic rocks were unsuccessful.

Figure 11

Field photographs of pendants in the Waddington core. (a) Typical gray-banded orthogneiss, 162.5 ± 2.1 Ma (Cecil et al., 2018). (b) Meta-conglomerate with clast-defined axial planar foliation. (c) Quartzofeldspathic gneiss with rare, kinematically late, melt-filled dextral shear zones. Outcrop surface is horizontal; foliation and shear zones are near-vertical. Pen for scale.

Schistosity and gneissic foliation are defined by amphibolite-facies metamorphic minerals, compositional segregations, and lithologic layering. Metamorphic minerals are moderately aligned within the foliation parallel to an elongation lineation and also partially overprint these fabrics. Sillimanite, recognized in only one area, occurs as small crystals and fine-grained fibrolite within the foliation and lineation. Aggregates of fibrolite prisms up to 4 cm long in the foliation appear to be pseudomorphs of andalusite. Fibrolite-rich layers also form microfolds with fibrolite along the axial planes. Biotite, white mica, and amphibole lie within and across the foliation. Garnet porphyroclasts vary from irregular and highly embayed to idioblastic; inclusions of quartz and plagioclase are unoriented to helicitic. Garnet Sm-Nd ages indicate metamorphism occurred between 82 and 65 Ma (Bollen et al., 2022). Textures in dated garnet cores and rims record deformation before or during garnet growth at 82 and 72 Ma followed by static garnet growth until 65 Ma (Bollen et al., 2022). Collectively, the textures in a variety of minerals and the garnet ages show metamorphism accompanied and outlasted development of the foliation, lineation, and small-scale folds.

The foliation and lineation are folded on all scales, from thin section to map. Folds visible in outcrops are tight to isoclinal, most verge to the northeast and have a half-wavelength of a cm to ∼10 m, although vaguely defined larger isoclines are locally visible in mountainsides The folds are non-cylindrical with hinge lines varying from gently plunging to vertical (Figure 10d). Axial planes of outcrop-scale folds strike northwest and dip steeply southwest, parallel to the most steeply dipping foliations (Figure 10d); this pattern is caused by both isoclinal folding and development of an axial planar foliation. Where a foliation is very strongly developed, folds are difficult to discern in outcrop and the dominant foliation is inferred to be the axial planar foliation. Because foliations can only be reliably distinguished where both occur in one outcrop, the foliation shown on maps and stereonets is considered a composite foliation that developed throughout contractional deformation. A more complex and poorly understood structural history is preserved in a thin (∼250 m thick) selvage of schist between undated plutons east of the Waddington Glacier. Refolded folds are common, axial planes strike northeast and hinge lines are vertical and parallel to hornblende-defined lineations in one of the adjacent plutons. This local structural complexity appears related to emplacement of the adjacent plutons but cannot be resolved at the scale of this study.

All structures are slightly modified by map-scale folds that are open, upright, gently NW-plunging with a half-wavelength of about 10 km. These regional folds are signaled by reversals of the dominant dip direction of the foliation (Figure 10) and vergence of smaller folds. Because they are nearly parallel to outcrop-scale folds, the regional folds cause only minor dispersion of outcrop-scale folds and lineations across the region and little disruption within km-scale domains.

Small, discrete ductile shear zones occur in the metamorphic rocks and plutons throughout the Waddington core. Within the metamorphic rocks the shear zones are generally <30 cm thick and are marked by mylonitic rocks or more commonly are filled by undeformed quartzofeldspathic dikelets (Figure 11). Shear zones in plutons generally consist of a few meters of protomylonite containing rare mylonitic zones less than 10 cm wide. Regardless of expression, the small ductile shear zones are northwest-striking and nearly vertical with sparse, down-dip lineations (Figure 10e). A sense of shear was determined in 16 small shear zones from mylonitic foliations (S-C), porphyroclasts shear zone-fold relations, and reconstruction of separation of layers in gneisses. Thirteen are reverse, two are reverse-dextral or dextral and one is reverse-sinistral. Overall, the Waddington core contains sparse small reverse shear zones that are parallel to axial planes of folds on all scales. Collectively, these structures record northeast-southwest shortening and steeply northwest-plunging elongation.

