Journal of Volcanology and Geothermal Research 185 (2009) 251–275

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Journal of Volcanology and Geothermal Research

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Evolution of an englacial volcanic ridge: Pillow Ridge tindar, Mount Edziza volcanic complex, NCVP, British Columbia, Canada

Benjamin R. Edwards a,⁎, Ian P. Skilling b, Barry Cameron c, Courtney Haynes a, Alex Lloyd a, Jefferson H.D. Hungerford b a Department of Geology, Dickinson College, 5 N. Orange Street, Carlisle, PA, 17013, USA b Department of Geology and Planetary Science, 200 SRCC Building, University of Pittsburgh, Pittsburgh, PA, 15260, USA c Department of Geology, University of Wisconsin-Milkwaukee, WI, USA article info abstract

Article history: Glaciovolcanic deposits are critical for documenting the presence and thickness of terrestrial ice-sheets, and Received 24 April 2008 for testing hypotheses about inferred terrestrial ice volumes based on the marine record. Deposits formed by Accepted 9 November 2008 the coincidence of volcanism and ice at the Mount Edziza volcanic complex (MEVC) in northern British Available online 24 November 2008 Columbia, Canada, preserve an important record for documenting local and possibly regional ice dynamics. Pillow Ridge, located at the northwestern end of the MEVC, formed by ice-confined, fissure-fed eruptions. It Keywords: comprises predominantly pillow lavas and volcanic breccias of alkaline basalt composition, with subordinate tindar fi ∼ ∼ glaciovolcanism ner-grained volcaniclastic deposits and dykes. The ridge is presently 4 km long, 1000 m in maximum volcano–ice interaction width, and ∼600 m high. Fifteen syn- and post-eruptive lithofacies are recognized in excellent exposures Mount Edziza volcanic complex along the glacially dissected western side of the ridge. We recognize five lithofacies associations: (1) poorly pillow lava sorted tuff breccia and dykes, (2) proximal pillow lava, dykes and tuff breccia, (3) distal pillow lava, poorly sorted conglomerate and well-sorted volcanic sandstone, (4) interbedded tuff, lapilli tuff, and tuff breccia units, and (5) heterolithic volcanogenic conglomerate and sandstone. Given the abundance of pillow lavas and the lack of surrounding topographic barriers capable of impounding water, we agree with Souther [Souther, J.G., 1992. The late Cenozoic Mount Edziza volcanic complex. Geol. Soc. Can. Mem., vol. 420. 320 pp] that the bulk of the edifice formed while confined by ice, but have found evidence for a more complex and variable eruption history than that which he proposed. Preliminary estimates of water-ice depths derived

from FTIR analyses of H2O give ranges of 300 to 680 m assuming 0 ppm CO2, and 857 to 1297 m assuming

25 ppm CO2. Variations in depth estimates among samples may indicate that water/ice depths changed during the evolution of the ridge, which is consistent with our interpretations for the origins of different lithofacies associations. Given that the age of the units are likely to be ca. 0.9 Ma [Souther, J.G., 1992. The late Cenozoic Mount Edziza volcanic complex. Geol. Soc. Can. Mem., vol. 420. 320 pp], Pillow Ridge may be the best documentation of a regional high stand of the Cordilleran Ice Sheet (CIS) in the middle Pleistocene, and an excellent example of the lithofacies and stratigraphic complexities produced by variations in water levels during a prolonged glaciovolcanic eruption. © 2008 Elsevier B.V. All rights reserved.

1. Introduction retreats for the North American Ice Sheet (NAIS) are well constrained from studies of end-moraines, geomorphology and isostatic effects One of the outstanding problems in studies of Earth's paleo- (Kutzbach, 1987; James et al., 2000; Clague and James, 2002; Fulton climate is the comparison of very detailed records of global et al., 2003), but the terrestrial record of pre-Last Glacial Maximum temperature from Pleistocene marine and ice core proxies to the (LGM; e.g., Illinoian, Kansan and Nebraskan) Pleistocene ice is much relatively sparse records on the extents and thicknesses of terrestrial less detailed, since much of the evidence was destroyed during the ice-sheets, particularly in the Northern Hemisphere. Data from marine LGM (Fulton, 1992; Jackson et al., 1996; Barendegt and Irving, 1998). cores and lacustrine sediments suggest multiple episodes of glacia- The products of volcano–ice interactions can provide critical tions during the Pleistocene (Raymo, 1992; Benson et al., 1998; constraints for the presence and characteristics of pre-LGM ice, as Marshall et al., 2002). The pattern and timing of the most recent they can be more resistant to erosion than other types of glaciogenic deposits (e.g. till). Two volcanic landforms, both capable of surviving ⁎ Corresponding author. Tel.: +1 717 254 8934; fax: +1 717 245 1971. multiple episodes of ice burial and erosion, are broadly recognized as E-mail address: [email protected] (B.R. Edwards). forming only in ice-confined environments: tuyas (Mathews, 1947)

