Mesozoic terranes of the central Cascades: Geology ...

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The Geological Society of America Field Guide 49

Mesozoic terranes of the central Cascades: Geology of the Hicks Butte complex, Easton Metamorphic Suite, Peshastin Formation and Ingalls ophiolite complex James H. MacDonald Jr. Department of Marine and Ecological Sciences, Florida Gulf Coast University, 10501 FGCU Boulevard, South Fort Myers, Florida 33965, USA Joe D. Dragovich Associated Earth Sciences, Inc., 1552 Commerce Street, Suite 102, Tacoma, Washington 98402, USA Mark E. Pecha Department of Geosciences, University of Arizona, Tucson, Arizona 85721, USA Glenn T. Thompson Shanna C. Stingu Alex D. Maruszczak Kristy M. Zalud Scott H. Milliken Department of Marine and Ecological Sciences, Florida Gulf Coast University, 10501 FGCU Boulevard, South Fort Myers, Florida 33965, USA ABSTRACT This paper reviews the Mesozoic terranes in the central Cascades, south of the Windy Pass thrust and east of the Straight Creek–Fraser River fault, and provides a guide to field locations for these units. These include the Easton Metamorphic Suite, Hicks Butte complex and higher-grade tectonic zone, the Peshastin Formation, and the Ingalls ophiolite complex (also known as the Ingalls terrane). Age data, whole rock and mineral chemistry, and structural data are reviewed. These oceanic- and arcaffinity terranes formed outboard of the North American craton during the Jurassic and accretion likely occurred during the Late Jurassic or Early Cretaceous. They were then dextrally translated north and emplaced in Washington State during the Late Cretaceous. A better understanding of these Mesozoic terranes will more closely constrain the tectonic development of the North American Cordillera.

MacDonald, J.H., Jr., Dragovich, J.D., Pecha, M.E., Thompson, G.T., Stingu, S.C., Maruszczak, A.D., Zalud, K.M., and Milliken, S.H., 2017, Mesozoic terranes of the central Cascades: Geology of the Hicks Butte complex, Easton Metamorphic Suite, Peshastin Formation and Ingalls ophiolite complex, in Haugerud, R.A., and Kelsey, H.M., eds., From the Puget Lowland to East of the Cascade Range: Selected Geologic Excursions in the Pacific Northwest: Geological Society of America Field Guide 49, p. 79–100, doi:10.1130/2017.0049(05). © 2017 The Geological Society of America. All rights reserved. For permission to copy, contact [email protected].

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INTRODUCTION

EASTON METAMORPHIC SUITE

Washington State includes numerous Mesozoic terranes (Fig. 1). While many of these units have been studied in detail (e.g., Misch, 1966; Brandon et al., 1988; Miller et al., 1993; Brown and Gehrels, 2007; MacDonald et al., 2008a; MacDonald and Schoonmaker, 2017), they are commonly omitted from Cordilleran tectonic syntheses. These Mesozoic terranes originated outboard of the North American craton and likely accreted during the Late Jurassic or Early Cretaceous. Many researchers suggest accretion of these terranes occurred south of their current position in Washington State (Jett and Heller, 1988; Garver, 1988; Miller et al., 1993; MacDonald et al., 2008a). This would require northward dextral translation of these terranes after accretion. It is well documented that these terranes were emplaced by thrust faulting in Washington State during the Late Cretaceous (Fig. 1) (Miller, 1985; Tabor, 1994; Brown, 2012; Sauer et al., 2014). The extent of northward translation of Washington Mesozoic terranes has long been debated. Some researchers propose that these terranes correlate to terranes in California and Oregon and have been translated 100s of km (Brown and Blake, 1987; Garver, 1988; Jett and Heller, 1988; Miller et al., 1993; MacDonald et al., 2008a). Other researchers, primarily using paleomagnetic data from Cretaceous plutons that intrude some of these terranes (e.g., Mount Stuart batholith), suggest they originated near the modern-day Baja California and were translated more than 3000 km (the Baja British Columbia hypothesis; Beck, et al., 1981; Cowan et al., 1997; Housen et al., 2003). Palinspastic restoration of post–mid-Cretaceous strike-slip faulting places Washington Mesozoic terranes near proposed correlative terranes in northern California and southern Oregon (e.g., Josephine ophiolite and overlying Galice Formation) (Wyld et al., 2006). This review concentrates on Mesozoic terranes south of the Windy Pass thrust and east of the Straight Creek–Fraser River fault (Fig. 1). These include: the Easton Metamorphic Suite, which is located in both the Kachess Lake and Hicks Butte inliers (Figs. 2 and 3); the Hicks Butte complex and higher-grade equivalent tectonic zone (Fig. 3); and the Peshastin Formation and serpentinites from the Ingalls ophiolite complex (also called Ingalls terrane) (Figs. 4 and 5). Detailed studies of these oceanic and arc affinity rocks have been undertaken by Stout (1964), Southwick (1974), Ashleman (1979), Miller (1985), Treat (1987), Miller and Mogk (1987), Miller et al. (1993), Metzger et al. (2002), MacDonald et al. (2008a, 2008b), MacDonald and Dragovich (2015), and Metzger et al. (2016). Absent from our review, but also occurring in this region, are the Mesozoic terranes within the Manastash and Rimrock Lake inliers (Fig. 1) (Stout, 1964; Miller, 1989; MacDonald and Schoonmaker, 2017). Understanding the origins of these terranes, due to their complex development and possible regional correlations, is critical to understanding the Mesozoic tectonic development of the North American Cordillera. We conclude with a road log and stops for a two-day field trip from Seattle that examines these rocks.

The Easton Metamorphic Suite (Tabor et al., 1993) consists of the Shuksan Greenschist, Darrington Phyllite, and semischist and phyllite of Mount Josephine (Misch, 1966; Brown, 1986; Tabor et al., 2003). However, the Mount Josephine unit does not occur in the central Cascades (Figs. 1, 2, and 3). In the northwest Cascades, the Easton Metamorphic Suite overlies rocks of the Excelsior and Welker Peak nappes along the Shuksan thrust (Fig.  1) (Misch, 1966; Brown and Blake, 1987; Brown, 2012; Tabor and Haugerud, 2016). In the Kachess Lake inlier, the Easton Metamorphic Suite structurally overlies the North Peak unit. The North Peak unit has been correlated with the Chilliwack Group of the Excelsior nappe (Fig. 2) (Ashleman, 1979; Tabor and Haugerud, 2016). The Shuksan Greenschist protolith is interpreted to have been stratigraphically under the Darrington Phyllite protolith (Haugerud et al., 1981; Brown, 1986). Fe-Mn metalliferous metasediments locally occur between the Shuksan Greenschist and Darrington Phyllite (Ashleman, 1979; Haugerud et al., 1981; Brown, 1986). Epidote-bearing blueschist occurs throughout the Easton Metamorphic Suite. It commonly transitions to greenschist. Brown (1974) suggested that the bulk composition of the protolith, particularly the Fe 3 + content, determined whether blueschist or greenschist mineral assemblages were stable. Higher-grade schists are present locally (Brown et al., 1982). The Easton Metamorphic Suite in the northwest Cascades (Fig. 1) has 163–150 Ma Jurassic protolith ages, 160–144 Ma higher-grade metamorphic ages, and 130–120 Ma regional blueschist metamorphic ages (Armstrong, 1980; Brown et al., 1982; Armstrong and Misch, 1987; Gallagher et al., 1988; Dragovich et al., 1998; Brown and Gehrels, 2007). No age data have been generated from the Easton Metamorphic Suite in the Kachess Lake or Hicks Butte inliers. Shuksan Greenschist in the Central Cascades Shuksan Greenschist in the central Cascades occurs in the Kachess Lake and Hicks Butte inliers (Figs. 2 and 3). It consists of interlayered cm-wide compositional bands of fine-grained, wellrecrystallized greenschist and epidote blueschist. Thin section textures are primarily lepidoblastic to nematoblastic. Generally, greenschist consists of albite-epidote-chlorite-quartz-actinolite, whereas blueschist has epidote-Na-amphibole-quartz-albite with lesser muscovite and chlorite (Lofgren, 1974; Ashleman, 1979; Treat, 1987; Miller et al., 1993; Tabor et al., 2000; Maruszczak et al., 2016). Ashleman (1979) observed rare relic igneous textures in Shuksan from the Kachess Lake inlier; however, no relic textures have been reported from the Hicks Butte inlier. Blueschist is more common than greenschist in the Kachess Lake inlier (Lofgren, 1974; Ashleman, 1979), whereas greenschist is more common than blueschist in the Hicks Butte inlier (Treat, 1987; Miller et al., 1993). In the Kachess Lakes inlier, foliations in the Shuksan trend NW-SE and generally dip steeply to the SW (Lofgren, 1974;



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= Late Jurassic arc rocks

HH = Helena-Haystack mélange ING = Ingalls ophiolite complex NWCT = other Northwest Cascade terranes SCFRF = Straight Creek-Fraser River fault

Manastash 47°00´N inlier

WA N

TS = Twin Sisters ultramafic complex WM = western mélange belt

BP US 97

BP = Blewett Pass

ING

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Figure 1. Simplified geologic map displaying pre-Cenozoic tectonic elements of the central and northwest Cascades, modified from Miller et al. (1993), Haugerud and Tabor (2009), and MacDonald and Schoonmaker (2017). Jurassic arc rocks are displayed with dark-gray patterned fill. Ophiolitic and ultramafic rocks are black. WA—Washington.

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Figure 2. Simplified geologic map of the Kachess Lake inlier. Modified from Lofgren (1974), Ashleman (1979), Tabor et al. (2000), and Haugerud and Tabor (2009). Note the locations of Stops 1-6 and 1-7 on this figure. Cross section A–B is shown on map.