Relations with intrusive bodies show progressive development of these structures in the Late Cretaceous. The oldest structures, the regionally developed foliation and lineation, are older than a small 77 Ma pluton (Cecil et al., 2018) which crosscuts them and contains stoped blocks of gneiss and orthogneiss near the pluton margin. This pluton lacks a penetrative foliation but is locally cut by small discrete mylonitic shear zones. Discrete mylonitic shears also occur in a 74 Ma granodioritic sill (14MR40, Table 1) about 50 m thick which lacks a penetrative foliation and cuts slightly obliquely across the gneissic foliation. Some outcrop-scale folding occurred after intrusion of two small sills at 74 Ma (14MR46, 14MR59, Table 1) which are folded with the gneiss. Both sills have many grains with high U/Th values which may indicate metamorphic growth ca. 74 Ma. The sills and structures are older than a granodiorite pluton which intruded them at 72 Ma (Cecil et al., 2018) at ~20 km depth (6.8 kb, 14MR61, Table 3). This post-kinematic pluton is irregularly ovoid in map view, mostly unfoliated with a local magmatic foliation parallel to its contacts and lacks the discrete mylonitic shears present in older intrusions. Its contact with the gneiss is steep, sharp and cuts across the gneiss foliation and outcrop-scale folds. Stoped blocks of gneiss are common at the contact and become increasingly bleached and diffuse within about 25 m of the contact. These relations between structures and plutons show crustal shortening began prior to 77 Ma and ended by 72 Ma, compatible with the syn- to postkinematic metamorphic ages (Bollen et al., 2022).

3.4.2.3 Mt Raleigh Pendant

The Mt. Raleigh pendant (Figure 10) preserves a metamorphosed remnant of a Late Cretaceous northeast-directed thrust belt (Bollen et al., 2022; Cecil et al., 2018; Woodsworth, 1979; Woodsworth et al., 2000). The structurally lowest unit is the granitic Sisyphus orthogneiss above which are two compositionally distinct thrust sheets of metamorphosed volcanic and sedimentary rocks (Figure 12). Along the lowest thrust fault, lenses of Sisyphus orthogneiss are imbricated with metavolcanic and metaclastic rocks of the overlying thrust sheet. Deformation within the thrust sheet precludes reconstruction of the stratigraphy but metasedimentary rocks are more common in the higher parts of the thrust sheet (Woodsworth, 1979). Lithologic similarities support correlation of these units to Early Cretaceous clastic and intermediate composition volcanic rocks of the Cloud Drifter and Ottarasko formations in the EWTB (Mustard et al., 1994; Rusmore & Woodsworth, 1993, 1994). The structurally highest thrust is defined by an abrupt contrast in rock types (Woodsworth, 1979). Above it, the highest thrust sheet consists of 2–3 km of mafic metavolcanic rocks that pass upward into marble and schist; these units are likely correlative to Upper Triassic Mt. Moore volcanics and carbonate units involved in the EWTB (Rusmore & Woodsworth, 1993; Tipper, 1978; Umhoefer et al., 1994). Syn-kinematic plutons, one of which is 84 Ma, occur in this structurally highest thrust sheet (Woodsworth, 1979; Woodsworth et al., 2000).

Figure 12

Photographs, Mt. Raleigh pendant. (a) Field view thrust faults and the three thrust sheets. View is toward the southeast. (b) Andalusite crystals in weakly foliated pelite in the eastern pendant, view is down onto foliation plane. (c) Foliated sillimanite-andalusite schist with weakly aligned sillimanite + andalusite clots with post-kinematic garnet, 5 km SW of (b). (d) Strongly foliated and lineated sillimanite-bearing schist adjacent to (c); compare to (b); white clots are sillimanite pseudomorphs of andalusite, view is down onto foliation plane.

The Sisyphus orthogneiss preserves the oldest deformation and metamorphism recognized in the Interior domain. Pinkish coarse-grained granitic orthogneiss is the most common and distinctive lithology but the unit includes gray orthogneiss and amphibolite. The granitic protolith is 110 Ma (Cecil et al., 2018) and has a penetrative gneissic foliation which is crosscut by unfoliated felsic garnet-bearing dikes and veins. Bollen et al. (2022) report Sm-Nd ages on two sizes of garnet from one dike. The larger garnet are ca. 99 Ma; smaller garnet yielded ca. 90 Ma ages interpreted to reflect diffusive re-equilibration (Bollen et al., 2022). Small discrete ductile shear zones and discontinuous shear bands cut these garnet-bearing dikes and the gneissic foliation. These relations indicate that penetrative foliation in the orthogneiss formed between 110 and 99 Ma and that metamorphism was either polyphase or lasted until at least ca. 90 Ma. Subsequent deformation, including development of the shear bands, incorporation into the thrust belt, and folding obscure the origins of the earliest deformation which is not shared by structurally higher thrust sheets.

Development of the thrust belt is illuminated by the structural and metamorphic evolution of the metaclastic-volcanic thrust plate. Deformation within this thrust plate is linked to metamorphism which increases from east to west across the pendant through a series of isograds (Figure 10) and likely occurred in two pulses (Bollen et al., 2022; Woodsworth, 1979). An early andalusite-bearing assemblage preserved in the eastern part of the pendant reached peak conditions of 3.4 ± 0.3 kbar and 527 ± 50°C to 560 ± 20°C at ca. 90 Ma (Bollen et al., 2022). To the west, this assemblage is replaced by sillimanite-bearing assemblages with peak conditions of 4.0–4.4 kbar and 640–670°C at ca. 72 Ma (Bollen et al., 2022; Woodsworth, 1979). Similarly, rocks within thrust sheet become more strongly deformed from east to west. The eastern part of the pendant is only lightly deformed, with weak foliation locally developed and a rare weakly developed elongation lineation. Large (∼5–10 cm long) andalusite crystals appear unoriented in pelitic outcrops where foliation is absent and are weakly oriented where foliation is stronger (Figure 12). Where a weak mica-defined lineation is present, andalusite crystals are not notably aligned with the lineation but those parallel to the lineation are mildly boudinaged. These textures suggest the metamorphism that produced andalusite at ca. 90 Ma was a largely static event which has been lightly overprinted by younger deformation in the eastern exposures.

This early static fabric is progressively modified as metamorphic grade increases to the west. The foliation and lineation become more strongly developed and are penetrative in the sillimanite-grade rocks (Figure 12). Sillimanite-bearing assemblages define a moderate to strong foliation and lineation and andalusite is a relic phase. Garnet both lies within and overprints the foliation. Sillimanite and biotite wrap around garnets within the foliation and form well developed pressure shadows. Microfolds of sillimanite-defined foliations are locally preserved in the pressure shadows, suggesting the foliation is a composite feature that formed during sillimanite-garnet metamorphism. This observation is consistent with outcrop-scale isoclinal to tight folds with axial planes parallel to foliation and a locally developed axial-planar foliation defined by sillimanite and is difficult to distinguish from the older foliation. Hinge lines of these folds are parallel to a widespread elongation lineation (Figures 10f and 10g). In contrast to the lower grade rocks, the sillimanite and relict andalusite define this lineation (Figure 12) and locally form L-tectonites. The increase in the alignment of relic andalusite crystals and sillimanite pseudomorphs is attributed to rotation of the crystals into the elongation direction during synkinematic metamorphism. The parallelism of the hinge lines to lineation also likely reflects this rotation. Development of these structures ended before the peak of metamorphism at 72 Ma which produced statically grown garnet (Figure 12) and vertical northwest-striking sillimanite-in and cordierite-in isograds that crosscut the foliation (Figure 10) (Bollen et al., 2022; Woodsworth, 1979).

Throughout the Mt Raleigh pendant, foliation, lineation, and thrust faults have been systematically reoriented by a post-metamorphic synform mapped by Woodsworth (1979) and a newly recognized antiform to its west (Figure 10). The folds are evident in the map pattern of thrust faults and variations in the average orientation of foliation and lineation in domains across the synform and antiform (Figure 10). These folds, with a half wavelength of ∼5 km, are open, steeply inclined to upright, and plunge ∼60° south. The age of these folds is limited by map patterns of isograds. The easternmost and lowest-grade isograd, the staurolite-in isograd, traces the eastern half of the synform and appears folded, consistent formation of staurolite with andalusite at ca. 90 Ma. The cordierite and sillimanite-in isograds that represent the 72 Ma peak of metamorphism cross the western limb of the synform, limiting the age of folding to >72 Ma. Similarly, the folds and a thrust fault are intruded by the 72 Ma Bishop River (Woodsworth, 1979). In summary, thrust faulting, penetrative deformation, and subsequent map-scale folding began after 90 Ma, was likely ongoing at 84 Ma and ended before 72 Ma. Metamorphism preceded, accompanied, and outlasted this deformation.