Ashleman, 1979; Tabor et al., 2000). Foliation generally follows the anti- and synformal structures that deform the overlying Eocene strata (Fig. 2). In the Hicks Butte inlier, foliation in the Shuksan also strikes NW-SE and dips moderately to the SW (Treat, 1987; Miller et al., 1993; Tabor et al., 2000). Synmetamorphic folds, crenulations, and crosscutting shear zones deform the Shuksan foliation in the Hicks Butte inlier, and synmetamorphic mineral lineations in the Shuksan are subparallel to foliation. Maruszczak et al. (2016) conducted electron probe microanalyses (EPMA) of amphiboles and epidotes from a Shuksan blueschist in the Kachess Lake inlier and classified amphiboles using Hawthorne et al.’s (2012) scheme. The analyzed amphiboles

are predominantly glaucophane with one winchite (Fig. 6A). These amphiboles have low TiO2, high NaM4 (Fig. 6B), little AlIV substitution for Si (Fig. 6B), and XMg (Mg /[Mg+Fe2+]) averages 0.52 (Maruszczak et al., 2016). Epidotes from the Kachess Lakes blueschist have high X Ep (≥0.83; Fe3+/[Fe3++Al+Cr-2]) (Fig.  7) and, an average X Fe of 0.28 (Fe3+/[Fe3++Al+Cr+Mn]) (Maruszczak et al., 2016). An epidote overgrowth has X Fe of 0.70 (Maruszczak et al., 2016). Maruszczak et al. (2016) also noted a retrograde greenschist-facies metamorphism overprinting the blueschist mineral assemblages from the Kachess Lake inlier sample. Treat (1987) conducted EPMA of amphiboles from Shuksan greenschist and blueschist in the Hicks Butte inlier.



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Figure 3. Simplified geologic map of the Hicks Butte inlier. Modified from Treat (1987), Miller et al. (1993), Tabor et al. (2000), and Haugerud and Tabor (2009). The eastern position of the Ainsley Canyon anticline is modified from Cheney and Hayman (2009). Note the locations of Stops 1-1 to 1-5 on this figure. Cross section A–B is shown on map. Pz—Paleozoic.

Amphiboles from Shuksan greenschist in the Hicks Butte inlier range from actinolitic hornblende to magnesio-hornblende, while amphiboles from Shuksan blueschist in the Hicks Butte inlier are glaucophane with one winchite (Treat, 1987). Ashleman (1979) conducted whole-rock major and trace element geochemistry of Shuksan greenschist and blueschist from the Kachess Lake area with X-ray fluorescence (XRF).

He ground XRF samples in both iron and tungsten mills, which likely resulted in variable contamination with FeO, Cr, Co, and Nb. In addition, many immobile elements reported by Ashleman (1979) are at or below XRF detection limits. This indicates his data set should be used with care. Based on his results, Ashleman (1979) suggests the Shuksan greenschist and blueschist in the Kachess Lake inlier have ocean-floor geochemical affinities.

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Figure 4. Simplified geologic map of the southern portion of the Ingalls ophiolite complex. Map modified from MacDonald et al. (2008a) and MacDonald and Dragovich (2015). Cross section modified from Miller (1985) and MacDonald et al. (2008a). Gray dashed-line box outlines the area of Figure 5. Cross section A–B is shown on map. MORB-IAT—mid-ocean ridge basalt–island arc tholeiite; WPB— within-plate basalt.



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Figure 5. Simplified geologic map of the eastern portion of the Ingalls ophiolite complex. Units are the same as in Figure 4. Map modified from Harper et al. (2003), MacDonald et al. (2008a), and MacDonald and Dragovich (2015). Cross section C–D (shown on map) drafted by A.N. Mlinarevic and used with permission. Note the locations of Stops 2-1 to 2-4 on this figure. I—Iron Mountain; K—King Creek area.

Darrington Phyllite in the Central Cascades Darrington Phyllite in the central Cascades consists predominantly of carbonaceous quartzose phyllite and muscovitechlorite-albite-quartz schist that locally contains lawsonite (Lofgren, 1974; Goetsch, 1978; Ashleman, 1979). Ashleman (1979), Treat, (1987), and Miller et al. (1993) proposed that the protolith was a carbonaceous mudstone. Dungan et al. (1983) reported locally preserved primary structures in Darrington Phyllite west of the Straight Creek–Fraser River fault (Fig. 1); however, none are reported from the central Cascades. Foliations in the Darrington Phyllite located in the Kachess Lakes inlier strike NW-SE and generally dip steeply to the SW (Lofgren, 1974; Ashleman, 1979; Tabor et al., 2000). Similar to the Shuksan in this inlier, the orientation of foliation follows the large folds that deform the inlier and the overlying Eocene rocks (Figs. 2 and 3). Again, the foliation of the Darrington Phyllite in the Hicks Butte inlier strikes NW-SE, and generally dips moderately to the NE or SW (Treat, 1987; Miller et al., 1993; Tabor et al., 2000). Small folds deform the foliation and create

a crenulation lineation. Open, possibly late (possibly Cenozoic), folds deform the original fabric of the Darrington Phyllite. The foliation and folding is consistent with the antiformal structure folding the Darrington Phyllite (Fig. 3). Some local cataclasites occur in the Darrington Phyllite in the Hicks Butte inlier. HICKS BUTTE COMPLEX The Hicks Butte complex of Treat (1987) (Fig. 3) consists of variably deformed hornblende tonalite and hornblende quartz diorite with lesser hornblende gabbro, diabase, dacite, and rare hornblendite and trondhjemite (Stout, 1964; Treat, 1987; Miller et al., 1993; Tabor et al., 2000; Stingu et al., 2015). Diabase, dacite, hornblendite, and trondhjemite cut other lithologies as 1-cm–0.5-m-thick intrusive dikes. Dacite is allotrimorphicgranular in thin section, while the other lithologies are predominantly hypidiomorphic-granular. Adjacent to the Hicks Butte complex is a tectonic zone, which consists of tonalitic gneiss, amphibolite, gabbroic gneiss, and dacitic gneiss. The gneisses are porphyroclastic and

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A

2.5 Eckermannite

1.5 Al-Barroisite / Mg-Taramite

B

Na

2.0

Figure 6. (A) Kachess Lake inlier Shuksan Na-amphiboles plotted on the classification diagram modified from Schumacher (2007). The black squares represent the compositions of the adjacent amphibole names. apfu—atoms per formula unit. (B) Kachess Lake inlier Shuksan Na-amphiboles plotted on the NaM4 versus AlIV amphibole diagram modified from Brown (1977). (C) Hicks Butte complex Ca-amphiboles plotted on the classification diagram modified from Hawthorne and Oberti (2007).

Glaucophane

1.0

Richterite

Winchite

Tremolite

0.0 -1.0

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Ca2Al2Cr3+Si3O11(O/OH)

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Tschermakite Actinolite

Magnesio-Hornblende

0.5 FerroActinolite

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Ca2Al2Fe3+Si3O11(O/OH)

Figure 7. Kachess Lake inlier Shuksan epidote plotted on the epidote classification diagram.

4 kb

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Ferro-Hornblende

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Ferro-Tschermakite

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6.0

Si (apfu) Kachess Lake inlier Shuksan blueschist

Hicks Butte complex Intrusive high Sr/Y Extrusive high Sr/Y Intrusive low Sr/Y

5.5

granoblastic in texture, while the fine-grained amphibolite is primarily nematoblastic. Though he did not map it separately, Stout (1964) interpreted this tectonic zone to be a higher metamorphic grade equivalent of the igneous rocks of the complex. Frizzell et al. (1984) and Tabor et al. (2000) defined and separated the tectonic zone from the Hicks Butte complex as a recognizable unit in the field (Fig. 3). Like Stout, they indicated the tectonic zone was derived from the Hicks Butte complex. Treat (1987) also suggested the tectonic zone consisted of higher metamorphic grade equivalents of Hicks Butte complex lithologies. However, Miller et al. (1993) suggested the amphibolite of the tectonic zone was formed by the intrusion of the Hicks Butte complex possibly into the Easton Metamorphic Suite. Miller et al. (1993) also interpreted the lithologies of the tectonic zone to have been derived from both the Easton Metamorphic Suite and the Hicks Butte complex. The foliation of the Hicks Butte complex and tectonic zone strikes NW-SE with a steep to moderate SW dip. Common stretching lineations have a NW trend and a shallow to moderate NW plunge. The amphibolites are L-tectonites. Boudins in the Hicks Butte complex and tectonic zone indicate stretching was layer-parallel. Less deformed igneous dikes are folded into strongly lineated shear zones in the Hicks Butte complex. The