3.4.2.4 Eastern Waddington Thrust Belt

Prior work showed that the EWTB formed through northeast-directed thrust faulting and related deformation that was active by 87 Ma and ended before 68 Ma (Mustard et al., 1994; Rusmore & Woodsworth, 1994). Metamorphism outlasted deformation (Rusmore & Woodsworth, 1994) with peak amphibolite facies metamorphism at 65 Ma (Bollen et al., 2022). New mapping and dating of post-kinematic plutons revise this timing, showing that development of the thrust belt ended by 75–78 Ma. In the northern part of the field area, two post-kinematic plutons cut across most of the EWTB, forming a nearly continuous northeast-trending band <30 km long (Figure 10) (Roddick & Tipper, 1985; Rusmore & Woodsworth, 1993, 1994). The plutons are granodioritic to tonalitic in composition, medium grained, unfoliated, and internally homogeneous with few mafic enclaves or syn-intrusive mafic dikes. Their contacts are sharp and well defined in the field and on satellite imagery; locally dikes emanate from the pluton into wall rocks. No mylonitic shear zones or significant brittle faults were observed during mapping at 1:10,000–1:5000 scale within these plutons. The westernmost pluton is 75 Ma (Cecil et al., 2018) and on its western edge, intrudes the 76 Ma Determination tonalitic orthogneiss (Cecil et al., 2018; Rusmore & Woodsworth, 1994) (Figure 10). This 75 Ma pluton continues east >20 km, intruding across the EWTB imbricate zone, the 87 Ma Pagoda tonalitic orthogneiss (Rusmore & Woodsworth, 1994), and a regional antiform (Figure 10) (Roddick & Tipper, 1985). Less than 0.5 km from the eastern end of this pluton, a 78 Ma pluton intrudes across thrust faults, isograds, and regional synform (Figure 10) (Cecil et al., 2018; Roddick & Tipper, 1985; Rusmore & Woodsworth, 1993, 1994). Regional mapping shows a similar relation ∼50 km southeast along strike of the EWTH, where a 75 Ma granodioritic pluton intrudes amphibolite gneiss and the imbricate zone east of the Mantle Glacier pluton (Figure 10) (Roddick & Tipper, 1985). These relationships between structures, metamorphic rocks, and post-kinematic plutons show that the eastern Waddington thrust belt was active between 87 and 76 Ma and ended by 75 Ma. Locally, deformation ended slightly earlier (by 78 Ma) toward the foreland, consistent with out-of-sequence faulting noted by Rusmore and Woodsworth (1994). These new age constraints show that peak metamorphic conditions were reached 10–12 m.y. after contractional deformation ceased within the EWTB.

3.4.3 Structural Evolution of Interior Domain: Waddington Thrust Belt

Similarities in the style and timing of contractional deformation support the interpretation that the Waddington core, the Mt. Raleigh pendant, and the eastern Waddington thrust belt form a single northeast-directed Late Cretaceous thrust belt >130 km long and 75 km wide, which we refer to as the Waddington thrust belt (Figure 10). Integration of results from these three regions reveals the onset and duration of crustal thickening and contraction in the Waddington thrust belt (Figure 13). Thrust-related crustal loading began after the static low-pressure metamorphism at 90 Ma in the Mt. Raleigh pendant (Bollen et al., 2022). Emplacement of synkinematic plutons at 87 Ma and 84 Ma signals the onset of thrust faulting (Rusmore & Woodsworth, 1994; Woodsworth et al., 2000). By 80 Ma, metamorphic rocks in the Waddington core had been buried to ∼20–25 km (Bollen et al., 2022) which likely represents the peak of thrust faulting; penetrative structures in the core formed before 77 Ma and synkinematic plutons in the Waddington thrust belt are no younger than 76 Ma. By 72 Ma, contractional deformation had ceased throughout the thrust belt and pressures were lower, although comparison across different geobarometric systems adds uncertainty to estimates (Figure 13). In the Waddington core, hornblende geothermobarometry on two 72 Ma post-kinematic plutons gives results of 6.8 ± 0.3 kb (14MR61, Table 3; Cecil et al., 2018) where garnet equilibration pressures in adjacent gneisses are 4.5–6.5 kb (Bollen et al., 2022). The other pluton yielded pressures of 5.6 ± 1 kb (14MR72, Table 3; Cecil et al., 2018) and nearby gneisses yield garnet equilibrations pressures of 5.5–7.25 kb (Bollen et al., 2022). At Mt. Raliegh, post-kinematic garnet records pressures of 4.2 kb at ca. 72 Ma (Bollen et al., 2022), an increase from 90 Ma prior to onset of thrusting. Comparison along the strike of the thrust belt suggests post-kinematic pressures change by ∼2.5 kb over ∼70 km, requiring a slope of <1%. This along-strike decrease in pressure thus records minor a variation in exhumation after the end of contractional deformation. Overall, pressure estimates show moderate burial during development of the thrust belt between ca. 90 and 72 Ma, with the deepest rocks exposed in Waddington core, the rearward part of the Waddington thrust belt.

Figure 13

Pressure variation during development of Waddington thrust belt. Colors denote regions within the thrust belt. Pressure inferred from hornblende (Table 3) and garnet geobarometry (Bollen et al., 2022); ages are garnet Sm-Nd (Bollen et al., 2022) and U-Pb zircon ages of plutons (Cecil et al., 2018).

Metamorphism continued for about 10 m.y. after the end of contractional deformation and pressure remained steady or decreased slightly toward the foreland. Pressure estimates at ca. 72 Ma in the Waddington core are indistinguishable from a result of 5.7 ± 0.7 kb from the ca. 62 Ma Mantle Glacier pluton in the EWTB (Figures 10 and 13). Pressures appear lower toward the foreland, where a 60 Ma post-kinematic pluton in the easternmost EWTH yields pressure of 3.9 ± 0.5 kb (sample 2727R08, Figure 2, Table 3). In contrast, a ∼2 kb pressure increase is recorded in a gneiss from the Waddington core (Bollen et al., 2022). This sample is from the structurally complex selvage of gneiss discussed above and the pressure increase is interpreted to reflect local processes related to emplacement of adjacent plutons. Overall, the data show that the end of contractional deformation at 72 Ma was not followed by significant changes in pressure as metamorphism peaked ca. 65 Ma.

3.5 Boundary Between the Coastal and Interior Domains: The CSZ

Significant geologic contrasts between the Coast and Interior domains show they are juxtaposed along a crustal-scale shear zone which is nearly vertical and strikes northwest for >100 km across our study area. It projects northwestward along strike for an additional ∼100 km (Figure 1) into the southernmost known exposure of the CSZ (Rusmore et al., 2001). Based on its location and history as described below, we infer the CSZ is the boundary between the Coastal and Interior domains, expanding the known length of the CSZ to more than 1,400 km. The CSZ likely continues southeastward (Figure 1) from our area within the active Paleocene arc as postulated by Monger and Brown (2016).

In our study area, the CSZ is recognized by abrupt changes in rock type, metamorphic grade, thrust vergence, and disruption of geochronologic trends but the shear zone itself is nowhere exposed (Figures 2, 10, and 14). Truncation of the diorite complex marks the shear zone for much its ∼100 km mapped length. The diorite complex is a fundamental and widely exposed component of the Coastal domain and is absent in the Interior domain (Figure 2). Adjacent to the CSZ, the diorite complex is 81 Ma and is undeformed with a locally developed chlorite-epidote overprint (Figures 2, 4, and 10). Contrasting metamorphic grade also defines the CSZ; greenschist facies or lower grade rocks are adjacent the southwest of the CSZ; to the northeast are schist and gneisses of the Waddington thrust belt where amphibolite facies metamorphism persisted until 65 Ma (Bollen et al., 2022). Late Cretaceous structures are also affected by the CSZ. The dominant direction of thrusting reverses across the shear zone and the thrusts and late-stage regional folds of the Waddington thrust belt appear to be truncated by the shear zone (Figures 10 and 14).

Figure 14

Cross section between points A and B on Figure 2. Asz = Alpha Bluff shear zone, Bsz = Blind Creek shear zone, CSZ = Coast shear zone, Cts = Cumsack thrust system, Wsz = Wahkash shear zone. WTB = Waddington thrust belt, includes eastern Waddington thrust belt, Waddington core, and Mt. Raleigh pendant. Units match Figure 2.

Regional patterns in the magmatic and cooling ages of plutons are disrupted by the CSZ (Figure 15). The northeastward migration of the arc is recorded by the decrease in pluton ages to the northeast (Cecil et al., 2018). This trend is best seen in the Coastal domain where the youngest U-Pb ages decrease steadily to the northeast (Cecil et al., 2018), although a scarcity of ages between −40 and −20 km reduces the fidelity of this pattern. At the CSZ, the pattern changes; plutons within the Interior domain do not systematically become younger to the northeast and the CSZ is the southwestern limit of plutons younger than ca. 70 Ma (Figure 15). Similarly, throughout the Coastal domain, ⁴⁰Ar/³⁹Ar and K/Ar ages become younger to the northeast, mimicking the pattern of the youngest magmatic ages (Rusmore et al., 2013). At the CSZ however, biotite ⁴⁰Ar/³⁹Ar and K/Ar ages step from ca. 70 Ma on the southwest to 52-47 Ma to the northeast. Moreover, both existing and newly measured biotite ages show no systematic spatial change within the Interior domain. We interpret this contrast to transient heating-induced Ar loss related to pluton intrusion during the 61- 48 Ma high-flux magmatic event (Cecil et al., 2018) east of the CSZ.

Figure 15

Geochronology shown by distance from Coast shear zone. New ⁴⁰Ar/³⁹Ar ages (Table 2) shown with heavy outlines, others from compilation in Rusmore et al. (2013). U-Pb zircon ages from Cecil et al. (2018).

The geologic contrasts documented across the CSZ suggest it was active in the Paleocene and accommodated significant dextral displacement. Strike-slip displacement is required by truncation of the diorite complex; its absence in the Interior domain indicates strike-slip displacement exceeds the ∼100 km mapped length of the shear zone. This displacement is younger than the 65 Ma metamorphism of rocks juxtaposed against the diorite along the CSZ and may have begun after formation of the diorite complex (>77 Ma). The disruption of geochronologic patterns across the CSZ support displacement after 70 Ma and before 50 Ma, and sparse mylonitic fabrics in 64–62 Ma plutons adjacent to and east of the shear zone may reflect distributed strain related to displacement on the CSZ. Slip on the CSZ likely ended prior to 53 Ma when the Tiedemann pluton intruded these plutons. The displacement is likely to be dextral based on Late Cretaceous—Paleocene plate models that favor dextral displacement on northwest-striking shear zones (e.g., Engebretson et al., 1985; Stock & Molnar, 1988) and a sparse record of dextral slip on the northern CSZ (Andronicos et al., 1999; Boghossian & Gehrels, 2000; Gehrels, 2000; Ingram & Hutton, 1994; McClelland & Mattinson, 2000; Stowell & Hooper, 1990).