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A

500 400

Age (Ma)

strength of the mineral lineation increases toward the contact between the tectonic zone and the Easton Metamorphic Suite (Treat, 1987; Miller et al., 1993). In places, the tectonic zone and the Easton Metamorphic Suite appear to grade into each other (Treat, 1987; Miller et al., 1993; Tabor et al., 2000). Miller et al. (1993) reported a multigrain zircon age of 153 ± 3 Ma U-Pb from a tonalite gneiss in the tectonic zone (Fig. 3). MacDonald and Pecha (2011) reported laser-ablation inductively coupled plasma-mass spectrometry (LA-ICP-MS) U-Pb zircon ages from a hornblende quartz diorite and a dacite from the Hicks Butte complex (Figs. 3 and 8). Either cores or rims of zircons were analyzed. All zircons, except one ca. 154 Ma age zircon from the dacite, have low U/Th ratios. Eleven zircons from the hornblende quartz diorite yielded a mean age of 150.0 ± 6.8 Ma (2s; MSWD = 0.2) (Fig. 8A). Three zircon ages from this hornblende quartz diorite are Paleozoic (317 Ma, 356 Ma, and 464 Ma) (Fig. 8A). Twenty-one zircons from the dacite yielded a mean age of 144.0 ± 2.4 Ma (2s; MSWD = 1.1) (Fig. 8B). Five zircons from this sample have Late Jurassic U-Pb ages, and two have Middle Jurassic U-Pb ages (Fig. 8B). Tabor et al. (2000) obtained a hornblende K-Ar age of 127.7 ± 16 Ma from a Hicks Butte complex tonalite gneiss (Fig. 3). Stingu et al. (2015) conducted geochemistry of the Hicks Butte complex and tectonic zone (Figs. 3 and 9). Samples are subalkaline (Fig. 9A), calcic, low-K, and predominantly magnesian (Stingu et al., 2015). The extrusive samples are dacite with one rhyolite (Fig. 9A). The molecular Na 2O+K 2O/Al 2O3 of this rhyolite is much less than 1, indicating it is not peralkaline. The extrusive samples are predominantly felsic (Fig. 9A); mostly peraluminous; have high Sr/Y values (Fig. 9B); and have low Cr, Ni, and Mg# (Stingu et al., 2015). The extrusive samples predominantly plot in the field defined by adakites on the Sr/Y versus Y diagram (Fig. 9B). Two intrusive samples have the same geochemical affinities, including the high Sr/Y ratios (Fig. 9B). The remaining intrusive samples and gneisses and amphibolites are mafic to intermediate (Fig. 9A)—one sample is felsic and metaluminous—and have low Sr/Y values (Fig. 9B) (Stingu et al., 2015). Thompson et al. (2015) conducted EPMA analysis of amphibole and plagioclase from both high Sr/Y and low Sr/Y Hicks Butte complex samples (Figs. 6C and 10) using Hawthorne et al.’s (2012) amphibole classification. All amphiboles are calcium-rich. Amphiboles from the high Sr/Y dacite are ferritschermakite, whereas amphiboles from a high Sr/Y hornblende tonalite are magnesio-ferri-hornblende (Fig. 6C) (Thompson et al., 2015). The amphibole from low Sr/Y intrusive samples are magnesio-ferri-hornblende (Fig. 6C) (Thompson et al., 2015). Treat (1987) also analyzed amphibole from the Hicks Butte complex; however, she did not have the benefit of age or wholerock geochemical data. Treat (1987) indicates the amphiboles from the Hicks Butte complex range from magnesio-hornblende to tschermakite with minor cummingtonite. Additionally, Treat (1987) analyzed amphiboles from the higher-grade tectonic zone. These range in composition from magnesio-hornblende to ferro-tschermakite.

300

Hbld Qtz Diorite Low Sr/Y ratio 150.0 ± 6.8 Ma (2σ) Total error = 1.2% MSWD = 0.2

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190

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150 140 130 120

Figure 8. (A) LA-ICP-MS U-Pb zircon ages from a Hicks Butte complex hornblende quartz diorite. Error bars are 1s. The gray horizontal line represents the best-fit age of the 11 samples and its thickness includes the 2s error. Three ages are excluded from the best-fit age. Hbld—hornblende; MSWD—mean square of weighted deviations; Qtz—quartz. (B) LA-ICP-MS U-Pb zircon ages from a Hicks Butte complex dacite. Error bars are 1s. The gray horizontal line represents the best-fit age of the 21 samples and its thickness includes the 2s error. Seven ages are excluded from the best-fit age.

Thompson et al. (2015) also analyzed plagioclase from the Hicks Butte complex (Fig. 10). Plagioclases from the high Sr/Y rocks are primarily andesine, with lesser labradorite. Plagioclases from the low Sr/Y rocks are primarily bytownite with lesser labradorite. Treat (1987) analyzed plagioclase from both the Hicks Butte complex and the higher-grade tectonic zone and found labradorite in granitoids and oligoclase in an amphibolite from the tectonic zone. Thompson et al. (2015) used mineral compositions to estimate the temperate and pressure of crystallization of the Hicks Butte complex (Table 1). Temperature estimates from amphiboles follow Ridolfi et al. (2010) methods. Amphibole from the low Sr/Y samples yields temperatures of 828°–864 °C. Amphibole from high Sr/Y dacites yield temperatures of 934°–945 °C. The high Sr/Y hornblende tonalite yields a temperature of

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Figure 9. (A) Total alkali versus silica diagram for the Hicks Butte complex and tectonic zone. (B) Sr/Y versus Y diagram for the Hicks Butte complex and tectonic zone modified from Defant and Drummond (1990).

Sr/Y

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“Normal” island arc lavas

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Y Hicks Butte complex & tectonic zone High Sr/Y Extrusives Tonalite or Trondhjemite

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847° ± 56 °C (Table 1). Pressure estimates using the amphibole thermometer and barometer of Ridolfi et al. (2010), and the Al-in-hornblende barometer calibration of Anderson and Smith (1995), are reported in Table 1. Although the Hicks Butte complex lacks the buffer assemblage required for the Anderson and Smith (1995) calibration, its application yields hydrous intermediate to felsic magmas solidus to slightly hypersolidus pressures for Hicks Butte samples (Table 1). Hydrous intermediate pressures are expected for these samples based on their whole-rock major element geochemistry (Fig. 9A). The Ridolfi et al. (2010) barometer yields pressures that would cause these intermediate to felsic samples to plot closer to the solidus for Or

KAlSi3O8

Intrusive high Sr/Y Extrusive high Sr/Y Intrusive low Sr/Y

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hydrous mafic magmas—which has pressure and temperatures too high for the whole-rock magmatic compositions for the Hicks Butte complex. Further, the compositions and lack of zoning in the plagioclase (Fig. 10; Thompson et al., 2015)—compared to the wholerock geochemistry (Fig. 9A)—indicates the mineral assemblage may have been in equilibrium with the melt during crystallization. Two high Sr/Y samples did not meet the chemical requirements of Ridolfi et al. (2010) amphibole barometer (Table 1). In addition, the barometer of Ridolfi et al. (2010) estimates a pressure for the extrusive dacite that is noticeably higher than that of all intrusive rocks (~5.7 kbar for the dacite versus an average of ~1.9 kbar for intrusive rocks; Table 1). We report both Ridolfi et al. (2010) and Anderson and Smith (1995) pressures in Table 1, but only reference the latter herein. The low Sr/Y phase of the Hicks Butte complex yields pressures ranging from 0.66 to 1.80 kbar. The high Sr/Y dacite yields pressures between 0.21– 0.26 kbar. The high Sr/Y hornblende tonalite yields a pressure of 1.81 ± 0.06 kbar (Table 1).

Andesine

Labradorite

Bytownite

Anorthite

An

Ab

CaAl2Si2O8

NaAlSi3O8

Figure 10. Hicks Butte complex plagioclase plotted on the plagioclase classification diagram.

The polygenic Ingalls ophiolite complex, called the Ingalls terrane by Tabor and Haugerud (2016), is a nearly complete, yet highly dismembered ophiolite that primarily consists of ultramafic rocks (Fig. 4) (Miller, 1985; MacDonald et al., 2008a). It is intruded by the ca. 96–91 Ma Mount Stuart batholith and is locally overlain by Eocene sedimentary rocks (Fig. 4). Two large east-west–striking and north-dipping serpentinite fault zones, the Navaho Divide and Cle Elum Ridge fault zones (Fig. 4), disrupt the ophiolite into a large-scale mélange (Frost, 1975; Miller, 1985). The Ingalls ophiolite complex was thrust over the Cascade crystalline core along the Cretaceous Windy Pass thrust (Fig. 4). Miller (1985) demonstrated that thrust faulting was synchronous with the emplacement of the Mount Stuart batholith.

TABLE 1. GEOTHERMOBAROMETRY DATA FOR HICKS BUTTE COMPLEX Sample HB12U/Pb-1 HB12U/Pb-2A

Lithology

Sr/Y ratio

Amphibole classification*

Temperature (°C)†

Pressure (kbar)†

Pressure (kbar)#

dacite

High

ferri-tschermakite

934 ± 22

5.67 ± 1.42

0.21 ± 0.05

dacite

High

ferri-tschermakite

945 ± 56

n.d.

0.26 ± 0.05

HB23-BJ

hbl tonalite

High

magnesio-ferri-hornblende

847 ± 56

n.d.

1.81 ± 0.06

HB12A-1

hbl qtz diorite

Low

magnesio-ferri-hornblende

841 ± 22

1.70 ± 0.19

0.93 ± 0.16

HB12A-2

hbl qtz diorite

Low

magnesio-ferri-hornblende

828 ± 22

1.33 ± 0.15

0.66 ± 0.16

HB12B-1

hbl gabbro

Low

magnesio-ferri-hornblende

864 ± 22

2.12 ± 0.53

0.75 ± 0.16

HB12B-2

hbl gabbro

Low

magnesio-ferri-hornblende

864 ± 22

2.30 ± 0.57

0.96 ± 0.16

HB12B-3

hbl gabbro

Low

magnesio-ferri-hornblende

857 ± 22

2.17 ± 0.54

1.04 ± 0.16

HB22B-1-amph

hbl diorite

Low

magnesio-ferri-hornblende

853 ± 22

2.17 ± 0.54

1.48 ± 0.07

HB22B-4-GT

hbl diorite

Low

magnesio-ferri-hornblende

833 ± 22

1.77 ± 0.44

1.80 ± 0.23

*Amphibole classification based on Hawthorne et al. (2012) Amphibole thermometry and barometry of Ridolfi et al. (2010). # Amphibole-feldspar barometry of Anderson and Smith (1995). Abbreviations: hbl—hornblende; qtz—quartz; n.d.—not able to be determined due to failure to meet chemical criteria. †

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Three distinct mantle peridotite units, which have been variably serpentinized, were recognized by Miller (1985) and Miller and Mogk (1987) (Fig. 4). These include a southern dunite and harzburgite unit, lherzolite and hornblende peridotite that are overprinted by the high-temperature shear zone of the Navaho Divide fault zone, and a northern lherzolite unit (Fig. 4). Relatively intact oceanic crustal rocks exist as m- to km-scale fault blocks within the mélange of the Navaho Divide fault zone (Fig. 4). Miller (1985) and MacDonald et al. (2008a, 2008b) utilized lithology, age, and geochemistry to divide crustal parts of the Ingalls complex into the: (1) Early Jurassic Iron Mountain unit, (2) Late Jurassic Esmeralda Peaks unit, and (3) Jurassic Peshastin Formation (Fig. 4). The Iron Mountain unit consists of pillow basalt and broken-pillow breccia, with minor rhyolite, hyaloclastite, oolitic limestone and chert (MacDonald et al., 2008b). A U-Pb zircon age from an Iron Mountain unit rhyolite is ca. 192 Ma (Figs. 4 and 5) (MacDonald et al., 2008b). Radiolarian ages from chert within this unit are Pliensbachian (Early Jurassic; Fig. 4) (MacDonald et al., 2008b). This unit has within-plate basalt geochemistry that is transitional to enriched mid-ocean ridge basalt (Metzger et al., 2002; MacDonald et al., 2008b). MacDonald et al. (2008b) interpreted the Iron Mountain unit to have formed as an Early Jurassic seamount. The Esmeralda Peaks unit consists of basaltic flows and pillows, massive diabase with minor sheeted dikes, gabbro and rare tonalite, and trondhjemite (Miller, 1985; MacDonald et al., 2008a). Minor chert, as well as isolated and broken-pillow breccia, also occur within this unit (Miller, 1985; MacDonald et al., 2008a). A U-Pb zircon age from an Esmeralda Peaks unit gabbro is ca. 161 Ma (Figs. 4 and 5) (Miller et al., 2003). This unit has transitional island arc tholeiite to mid-ocean ridge basalt geochemical affinities with rare boninites (Metzger et al., 2002; MacDonald et al., 2008a). MacDonald et al. (2008a) interpreted the Esmeralda Peaks unit to have formed in a Late Jurassic back-arc basin, and suggested it may have originated from forearc rifting. Peshastin Formation The Peshastin Formation is mostly located in the eastern portion of the Ingalls ophiolite complex (Figs. 4 and 5). It consists of massive argillite with lesser chert, sandstone (lithic graywacke), pebble conglomerate, and pebbly mudstone as well as minor sedimentary serpentinite and ophiolitic breccia (Southwick, 1974; Miller, 1985; Harper et al., 2003; MacDonald and Dragovich, 2015). Peshastin argillite conformably overlies the ca. 192 Ma Iron Mountain unit and the ca. 161 Ma Esmeralda Peaks unit (Figs. 4 and 5) (Harper et al., 2003; MacDonald et al., 2008b). Peshastin chert that contains radiolarians of late Pliensbachian to Middle Toarcian age (“Early Jurassic” on Fig. 4) conformably overlies the Iron Mountain unit (Harper et al., 2003; MacDonald et al., 2008a). Peshastin chert that contains Oxfordian age radiolarians (“Late Jurassic” on Fig. 4) conformably overlies the Esmeralda Peaks unit (Miller et al., 1993; MacDonald et al.,

2008a). Detrital zircon from a sandstone yielded two U-Pb age peaks at ca. 152 and ca. 232 Ma (Miller et al., 2003). This sandstone is rich in volcanic detritus and compositionally immature, therefore Harper et al. (2003) interpreted the ca. 152 Ma age peak to approximate the rock’s depositional age. A sandstone from this U-Pb locality was collected and chemically analyzed by MacDonald and Dragovich (2015; see below). The Peshastin Formation was deformed and metamorphosed by emplacement of the Mount Stuart batholith (Fig. 4) (Paterson et al., 1994; Albertz et al., 2005). This is expressed by penetrative transposition of preexisting fabrics in a structural and thermal aureole extending more than 2 km out from the batholith (Paterson et al., 1994; Albertz et al., 2005). Lithic graywacke of the Peshastin Formation consists mostly of angular grains of chert, monocrystalline quartz, albite, and mafic to felsic volcanic and metavolcanic lithic clasts (Southwick, 1974; Miller, 1985; MacDonald and Dragovich, 2015). Several monocrystalline quartz clasts are bipyramidal (MacDonald and Dragovich, 2015). Other clasts occurring in the graywacke are lithic shale, siltstone, felsic to mafic plutonic rocks, and minor grains of clinopyroxene, amphibole, altered volcanic glass, diabase, perthite, epidote, and chromite (Southwick, 1974; Miller, 1985; MacDonald and Dragovich, 2015). Peshastin argillite consists of ~10%–30% very angular silt and sand that consists of quartz, plagioclase, and shale clasts (Southwick, 1974). MacDonald and Dragovich (2015) reported chemical analysis of sandstone and argillite from the Peshastin Formation (Fig. 11). Samples have low Th, Hf, and Zr; high V, Ni, and Sc; and plot in or near fields defined by modern back-arc basins. MacDonald and Dragovich (2015) noted that the composition of the Peshastin Formation was similar to that of sediments that have intermediate to mafic igneous provenance (Fig. 11). Serpentinite from the Ingalls Ophiolite Complex Serpentinite in the Ingalls ophiolite complex is both massive and highly sheared. Massive serpentinite occurs as completely or partly recrystallized peridotite that displays primary igneous textures (e.g., bastite serpentine pseudomorphs after pyroxene) (Frost, 1975; Miller, 1985). Sheared serpentinite occurs in fault zones that cut other ultramafic rocks locally and regionally (e.g., Navaho Divide fault zone; Fig. 4) (Frost, 1975; Miller, 1985; Miller and Mogk, 1987). Using the serpentinite thin section textures outlined by O’Hanley (1996), the Ingalls serpentinite have pseudomorphic, nonpseudomorphic, and transitional textures. Mesh, hourglass, and interlocking textures are common in thin section. Relic olivine and pyroxene have hypidiomorphicgranular textures. Some serpentinite close to the Mount Stuart batholith in the eastern portion of the ophiolite (Fig. 5) have tremolite overprinting older metamorphic textures. Cr-spinels can be well preserved or totally replaced by magnetite (Zalud and MacDonald, 2017). Krevor et al. (2009) identified the ultramafic rocks from the Ingalls ophiolite complex as a potential resource for mineral carbon dioxide sequestration.



Mesozoic terranes of the central Cascades

91

V

A

Mafic protoliths

Felsic protoliths

Ultramafic protoliths

Ni

(56.5*TiO2/Al2O3) - (10.879*Fe2O3T/Al2O3) + (30.875*MgO/Al2O3) - (5.404*Na2O/Al2O3) + (11.112*K2O/Al2O3) - 3.89

B

Th*10 8 6 4

Quartzose sedimentary provenance Mafic igneous provenance

2 0

Modern back-arc basin turbidites (n = 23)

-2 -4 -6

Intermediate igneous provenance

-8 -10

-10

Lau backarc basin volcaniclastic sediments (n = 45)

-8

-6

-4

-2

0

2

4

Felsic igneous provenance

6

(30.638*TiO2/Al2O3) - (12.541*Fe2O3T/Al2O3) + (7.329*MgO/Al2O3) + 12.031*Na2O/Al2O3) + (35.402*K2O/Al2O3) - 6.382 Darrington Phyllite metashale metasandstone

Peshastin Formation argillite sandstone

8

10

Figure 11. (A) Darrington Phyllite and Peshastin Formation samples plotted on the V-Ni-Th provenance diagram modified from Bracciali et al. (2007). (B) Darrington Phyllite and Peshastin Formation samples plotted on the discrimination function diagram of Roser and Korsch (1988). Also on this diagram are fields for modern back-arc basins. Both diagrams modified from MacDonald and Dragovich (2015).

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MacDonald et al.

1.0

Ingalls dunite & harzburgite

0.8

0.6

Forearc Peridotite

Abyssal Peridotites Ingalls lherzolite

g

0.4

0.2

Forearc peridotites

part

ial m

eltin

MgO/50

tinized pyroxene (bastite) has higher Ni concentrations than serpentinized olivine (Milliken and MacDonald, 2013). Burkhard (1993) suggested spinel group minerals would have their SiO2 greatly increased via serpentinization (SiO2 is normally very low to absent in spinel group minerals). The spinel group minerals from the Ingalls serpentinites have very low SiO2 values, suggesting serpentinization did not greatly affect their chemistry (Zalud and MacDonald, 2017). Spinels range from Cr-spinel to Al-chromite (Zalud and MacDonald, 2017). Zalud and MacDonald (2017) note the serpentinite spinels plot mostly in the overlap fields between abyssal peridotite and forearc peridotite, while spinels from two samples plot in the field defined by forearc peridotite (Fig. 13).