4 Discussion

4.1 Deformation and Magmatic Tempo Across the Batholith

Our new results on the deformation of the batholith provide insight into the relative roles of magmatism and deformation on batholith growth. Deformation along this transect occurred in distinct phases summarized on Figure 16. Occurring during growth of the batholith, these events are temporally and spatially linked to focused periods of high magmatic flux between 120 and 50 Ma, as calculated by Cecil et al. (2018) (Figure 16). From 114 Ma to 102 Ma, high flux magmatism occurred in two regions, both of which were deformed during this time. Sinistral transpression occurred after 117 Ma and before 103 Ma in the Coastal domain and domain and granitic orthogneiss was deformed between 110 and 99 Ma in the Interior domain. Neither high-flux magmatism nor deformation is recorded outside of these two regions (Figure 16). Following a lull in both deformation and magmatism, significant crustal contraction took place during the next HFE at ca. 85–70 Ma. This HFE and contraction are best developed in the Interior domain, where the Waddington thrust belt (87–72 Ma) formed in the region affected by the high-flux magmatism (Figure 16). In the Coastal domain, the Cumsack thrust system occurs in the location of high flux magmatism and appears to be coeval with it, although the timing of deformation is not well constrained. The final HFE (61–48 Ma) affects only the Interior domain and the CSZ forms its western boundary. Deformation during this HFE includes post-65 Ma slip on the CSZ and mild strain recorded in 62–64 Ma plutons, but the spatial extent of this final phase of deformation is poorly known. Removing significant dextral displacement on the CSZ would separate the Coastal and Interior domains but leaves intact the spatial-temporal correlation of deformation and magmatism within each domain. The northeastward migration of these events, however, is only seen in the Coastal domain; deformation and magmatism remained stationary within the Interior domain from ca. 114 to 48 Ma. Overall, these results show that active deformation along the transect was limited to areas undergoing high-flux magmatic events and matches their tempo throughout ∼60 m.y. of batholith growth.

Figure 16

Comparison of age and location of periods of deformation and high magmatic flux from Cecil et al. (2018) across the batholith transect. Shading of colors shows uncertainty in timing of deformation.

A correspondence between magmatism and deformation has long been noted in Cordilleran batholiths and has supported models for coevolution (e.g., Andronicos et al., 1999; Hamilton, 1988; Hollister & Crawford, 1986; Ingram & Hutton, 1994; Paterson & Miller, 1998; Tikoff & Teyssier, 1992; Tobisch et al., 1995). Particularly pertinent are models accounting for high flux events to be controlled by upper-plate deformation; either by retroarc or forearc process (Cao et al., 2015; DeCelles et al., 2009; Ducea et al., 2009; Paterson & Ducea, 2015). In these models, crustal shortening introduces fluid and silica-rich supracrustal sediments to the magmagenesis zone which triggers high-flux magmatic events (Cecil et al., 2011, 2012; Girardi et al., 2012; Lackey et al., 2012; Samson et al., 1991; Wetmore & Ducea, 2011). The magma produced bears a crustal signature and HFEs occur during periods of crustal thickening within the batholith. Throughout the evolution of the southern CMB, however, magmas are derived from mantle sources even during HFE accompanied by thrust faulting and crustal thickening, suggesting that upper plate deformation does not control magmatic tempo in the SCMB (Cecil et al., 2018, 2021). Instead mantle processes such as cycles of mantle dehydration likely control the location and tempo of magmatism in the southern CMB (Cecil et al., 2021). We suggest that in turn, magmatic tempo controls the timing and location of deformation within this segment of the batholith. In this interpretation, magmatic tempo is controlled by mantle processes and spatially focused magmatic flare-ups serve to weaken the crust and localize upper plate deformation into the area of active magmatism. Localization of strike-slip faults in intrusions has been recognized in Cordilleran batholiths (e.g., Ingram & Hutton, 1994; Seymour et al., 2020; Tikoff & de Saint Blanquat, 1997; Weinberg et al., 2004); we extend this relationship to magmatic tempo and expand the influence of magmatic tempo to crustal shortening.

Comparison to the central segment of the CMB helps define the relation between magmatic tempo and deformation and its influence on the architecture of the batholith. Early Cretaceous sinistral shear zones in the central segment of the CMB occur in the westernmost batholith, west of and not connected to the coeval sinistral faults on the Bute-Waddington transect (Figure 1). The evolution of the faults, however, is similar: they were active between ca. 115 and 100 Ma during the early part of a protracted HFE from 120 to 78 Ma (Angen et al., 2014; Chardon et al., 1999; Gehrels et al., 2009; Nelson et al., 2011, 2012; Wang et al., 2022). The sinistral shear zones lie within synkinematic plutons intruded at 15–20 km depths and become younger to the northeast in synch with migration of the magmatic arc (Wang et al., 2022). This spatial-temporal linkage is interpreted as magma-enhanced deformation during a HFE in the upper plate of an obliquely convergent margin (Wang et al., 2022). Thus, on both transects, development of the shear zones was restricted to areas undergoing high-flux magmatism. Differences in the location of the high-flux magmatism affected the location of the shear zones; in the central segment of the batholith, the HFE is farther west than in the southern batholith. This difference caused the shear zones to form in segments, rather than as a continuous shear zone. Segmentation of strike-slip faults is common in Cordilleran batholiths forming through oblique-subduction (e.g. Glazner, 1991; Tobisch et al., 1995; Tikoff & de Saint Blanquat, 1997; Seymour et al., 2020). We suggest this segmentation may be fostered by the spatial-temporal linkage of faulting with HFE. Regionally, this relationship may also explain the distribution of segments of Early Cretaceous sinistral shear zones east of the CMB where they formed in areas of active arc magmatism (Greig, 1992; Hurlow, 1993; Israel et al., 2006; Wang et al., 2022), although the magmatic tempo is not known. Regardless, the distribution of the shear zones shows the influence of sinistral oblique convergence across a wide part of the continental margin.

A similar spatial-temporal dependence of deformation on magmatic tempo is difficult to discern in the central CMB during Late Cretaceous contraction. Unlike the Bute—Waddington transect, sinistral and contraction phases of deformation are not temporally nor spatially distinct; thrust faulting is at least as old as 110 Ma and occurs in the same region as the sinistral faults (Angen et al., 2014; Cook & Crawford, 1994) and the timing of contraction does not appear to migrate with the active magmatic front across the >100 km wide thrust belt (Andronicos et al., 2003; Angen et al., 2014; Cook & Crawford, 1994; Cook et al., 1991; Friedman et al., 2001; Gareau, 1991; Rusmore et al., 2005; Woodsworth et al., 2020). The along-strike variation in the relation between contraction and magmatic tempo suggests that other processes, such as crustal depth, contribute to the development structures. Unlike the sinistral faults where the influence of magmatic tempo is well defined, the thrust belts on these transects are exposed at different crustal levels. On the Bute—Waddington transect, thrusts in the locus of the HFE formed ~3–7 kb, significantly shallower than those in the central CMB (~8–10 kb). At this shallower crustal level, the focused addition of heat and fluid from high-flux magmatism may locally weaken the crust and foster deformation, such as suggested in the Atacama region and Sierra Nevada (Seymour et al., 2020; Tikoff & de Saint Blanquat, 1997). These effects would be muted at deeper crustal levels where conditions favor ductile deformation, thus obscuring the temporal-spatial patterns between magmatic tempo and deformation. At slightly shallower crustal levels, such as in the central Sierra Nevada, the effects may also be muted by the overall lower temperature and heterogenous crustal state (Attia et al., 2022).

One implication of our interpretation regards the fidelity of deformation as a recorder for relative plate motions. If, as we propose, the timing and location of deformation at moderate crustal levels is controlled by high-flux magmatism, then deformation in the batholith will reflect both the magmatic processes and the relative plate motions along the margin. The dependence of the timing of deformation on magmatic tempo limits the utility of deformation to detect the timing of shifts in relative plate motions, especially on a single segment of the batholith. For example, the lack of sinistral displacement prior to ca. 120–115 Ma would not signal onset of sinistral oblique subduction at ca. 120 Ma, but instead reflect the onset of HFE magmatism. Recognition of this limitation removes a mismatch in plate models which call for sinistral motion as old as 160 Ma and the rock record (Monger & Gibson, 2019). Similarly, the variation in the timing of contractional deformation along the arc would reflect variations in magmatic patterns driven by lower-plate processes, not along-strike shifts in subduction rates or angles. Given the uncertainties in plate reconstructions between ca. 120 and 50 Ma, our results suggest that timing of deformation within the batholith is not a reliable constraint on the timing of changes in plate motions.