Cr/(Cr+Al)

Frost (1975) studied the contact metamorphism of Ingalls serpentinite and other ultramafic rocks in the western portion of the ophiolite complex (Fig. 4). He noted that contact metamorphism drove off water and converted hydrous serpentinite into forsterite-rich peridotite. The olivine in the contact metamorphosed ultramafic rocks became Mg-rich due to Fe loss by fluid interaction. In addition, Frost (1975) noted that the temperatures produced by the intrusion of the Mount Stuart batholith into the ophiolite (Fig. 4) were ~700 °C. Milliken and MacDonald (2013) analyzed Ingalls serpentinites for major and trace elements (Fig. 12). Major element ratios of serpentinites are similar to harzburgite. O’Hanley (1996) and Hattori and Guillot (2007) suggested serpentinization did not greatly alter rock chemistry due to the metamorphism being both low temperature and pressure. Milliken and MacDonald (2013) noted major element ratios of serpentinites were similar to harzburgite. Ingalls serpentinites plot in the fields defined by modern abyssal and forearc peridotites (Fig. 12) (Milliken and MacDonald, 2013). Several samples have high Al2O3 (Fig. 12). Milliken and MacDonald (2013) and Zalud and MacDonald (2017) analyzed Ingalls serpentinite minerals with electron probe microanalyses (Fig. 13). The serpentine is geochemically similar to lizardite, and displays minor substitution of Fe, Ni, and Cr for Mg, and minor Al substitution for Si (Milliken and MacDonald, 2013). X-ray diffraction analysis of powdered samples confirms lizardite is the serpentine polymorph present. Serpen-

0.0 1.0

0.5

2+

Mg/(Mg + Fe )

0.0

Spinel from serpentinite

Abyssal peridotites me

co

JHM-06-09 JHM-12-10 JHM-14-10

nta

tam

Y

High-T shear zone over Ingalls lherzolite & hbld peridotite

or

ph

ct

ism

Al2O3

Figure 12. Ingalls serpentinites plotted on an MgO/50 versus Al2O3 versus Y diagram. Field for abyssal peridotites is derived using data from Niu (2004). Field for forearc peridotites is derived using data from Parkinson and Pearce (1998). Contact metamorphic samples have tremolite in them. See Frost (1975) for discussion of Al enrichment of contact metamorphosed serpentinite.

JHM-17-10 JHM-19-10 JHM-26-10

Figure 13. Spinels from Ingalls serpentinites plotted on the Cr/(Cr+Al) versus Mg/(Mg+Fe2+) diagram modified from Dick and Bullen (1984). Fields for spinels from the Ingalls ophiolite complex peridotites (dunite; harzburgite; lherzolite; and high-T shear zone lherzolite and hornblende peridotite) are derived from Miller and Mogk (1987). Abyssal peridotite and forearc peridotite fields are modified from Dick and Bullen (1984), Metzger et al. (2002), and Arai et al. (2010). hbld— hornblende.



Mesozoic terranes of the central Cascades

DISCUSSION Easton Metamorphic Suite Rocks of the Easton Metamorphic Suite in the northwest and central Cascades are lithologically similar. This similarity has led most workers to suggest these rocks are correlative and offset by the Straight Creek–Fraser River and Darrington–Devils Mountain faults (Fig. 1). We therefore assume the metamorphic and protolith ages from the Easton Metamorphic Suite of the northwest Cascades are applicable to the central Cascades (Armstrong, 1980; Brown et al., 1982; Armstrong and Misch, 1987; Gallagher et al., 1988; Dragovich et al., 1998; Brown and Gehrels, 2007). The unnamed thrust fault separating the Easton Metamorphic Suite from the North Peak unit in the Kachess Lake inlier (Fig. 2) is here proposed to be a structural equivalent of the Shuksan thrust in the northwest Cascades (Fig. 1). Alternatively, this fault is the Excelsior Ridge thrust, and the Shuksan thrust is missing in this inlier. Another possibility is that the contact between the Easton Metamorphic Suite and North Peak unit is an early Cenozoic extensional fault, as suggested by Tabor and Haugerud (2016), for the overlying contact between the inlier and Eocene strata (see their figure 11.9, p. 149). The chemistry of amphibole and epidote from the central Cascade Shuksan is very similar to that from the northwest Cascades (Fig. 6B) (Brown, 1986; Maruszczak et al., 2016). Maruszczak et al. (2016), following the methodology of Brown (1977) and Evans (1990), used amphibole and epidote geochemistry, as well as mineral assemblages, to estimate that blueschist of the Shuksan in the Kachess Lake inlier crystalized between 300 and 400 °C and 6–7 kbar (Fig. 6B). In the Hicks Butte inlier, Treat (1987) used amphibole chemistry to suggest blueschist crystallized at 450° ± 100 °C and 9 ± 2 kbar, whereas greenschist crystalized at 500° ± 100 °C and 5 ± 1 kbar. Davis and Lindmark (2015) used PERPLEX modeling of bulk mineral compositions to propose that blueschist crystallized at temperatures 8 kbar, while greenschist crystallized at ~500 °C and 5–7 kbar. The pressure–temperature estimates for the Shuksan blueschist in the Kachess Lake and Hicks Butte inliers corroborate pressure–temperature estimates from older studies (Haugerud et al., 1981; Brown, 1986). The blueschist and greenschist from Treat’s (1987) and Davis and Lindmark’s (2015) studies are interlayered. The probability of interlayered Shuksan schists in the Hickes Butte inlier forming 4 kbar apart in pressure is not high. Dungan et al. (1983) indicated the Shuksan Greenschist in the northwest Cascades has mid-ocean ridge basalt (MORB) geochemical affinities. This agrees well with the ocean-floor determination of Ashleman (1979) for the Kachess Lake inlier of the central Cascades. However, Dragovich et al. (1998, 1999, and 2000) and Metzger et al. (2016) report that in addition to normal and enriched MORB, the Shuksan Greenschist also has calc-alkaline basalt, island-arc tholeiite, Fe-Ti basalts, and rare within-plate basalt geochemical affinities. Using geochemical

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affinities and lithologies, Brown and Blake (1987) proposed the Shuksan Greenschist originated in a back-arc basin. MacDonald and Dragovich (2015) demonstrated geochemistry of the Darrington Phyllite from the northwest Cascades formed in an arc-proximal back-arc setting that could have initiated by forearc rifting (Fig. 11). This suggests the Easton protolith could have originated as a Late Jurassic back-arc basin, and blueschist facies metamorphism occurred during Cretaceous subduction of this oceanic crust. The consistent NW-SE structural trend of rocks in the Kachess Lake and Hicks Butte inliers is intriguing. It could be that this fabric is related to Cretaceous emplacement of these inliers. This fabric appears congruent with Eocene syn- and antiformal structures, suggesting younger deformation is overprinting the emplacement structures of these inliers (Figs. 2 and 3). Therefore, it is plausible the NW-SE trend is the result of Eocene folding. Hicks Butte Complex The ca. 150 Ma U-Pb zircon age from a hornblende quartz diorite, combined with the Miller et al. (1993) ca. 153 Ma U-Pb zircon age from a tonalite gneiss, are taken to represent the crystallization age of the low Sr/Y portion of the Hicks Butte complex and the higher-grade equivalent tectonic zone (Figs. 3, 8A, and 9). The low U/Th ratios of zircons, indicating they are igneous, and the lack of differences between core and rim ages, both support this interpretation (MacDonald and Pecha, 2011). This suggests the low Sr/Y portion of the Hicks Butte complex was assembled incrementally between ca. 153 Ma and 150 Ma. Due to our single crystal LA-ICP-MS dating of zircons, we were able to identify Paleozoic zircon xenocrysts (Fig. 8A). These Paleozoic zircons also have low U/Th ratios, indicating they are igneous (MacDonald and Pecha, 2011). We suggest these Paleozoic zircons were entrained by the Late Jurassic magma during its emplacement, and, the Hicks Butte arc was built on older Paleozoic crust. The ca. 144 Ma zircon age (Fig. 8B) is interpreted to represent the crystallization age of the high Sr/Y portion of the Hicks Butte complex (Fig. 9). The LA-ICP-MS single crystal zircon dating allowed us to identify Middle and Late Jurassic age zircons in the sample (Fig. 8B). Except for one sample, these Jurassic age zircons have low U/Th ratios indicating they are igneous, and there are no discrepancies between core and rim ages (MacDonald and Pecha, 2011). This supports ca. 144 Ma age as a crystallization age. There are no other Early Cretaceous igneous ages in Washington State. However, Early Cretaceous K-Ar or Ar-Ar metamorphic ages are reported from the Easton Metamorphic Suite (146–141 Ma), the Hereford Meadow amphibolite, Manastash inlier (ca. 143 Ma; Fig. 1), and the Indian Creek complex, Rimrock Lake inlier (ca. 142; Fig. 1) (Armstrong and Misch, 1987; Miller et al., 1993; MacDonald and Harper, 2010). The tectonic zone is a higher-grade, pervasively deformed correlative of the low Sr/Y phase of the Hicks Butte complex.

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Chemistry of the low Sr/Y phase of the Hicks Butte complex and the tectonic zone follow the same fractionation trends shown on Figure 9. Stingu et al. (2015) also demonstrated that the Late Jurassic phase of the Hicks Butte complex and the tectonic zone have similar trace element geochemistry. Except for the degree of deformation, the lithologies of the tectonic zone are similar to the low Sr/Y phase of the Hicks Butte complex. Finally, the U-Pb zircon ages from the Hicks Butte complex and the tectonic zone are similar and within each other’s error. All of this further supports the correlation of the low Sr/Y phase of the Hicks Butte complex with the higher-grade tectonic zone. Thompson et al. (2015), based on amphibole and plagioclase mineral chemistry, suggested crystallization of the low Sr/Y phase of the Hicks Butte complex at 828°–864 °C and 0.66–1.80 kbar (Table 1). They also proposed the high Sr/Y extrusive samples crystallized between 934°–945 °C and 0.21–0.26 kbar, while an intrusive sample from the high Sr/Y phase crystallized at ~847 °C and ~1.81 kbar (Thompson et al., 2015). Using amphibole compositions, Treat (1987) estimated the temperate and pressure for the Hicks Butte complex to be 500°–550 °C and 5.4–6.0 kbar. However, she did not have knowledge of the polygenetic ages of the complex, and her estimates might include amphiboles that are millions of years apart in age. Davis and Lindmark (2015) used mineral assemblages to suggest that “gabbro-diorite” of the Hicks Butte complex cooled rapidly from 1100° to ~700 °C at pressures between 6 and 8 kbar. The lithology of this sample suggests it may have been from the low Sr/Y phase, however, this is not certain. These great variations in pressure and temperature for the Hicks Butte complex are problematic. More work is needed to clarify the crystallization depth of this complex. The geochemistry of the low Sr/Y phase of the Hicks Butte complex and tectonic zone (Fig. 9) are similar to modern island arcs (Frost et al., 2001). Stingu et al. (2015) proposed this arc was built on older rifted crust, as evidenced by the Paleozoic zircon ages (Fig. 8A). Volcanic rocks of the high Sr/Y phase of the Hicks Butte complex are adakites (Fig. 9B) (Defant and Drummond, 1990). Stingu et al. (2015) noted the low Cr, Ni, and Mg# of the adakites suggested they were derived from melting of lower arc crust. The entrained Middle to Late Jurassic zircons (Fig. 8B) may be related to the mafic source for these adakites (Stingu et al., 2015). The juxtaposition of the Hicks Butte complex, the amphibolite-facies tectonic zone, and the blueschist-facies Easton Metamorphic Suite (Fig. 3) has resulted in multiple interpretations (e.g., Treat, 1987; Miller et al., 1993; Tabor et al., 2000; Davis and Lindmark, 2015). The Hicks Butte complex does not contain the high P/T metamorphic mineral assemblage of the Easton Metamorphic Suite; however, several researchers suggest this is a result of its bulk composition not being able to express high P/T metamorphism (Miller et al., 1993; Tabor et al., 2000). Several researchers do suggest the Hicks Butte complex and the Easton Metamorphic Suite were faulted against each other at great depth (Tabor et al., 2000; Davis and Lindmark, 2015). The subparallel lineation to the foliation in the tectonic zone, which was also

stressed by Treat (1987) and Miller et al. (1993), suggests the tectonic zone experienced dextral as well as normal dip-slip motion. Eocene deformation needs to be accounted for when interpreting the structural relationship between these two units. More work is needed to clarify this relationship. Peshastin Formation A detailed study is needed to clarify Early and Late Jurassic ages of the Peshastin Formation (see discussion in MacDonald and Dragovich, 2015). Southwick (1974), Miller (1985), and MacDonald and Dragovich (2015) proposed that the Peshastin Formation was deposited in or near a volcanic arc. MacDonald and Dragovich (2015), using geochemistry (Fig. 11), further proposed that it was deposited in an arc-proximal back-arc basin that could have initiated by forearc rifting. MacDonald and Dragovich (2015)—based on the Late Jurassic and Late Triassic bimodal detrital zircon populations (Miller et al., 2003; Brown and Gehrels, 2007) and similar geochemistry (Fig. 11)—suggested the Peshastin Formation and Darrington Phyllite are correlative. Serpentinite from the Ingalls Ophiolite Complex Harzburgite may have been the primary protolith for analyzed serpentinites (Milliken and MacDonald, 2013). The chemistry of these serpentinites hints that their peridotite protoliths formed as the result of partial mantle melting under a mid-ocean ridge and/or a suprasubduction zone setting (Fig. 12) (Milliken and MacDonald, 2013). The majority of the crustal rocks in the Ingalls ophiolitic complex—the Late Jurassic Esmeralda Peaks unit and the Peshastin Formation—have been proposed to have formed in a back-arc setting, which would have both mid-ocean ridge and arc geochemical affinities (Metzger et al., 2002; MacDonald et al., 2008a; MacDonald and Dragovich, 2015). Therefore, we suggest the peridotite protoliths of the serpentinite also formed in a suprasubduction zone, possibly a back-arc setting based on whole rock and spinel geochemistry (Figs. 12 and 13). Miller and Mogk (1987) suggested the lherzolite and hornblende peridotite, which are overprinted by the Navaho Divide hightemperature shear zone (Fig. 4), originated as fracture zone. Tremolite in serpentinite samples in the eastern portion of the ophiolite complex suggests the ultramafics rocks underwent contact metamorphism similar to that described by Frost (1975) to the west. Milliken and MacDonald (2013) attributed the unusual Al-enrichment of serpentinite samples (Fig. 12) to contact metamorphism (see Frost, 1975, for this process). These Al-rich serpentinites also have tremolite in thin section. Spinels from the Ingalls serpentinite are well preserved and their chemistry was not affected by serpentinization (Zalud and MacDonald, 2017). Their compositions record varying degrees of partial melting (Fig. 13). Miller and Mogk (1987) analyzed spinel group minerals from the Ingalls peridotite units (Fig. 13). Spinels from serpentinites mostly plot near the field defined by the northern lherzolite unit. Two samples plot near or higher than



Mesozoic terranes of the central Cascades

the field defined by dunite within the southern harzburgite unit (Figs. 4 and 13). Zalud and MacDonald (2017) suggested that the northern lherzolite, as well as southern dunite and harzburgite units, were protoliths for the serpentinite. Summary Regional correlations of these Mesozoic terranes are fundamental for determining their original locations. The high Sr/Y ratio ca. 144 Ma (Figs. 8A and 9B) igneous rocks intruding low Sr/Y ratio ca. 153–150 Ma (Figs. 8B and 9B) igneous rocks in the Hicks Butte complex (Fig. 3) are analogous to the ratios and ages of the Dixie Butte complex and Bald Mountain batholith in the Blue Mountains of northeast Oregon. There, high Sr/Y 148–140 Ma plutons intrude or are associated with low Sr/Y 162–154 Ma plutons (Schwartz et al., 2011, 2014). The overall geochemistry of these Cretaceous and Jurassic rocks from the Blue Mountains is similar to the Hicks Butte complex (Schwartz et al., 2011, 2014; Stingu et al., 2015), indicating comparable original tectonic settings. Based on ages, crosscutting relations, and geochemistry, we suggest the Dixie Butte complex and Bald Mountain batholith are correlative with the Hicks Butte complex. MacDonald and Dragovich (2015), based on detrital zircon population ages and geochemistry, correlated the Darrington Phyllite of the Easton Metamorphic Suite with the Peshastin Formation (Figs. 2–5, 11). Based on identical Late Jurassic radiolarian faunal assemblages, similar detrital zircon population ages, and similar geochemistry, MacDonald et al. (2008a) and MacDonald and Dragovich (2015) correlated the Peshastin Formation with the Galice Formation of the Klamath Mountains in northern California and southern Oregon. The correlation of the Darrington Phyllite with the Peshastin Formation indirectly correlates the Darrington Phyllite with the Galice Formation. MacDonald et al. (2008a), based on similar ages and geochemical affinities, correlated the Late Jurassic Esmeralda Peaks unit of the Ingalls ophiolite complex with the Late Jurassic Josephine ophiolite, Klamath Mountains, California–Oregon. The Peshastin and Galice Formations conformably overlie these ophiolites, further supporting this correlation. Brown and Blake (1987) also correlated the Darrington Phyllite with rocks of northern California–southern Oregon. A detailed study of the regional relationships of the Shuksan Greenschist needs to be conducted to understand how it correlates to other Cordilleran Mesozoic terranes. The contact between Darrington Phyllite and Shuksan Greenschist protoliths is interpreted to have been conformable (Haugerud et al., 1981). If the correlation of the Darrington Phyllite and Peshastin Formation is correct, then the Esmeralda Peaks and Iron Mountain units of the Ingalls ophiolite complex may be possible lower-grade correlatives to the Shuksan Greenschist. Assuming that the lithologic and chronologic correlations we propose are correct, and that these fundamental similarities reflect original spatial proximity of the geologic units and/or terranes, then we can infer translation of the originally contiguous

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units. Numerous researchers have suggested the terranes of the central Cascades have been dextrally displaced along the Straight Creek–Fraser River fault ~100 km after their thrust fault emplacement (e.g., Vance and Miller, 1981; Tabor et al., 1984; MacDonald and Dragovich, 2015). Prior to early Tertiary displacement on the Straight Creek–Fraser River fault, the Hicks Butte complex may have translated ~330 km from the Blue Mountains to the Washington Cascades. The Easton Metamorphic Suite and Ingalls ophiolite complex may have translated between 550 and 600 km from northern California–southern Oregon to the Washington Cascades. The faults that displaced these and other Cordilleran terranes still need to be clarified. ROAD LOG Miller and Frost (1977), Johnson and Miller (1987), and Harper et al. (2003) provide field guides that overlap the Ingalls ophiolite complex covered by this review. Please refer to them for detailed information on the Ingalls ophiolite complex. Latitude and longitude coordinates in World Geodetic System 1984 (WGS84) datum are presented for each field stop in decimal degrees (hddd.dddd°). Google Earth and other web mapping applications use the WGS84 datum and can easily plot decimal degrees. Many of the stops will be on National Forest Service Roads (NFSR), which are classified by the U.S. Forest Service as “improved gravel roads.” They are passable for any type of vehicle; however, vehicles with high clearance are recommended. Begin by driving to the city of Cle Elum, Kittitas County, Washington. From Cle Elum, drive south through the town of South Cle Elum along S. Cle Elum Way, turning right (west) onto Madison Avenue, then left (south) onto 6th Street. 6th Street will become Westside Road. Drive along Westside Road until you come to Woods and Steele Road (~3.8 miles from the intersection of Madison Avenue and 6th Street in South Cle Elum). Turn left (south) onto Woods and Steele Road and set odometer at zero. ■■ Day