4.2 Dextral Displacement of the Western Batholith

Large dextral displacement on the CSZ would have significant impacts on both magmatic architecture and the Baja B.C. controversy. Our results require more than 100 km of dextral slip and higher displacements would be consistent with the great length of the shear zone and its position in an obliquely convergent margin. Upper limits on displacements are difficult to discern. Latest Cretaceous piercing points are not found along the >1,400 km-long shear zone. Any disruptions of magmatic trends caused by large displacements would be muted by the sub-parallel strike of the CSZ and arc trends and difficult to distinguish from along-strike variations within the batholith (Cecil et al., 2018, 2021). Geologic relations discussed below support moderate displacements, on the order of a few hundred km.

Restoration of a few hundred km of dextral slip would simplify the distribution of terranes in southern CMB, where fragments of the Intermontane terrane occur west of rocks with paleomagnetic inclinations typical of the Insular terrane (Cecil et al., 2018, 2021; Monger & Gibson, 2019; Rusmore et al., 2013; Tikoff et al., 2023). Because the CSZ obliquely transects the batholith, this restoration brings the western flank of the batholith southeastward into a more inboard position aligned with the larger Intermontane terrane (Rusmore et al., 2013) and reduces the paleomagnetically inferred northward translation of plutons at Knight Inlet to values typical of the Intermontane terrane. It also restores pre-120 Ma arcs into positions found in the central segment of the CMB (Gehrels et al., 2009) and helps align sinistral faults in the western batholith with coeval sinistral faults east of the batholith (Greig, 1992; Hurlow, 1993; Israel et al., 2006; M. G. Miller, 1988). Occurring after 70 Ma, this dextral displacement post-dated assembly of the Insular and Intermontane terranes and may be an expression of the “run” phase of dextral translation of terranes along western North America proposed by Tikoff et al. (2023). Because the CSZ lies between rocks with similar large (2,300–2,700 km) inferred northward translations, the proposed moderate displacement cannot resolve the Baja-BC controversy but does create a more coherent pattern of displaced terranes and sinistral shear zones in southern British Columbia to serve as a starting point for understanding prior translation and terrane amalgamation.

5 Conclusions

Cretaceous—Eocene deformation of the southern CMB occurred in three spatially distinct phases coincident with three high-flux magmatic events. Development of sinistral shear zones between ca. 114 Ma and 103 Ma coincides with the 114–102 Ma HFE in the Coastal domain and penetrative deformation of the Sisyphus orthogneiss (110–99 Ma) in the Interior domain. Outside of these two areas, neither high-flux magmatism nor deformation is recorded. The subsequent HFE (85–70 Ma) was accompanied by orogen-perpendicular contraction which is especially well developed in the Interior domain (87–72 Ma) and was followed by dextral slip on the CSZ during the final HFE from 61 to 48 Ma. The CSZ, active after 65 Ma and likely before 53 Ma, is the western edge of the area affected by this HFE. Plutons within the area of the 61–48 Ma HFE also record distributed strain between 64 – 53 Ma. The spatial and temporal coincidence of deformation and HFE's implies that generation of mantle-derived magmas during HFE's modulated the tempo and location of deformation in the southern CMB. We propose that these spatially focused magmatic flare-ups weaken the crust, localizing upper plate deformation into the area of active magmatism at the moderate crustal levels exposed in the southern CMB. In the central segment of the CMB, Early Cretaceous sinistral faulting is localized by a coeval HFE (Wang et al., 2022) but this relation is unclear during widespread Late Cretaceous contraction. These variations are attributed to differences in the depth of crust exposed, suggesting the response of deformation to magmatism is affected by crustal level. The last stages of batholith growth are marked by development of the intra-batholithic CSZ, which is now recognized for more than 1,400 km along the CMB. Our results strengthen evidence for >100 km of Late Cretaceous dextral slip along this shear zone, contributing to northward displacement of terranes westernmost British Columbia. Restoration of moderate dextral slip on the CSZ would remove an apparent doubling of pre-100 arcs and Intermontane terrane in southern segment of the batholith, a situation analogous to the Salinian block on the San Andreas fault system in California.

Acknowledgments

The authors thank Intan Yokelson and Occidental College undergraduates Robert Bogue, Charles Chisom, Marshall Troutman for assistance in the field, and Ma Chi (Caltech Micropobe Lab) laboratory assistance. This research would not have been possible without superb logistical support from SilverKing Ventures and Whitesaddle Air Services. We thank Paul Umhoefer and Robert Miller for valuable discussions about the southern CMB. Reviews by Dr. Dan Gibson and Dr. Felix Gervais greatly improved the manuscript, and we thank Associate Editor Dr. Dawn Kellet and Editor Dr. Taylor Schildgen for their comments and editorial assistance.

Open Research

Data Availability Statement

All data supporting the conclusions in this paper may be accessed at https://doi.org/10.5061/dryad.jsxksn0kt.