1 On Day 1, we observe the Easton Metamorphic Suite in the Hicks Butte and Kachess Lake inliers (Figs. 2 and 3), Hicks Butte complex and tectonic zone in the Hicks Butte inlier (Fig. 3), and North Peak unit in the Kachess Lake inlier (Fig. 2). We observe both greenschist (Hicks Butte inlier; Fig. 3) and blueschist (Kachess Lake inlier; Fig. 2) of the Shuksan Greenschist. Mileage Description 0.0 Corner of Westside Road and Woods and Steele Road. Proceed southwest on Woods and Steele Road. 0.7 Turn left off Woods and Steele Road onto NFSR (National Forest Service Road) 4510. This initial part of NFSR 4510 is paved and a local developer has renamed part of it All Season Drive. 1.3 NFSR 4510 forks with Snow Ridge Drive. Stay to the right to continue on NFSR 4510. Immediately after

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4.5

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MacDonald et al. the fork, NFSR 4510 becomes a gravel road. This fork in the road can be confusing, due to the cul-desac-like nature of the termination of the pavement on NFSR 4510. Stop 1-1: Shuksan Greenschist (N47.1445°, W121.0332°). We are now in the Hicks Butte inlier (Fig. 3). An outcrop of Shuksan Greenschist (Fig. 3) is located along the left. Here, epidote- and chloritebearing greenschist is strongly foliated, expressed by 5–10-mm-thick compositional layering, and lineated. Foliation strikes NW-SE and dips moderately to the SW. Stop 1-2: Tectonic zone (N47.1402°, W121.0398°). Located here are strongly foliated and lineated rocks of the tectonic zone (Fig. 3). Coarse- to fine-grained tonalitic to gabbroic gneiss, amphibolite, and rare mylonite can be seen in outcrop. Strong planar foliation here strikes NW-SE and dips moderately to the SW. Lineation trend (NW) here is consistent with that in the Shuksan Greenschist at the previous stop. Geochemistry and age data suggest the tectonic zone is a higher-grade equivalent of the Hicks Butte complex. See “Discussion.” Stop 1-3: Tectonic zone (N47.1394°, W121.0360°). Located here are medium-grained porphyroclastic tonalitic to gabbroic gneiss and amphibolite. Minor mylonite also occurs. Again, foliation strikes NW-SE and dips moderately to the SW. Mineral and stretching lineations trend NW and have a shallow plunge. Open rootless folds and shear structures (e.g., rotated mineral grains) are common. A very large (>1.5 m) boudin of amphibolite occurs here. Stop 1-4: Hicks Butte complex (N47.1357°, W121.0424°). Park and walk south up the hill. Located here are variably deformed hornblende tonalite, hornblende quartz diorite, hornblende diorite, hornblende gabbro, dacite, and rare rhyolite and hornblendite of the Hicks Butte complex (Fig. 3). High Sr/Y dacite at this site provided the ca. 144 Ma U-Pb age (Figs. 3 and 8B), whereas the other lithologies, based on low Sr/Y geochemistry, are ca. 153–150 Ma. Numerous amphibole (Fig. 6C) and plagioclase (Fig. 10) analyses from the low Sr/Y and high Sr/Y phases are from this stop. Low Sr/Y hornblende gabbro and rare hornblendite occur here as dikes cutting low Sr/Y hornblende tonalite, hornblende quartz diorite, and hornblende diorite. Dikes range in thickness from 1 cm to ~30 cm. High Sr/Y dacite and rhyolite occur as 30-cm-thick dikes cutting all other lithologies. Enclaves occur in the granitoid lithologies, suggesting magma mixing may have occurred. Dikes are commonly folded, faulted (extensional), or stretched into boudinage. Weak foliation and lineation here have the same general direction as at previous stops.

7.8 Turnaround: The Cretaceous Mount Stuart batholith, Jurassic Ingalls ophiolite complex (Day 2), and younger Eocene sedimentary and volcanic rocks form the vistas of the Wenatchee Mountains and Stuart Range to the north of the turnaround. 14.2 After the turnaround, follow NFSR 4510 back to Woods and Steele Road and turn right onto Woods and Steele Road. 14.9 Follow Woods and Steele Road back to the north, and turn left onto Westside Road. 18.0 Turn left onto Fowler Creek Road. 18.8 Stop 1-5: Darrington Phyllite (N47.1704°, W121.0714°). Darrington Phyllite occurs along Fowler Creek Road. Park at the intersection of Fowler Creek Road and NFSR 4511 and walk across the street to the outcrop (Fig. 3). Tight, upright F2 folds deform the S1 foliation here and are associated with a steep, spaced S2 foliation. The L2 crenulation lineation and S1 foliation have the same NW-SE striking, dipping moderately to the SW fabric as at the other stops in the Hicks Butte inlier. Good outcrops of well-foliated and folded, darkcolored Darrington Phyllite with quartz segregations occur along Fowler Creek Road as you go to the northwest toward Westside Road. Caution: Do not visit the outcrop on the north side of Fowler Creek Road and Westside Road. This is a hairpin turn and can be very dangerous. The same structures at this dangerous outcrop can be seen on Fowler Creek Road. Turnaround: Turn around at Stop 1-5, and drive northeast on Fowler Creek Road back toward its intersection with Westside Road. Mileage Description 19.6 Turn left onto Westside Road. 21.3 Turn right onto Golf Course Road. 21.5 Turn right to merge onto Interstate 90 East (I-90 E). 23.5 Take exit 80 toward Salmon La Sac/Roslyn. 25.6 Turn left onto Bullfrog Road. 27.8 At the second traffic circle, continue straight onto WA-903 North. WA-903 North will become Salmon La Sac Road. 39.5 Turn left off WA-903 North, Salmon La Sac Road, onto gravel road NFSR 4308. 44.9 Stop 1-6: North Peak unit (N47.3461°, W121.1744°). We are now in the Kachess Lake inlier (Fig. 2). Here, fine-grained, dark- to light-green meta-volcaniclastic rocks of the North Peak unit (Ashleman, 1979) outcrop along NFSR 4308 (Fig. 2). Ashleman (1979) reported lawsonite from the North Peak unit. Tabor and Haugerud (2016) suggested the North Peak unit was in the Excelsior Nappe. The Excelsior Nappe in



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the northwest Cascades (Fig. 1) includes the Chilliwack Group, which has lithologies and metamorphic mineral assemblages similar to those of the North Peak unit (see Tabor and Haugerud, 2016, and references therein). As we drive southwest from this stop, we cross the folded thrust fault contact between the Easton Metamorphic Suite and the North Peak unit (Fig. 2). The Darrington Phyllite, just southwest of this stop, is remarkably similar, both lithologically and structurally, to the rock at Stop 1-5. 51.3  Stop 1-7: Shuksan blueschist (N47.3346°, W121.1832°). This stop is blueschist in the Shuksan Greenschist (Fig. 2). The epidote blueschist here consists of mm- to cm-thick compositional layers. Open to tight folds deform the generally NW-SE–striking foliation, which dips moderately to steeply to the SW. Crenulation lineations are observable on the S1 foliation. This is the locality Maruszczak et al. (2016) sampled for Na-amphibole and epidote geochemistry (Figs. 6A, 6B, and 7). See “Shuksan Greenschist in the Central Cascades” section above. ■■ Day

2 On Day 2, we observe the sedimentary rocks of the Jurassic Peshastin Formation, serpentinite, and the Iron Mountain unit from the Ingalls ophiolite complex (Figs. 4 and 5). We work along U.S. Highway 97, which can have heavy traffic. Please use caution at all times when working along this highway. Drive north to Blewett Pass on U.S. Highway 97 (BP on Fig. 1) and reset odometer. Continue north after resetting odometer. Mileage Description 8.7  Stop 2-1: Ingalls ophiolite complex (N47.4066°, W120.6577°). This is Stop 2 from Harper et al. (2003). Here, a spectacular outcrop occurs along the west side of Highway 97 (Fig. 5). The outcrop can be observed from the highway, however, please be very careful if you decide to cross Peshastin Creek to better see it. The southern end of the outcrop is massive and sheared serpentinite. Foliation of the serpentinite is anastomosing, but generally strikes W-E and dips steeply to the south. Milliken and MacDonald (2013) analyzed serpentinite from this stop for both serpentine minerals and whole rock geochemistry. Major element ratios are consistent with a harzburgite protolith that plots in the abyssal peridotite field on Figure 12. Serpentinite is steeply faulted against the Early Jurassic Iron Mountain unit (MacDonald et al., 2008b). At this stop, the Iron Mountain unit consists of, from south to north, limestone, hyaloclastite, chert, breccia, and pillow basalt. The sedimentary rocks appear to be folded into the steep fault. Basalt from the Iron Mountain unit has within-plate geochemical affinities (MacDonald et al., 2008b).

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10.1  Stop 2-2: Ingalls serpentinite (N47.4218°, W120.6588°). We park at the site of the gold-mining town of Blewett (Fig. 5). Walk south to observe black to light-green, massive, faulted, and sheared serpentinite along the east side of Highway 97. The anastomosing foliation of the serpentinite generally dips steeply to the north. Serpentinization here is pervasive; however, some samples have well-preserved primary clinopyroxene and orthopyroxene. A few protoliths at this locality may have been olivine pyroxenite instead of peridotite. Johnson and Miller (1987) report rodingitized mafic rocks from this location. Rodingitization is Ca enrichment of non-ultramafic rocks during serpentinization. A large, north-dipping Eocene Teanaway dike cuts the serpentinite at the northern end of the outcrop. 10.7  Stop 2-3: Peshastin Formation (N47.4313° W120.6561°). Park at the old dead-end road to the right (east) and walk north to the outcrop on the east side of the highway (Fig. 5). This outcrop exposes siltstone, argillite, and rare chert. Bedding here strikes W-E and dips moderately to the north; some thin sedimentary beds resemble Bouma turbidite sequences. Miller et al. (2003) collected sandstone for detrital zircon U-Pb (ca. 232 and 152 Ma) at this site. The coarse-grained lithic graywacke contains abundant igneous lithic, argillite, and chert clasts. It also contains monocrystalline quartz grains, some of which are bipyramidal and contain altered fluid inclusions (MacDonald and Dragovich, 2015), in accord with the suggestion by Harper et al. (2003) that it is volcanically derived (Fig. 11). This locality is approximately the same as stop 7 of Harper et al. (2003). 13.1  Stop 2-4: Peshastin Formation (N47.4628°, W120.6602°). Park at the turnoff on the east side and walk back south to the second, larger outcrop of deformed sedimentary rocks along the east side of the highway (Fig. 5). Here, thin beds of predominantly argillite strike E-W and dip moderately to the south. A SE-NW cleavage, which dips to the SW, overprints the bedding. Massive siltstones and sandstones occur toward the south of the outcrop. The coarse-grained sandstone includes rip-up clasts of argillite, monocrystalline quartz, plagioclase, and chert lithic clasts in a finer matrix. The rock has been contact metamorphosed, as evidenced by overprinted web textures, by the intrusion of the Mount Stuart batholith and is cut by numerous faults. ACKNOWLEDGMENTS We thank Kirsten Sauer and Adam Schoonmaker for their thoughtful and helpful reviews, which greatly improved this chapter. The mineral geochemistry data discussed in this

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chapter were the result of course-based undergraduate research projects conducted by Florida Gulf Coast University students and funded by the NSF TUES grant 1323354 to MacDonald. Additional funding by the Florida Gulf Coast University Office of Research and Graduate Studies to MacDonald and Whitaker Center for STEM Education to Stingu are also acknowledged. We thank the GSA staff for their excellent editorial work, as well as Ralph Haugerud and Harvey Kelsey for being the GSA 2017 Annual Meeting field trip co-chairs. Haugerud also provided an excellent review of this chapter, which greatly improved it. The hardworking staff at the Florida Center for Electron Microscopy, Florida International University, greatly assisted with all EPMA work. APPENDIX

Plagioclase geochemical data from the Hicks Butte complex is available through EarthChem at http://get.iedadata.org/ doi/100599. Amphibole geochemical data from the Hicks Butte complex is available through EarthChem at http://get.iedadata.org/ doi/100600. Epidote geochemical data from the Shuksan blueschist, Kachess Lake inlier, is available through EarthChem at http://get.iedadata.org/doi/100634. Amphibole geochemical data from the Shuksan blueschist, Kachess Lake inlier, is available through EarthChem at http:// get.iedadata.org/doi/100635. REFERENCES CITED Albertz, M., Paterson, S.R., and Okaya, D., 2005, Fast strain rates during pluton emplacement: Magmatically folded leucocratic dikes in aureoles of the Mount Stuart Batholith, Washington, and the Tuolumne Intrusive Suite, California: Geological Society of America Bulletin, v. 117, p. 450–465, doi:10.1130/B25444.1. Anderson, J.L., and Smith, D.R., 1995, The effects of temperature and f O2 on the Al-in-hornblende barometer: The American Mineralogist, v. 80, p. 549–559, doi:10.2138/am-1995-5-614. Arai, S., Okamura, H., Kadoshima, K., Tanaka, C., Suzuki, K., and Ishimaru, S., 2010, Chemical characteristics of chromian spinel in plutonic rocks: Implications for deep magma processes and discrimination of tectonic setting: The Island Arc, v. 20, p. 125–137, doi:10.1111/j.1440 -1738.2010.00747.x. Armstrong, R.L., 1980, Geochronometry of the Shuksan Metamorphic Suite, North Cascades, Washington: Geological Society of America Abstracts with Programs, v. 12, no. 3, p. 94. Armstrong, R.L., and Misch, P., 1987, Rb-Sr and K-Ar dating of mid-Mesozoic blueschist and Late Paleozoic albite-epidote amphibolite and blueschist metamorphism in the North Cascades, Washington and British Columbia, and Sr isotope fingerprinting of eugeosynclinal rock assemblages: Washington Division of Geology and Earth Resources Bulletin, v. 77, p. 85–105. Ashleman, J.C., 1979, The Geology of the western part of the Kachess Lake quadrangle, Washington [M.S. thesis]: Seattle, University of Washington, 88 p. Beck, M.E., Jr., Burmester, R.F., and Schoonover, R., 1981, Paleomagnetism and tectonics of the Cretaceous Mt. Stuart batholith of Washington: Translation or tilt?: Earth and Planetary Science Letters, v. 56, p. 336– 342, doi:10.1016/0012-821X(81)90138-2. Bracciali, L., Marroni, M., Pandolfi, L., and Rocchi, S., 2007, Geochemistry and petrography of Western Tethys Cretaceous sedimentary covers (Cor-

sica and Northern Apennines): From source areas to configuration of margins, in Arribas, J., Critelli, S., and Johnson, M.J., eds., Sedimentary Provenance and Petrogenesis: Perspectives from Petrography and Geochemistry: Geological Society of America Special Paper 420, p. 73–93, doi:10.1130/2006.2420(06). Brandon, M.T., Cowan, D.S., and Vance, J.A., 1988, The Late Cretaceous San Juan Thrust System, San Juan Islands, Washington: A Case History of Terrane Accretion in the Western Cordillera: Geological Society of America Special Paper 221, 81 p., doi:10.1130/SPE221-p1. Brown, E.H., 1974, Comparison of the mineralogy and phase relations of blueschists from the North Cascades, Washington, and greenschists from Otago, New Zealand: Geological Society of America Bulletin, v. 85, p. 333–344, doi:10.1130/0016-7606(1974)852.0.CO;2. Brown, E.H., 1977, The crossite content of Ca-amphibole as a guide to pressure of metamorphism: Journal of Petrology, v. 18, p. 53–72, doi:10.1093 /petrology/18.1.53. Brown, E.H., 1986, Geology of the Shuksan Suite, North Cascades, Washington, U.S.A., in Evans, B.W., and Brown, E.H., eds., Blueschists and Eclogites: Geological Society of America Memoir 164, p. 143–154, doi:10.1130/MEM164-p143. Brown, E.H., 2012, Obducted nappe sequences in the San Juan Island–northwest Cascades thrust system, Washington and British Columbia: Canadian Journal of Earth Sciences, v. 49, p. 796–817, doi:10.1139/e2012-026. Brown, E.H., and Blake, M.C., Jr., 1987, Correlation of early Cretaceous blueschists in Washington, Oregon and northern California: Tectonics, v. 6, p. 795–806, doi:10.1029/TC006i006p00795. Brown, E.H., and Gehrels, G.E., 2007, Detrital zircon constraints on terrane ages and affinities and timing of orogenic events in the San Juan Islands and North Cascades, Washington: Canadian Journal of Earth Sciences, v. 44, p. 1375–1396, doi:10.1139/E07-040. Brown, E.H., Wilson, D.L., Armstrong, R.L., and Harakal, J.E., 1982, Petrologic, structural, and age relations of serpentinite, amphibolite, and blueschist in the Shuksan Suite of the Iron Mountain–Gee Point area, North Cascades, Washington: Geological Society of America Bulletin, v. 93, p. 1087–1098, doi:10.1130/0016-7606(1982)932.0.CO;2. Burkhard, D.J.M., 1993, Accessory chromium spinels; their coexistence and alteration in serpentinites: Geochimica et Cosmochimica Acta, v. 57, p. 1297–1306, doi:10.1016/0016-7037(93)90066-6. Cheney, E.S., and Hayman, N.W., 2009, The Chiwaukum structural low: Cenozoic shortening of the central Cascade Range, Washington State, USA: Geological Society of America Bulletin, v. 121, p. 1135–1153, doi:10.1130/B26446.1. Cowan, D.S., Brandon, M.T., and Garver, J.I., 1997, Geologic tests of hypotheses for large coastwise displacements: A critique illustrated by the Baja British Columbia controversy: American Journal of Science, v. 297, p. 117–173, doi:10.2475/ajs.297.2.117. Davis, P., and Lindmark, M., 2015, Evidence for lower crustal syntectonic forearc emplacement of the Hicks Butte pluton breaking the correlation of HP-LT rocks across the Cascades of Washington State: Geological Society of America Abstracts with Programs, v. 47, no. 7, p. 766. Defant, M.J., and Drummond, M.S., 1990, Derivation of some modern arc magmas by melting of young subducted lithosphere: Nature, v. 347, p. 662– 665, doi:10.1038/347662a0. Dick, H.J.B., and Bullen, T., 1984, Chromian spinel as a petrogenetic indicator in abyssal and alpine-type peridotites and spatially associated lavas: Contributions to Mineralogy and Petrology, v. 86, p. 54–76, doi:10.1007 /BF00373711. Dragovich, J.D., Norman, D.K., Grisamer, C.L., Logan, R.L., and Anderson, G., 1998, Geologic Map and Interpreted Geologic History of the Bow and Alger 7.5 Minute Quadrangles, Western Skagit County, Washington: Washington Division of Geology and Earth Resources Open File Report 98-5, 80 p., 3 plates. Dragovich, J.D., Norman, D.K., Lapen, T.J., and Anderson, G., 1999, Geologic Map of the Sedro-Woolley North and Lyman 7.5-Minute Quadrangles, Western Skagit County, Washington: Washington Division of Geology and Earth Resources Open File Report 99-3, 37 p., 4 plates. Dragovich, J.D., Norman, D.K., and Anderson, G., 2000, Interpreted Geologic History of the Sedro-Woolley North and Lyman 7.5-Minute Quadrangles, Western Skagit County, Washington: Washington Division of Geology and Earth Resources Open File Report 2000-1, 71 p., 1 plate. Dungan, M.A., Vance, J.A., and Blanchard, D.P., 1983, Geochemistry of the Shuksan greenschists and blueschists, North Cascades, Washington;



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Manuscript Accepted by the Society 13 July 2017

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