0377-0273/$ – see front matter © 2008 Elsevier B.V. All rights reserved. doi:10.1016/j.jvolgeores.2008.11.015 252 B.R. Edwards et al. / Journal of Volcanology and Geothermal Research 185 (2009) 251–275 and tindars (Jones, 1969). Tuyas are flat-topped volcanoes that tephra, and capped by subaerial lava-fed deltas (e.g., Mathews, 1947; comprise a distinctive subaqueous to emergent sequence of basal Jones,1969; Skilling, 2002; Smellie, 2006). Tindars are distinctly linear lavas and volcanic breccias, overlain by Surtseyan phreatomagmatic volcanic landforms that comprise lithofacies similar to those found at tuyas except that they generally lack extensive lava-fed deltas (Jones, 1969). They are thought to form during ice-confined, fissure-fed eruptions. Jones (1969) first proposed that the word ‘tindar’, which is an Icelandic word meaning ‘row of peaks’, be used as a general term to describe linear volcanic landforms in west central Iceland formed by eruptions beneath ice. Although strictly speaking tindar is the plural form of tindur, we follow Jones (1969) usage of ‘tindar’ as singular and ‘tindars’ as plural. Jakobsson and Guðmundsson (2008) have recently suggested that tindars be distinguished from tuyas by having at least a 2:1 length to width ratio. Linear glaciovolcanic landforms have also been referred to as ‘hyaloclastite ridges’ (e.g., Schopka et al., 2006) and ‘pillow ridges’ (e.g., Höskuldsson et al., 2006). Neither of these terms is satisfactory as a general term because the majority of the fragmental material found at tindars is not hyaloclastite sensu stricto (i.e. vitric fragments formed by quench fragmentation and mechanical ‘spalling’; Rittman, 1962; Honnorez and Kirst, 1975; McPhie et al., 1993), and the volumetric proportions of pillow lavas at tindars can vary from nearly 100% to almost 0%. The word ‘tindar’ has also been used more generally to denote the explosion-dominated stage of iced-confined basaltic eruptions regardless of vent geometry (Smellie and Skilling, 1994; Smellie, 2000). Published studies of tindars indicate that their lithostratigraphy is highly variable between two extremes, one dominated by extensive magmatic fragmentation, leading to a predominance of fragmental lithofacies (Guðmundsson et al., 2002a,b; Schopka et al., 2006) and the other dominated by pillow lavas (Höskuldsson et al., 2006). It is very likely that the variations between the two end members are controlled by variations in ice thickness, sub-ice hydrology at the time of the eruptions, eruption rates, eruption duration and volume, and possibly pre-eruption volatile contents. Tindars with significant proportions of fragmental and coherent lithofacies have been described in Iceland (Kalfstindar; Jones, 1969, 1970) and in northern British Columbia, Canada (Pillow Ridge; Souther, 1992). This paper focuses on the physical evolution of Pillow Ridge, a tindar at the northwestern end of the Mount Edziza volcanic complex (Fig. 1). Our goals are to: (1) give detailed descriptions of the lithofacies and their associations, (2) present viable interpretations for their origins, (3) evaluate Souther's (1992) model for the evolution of the ridge, and (4) explore the implications of tindar formation for local and regional ice/meltwater dynamics at ∼1 Ma in northern British Columbia.

2. Geological setting

Pillow Ridge is part of the northern Cordilleran volcanic province (NCVP), which comprises isolated cones, lava flow remnants, and a few larger volcanic complexes located in northwestern British Columbia (BC), the western Yukon Territory, and east-central Alaska (Fig. 1A; Edwards and Russell, 1999, 2000). The NCVP formed mainly during the Neogene and Holocene in response to extensional tectonic stresses spread across the Canadian Cordillera (Souther, 1992; Edwards and Russell, 1999). Much of the eruption activity in the NCVP occurred during the past 2 Ma (cf. Edwards and Russell, 2000), when western Canada was periodically inundated by the Cordilleran

Fig. 1. Maps showing the location of the Mount Edziza volcanic complex (MEVC) and Pillow Ridge. A) Map of the Canadian Cordillera showing the location of the northern Cordilleran volcanic province (NCVP) and the MEVC. Hillshade is derived from 30 m Digital Elevation Models (DEM) available from the National Topographic Atlas. B) Map of the MEVC showing underlying topography, general outline of the distribution of formations within the MEVC (light grey), and locations of remnant ice still extant (after Souther, 1992, Fig. 4). The general locations of Pillow Ridge (dashed oval), the peak of Mount Edziza (northern black star), and the approximate, hypothesized location of ‘Ice Volcano’ (southern black star) are shown for reference. B.R. Edwards et al. / Journal of Volcanology and Geothermal Research 185 (2009) 251–275 253

Ice Sheet (CIS; Ryder and Maynard, 1991), which formed between the 3. Pillow Ridge morphology coastal mountains of western British Columbia and the Rocky Mountains of Alberta (Fulton, 1984; Clague, 1989; Ryder and Maynard, Pillow Ridge is an asymmetric, sigmoidal ridge whose crest 1991; Stumpf et al., 2000). Due to the coincidence in time and space of extends north-northwest from the northwestern edge of the Mount NCVP activity and the CIS, many glaciovolcanic deposits and landforms Edziza icecap (Figs. 1B, 2, 3). It has a length of at least 4 km, and a have been documented in northern British Columbia (Mathews, 1947; maximum width of ∼1 km. Elevations of the exposed base vary from Allen et al., 1982; Souther, 1992; Edwards et al., 1995; Moore et al., 1675 to 2130 m above sea level (ASL), and the highest part of the ridge 1995; Simpson, 1996; Bye et al., 2000; Hickson, 2000; Edwards and crest is ca. 2400 m ASL. The height from base to ridge crest is at Russell, 2002; Edwards et al., 2002; Harder and Russell, 2005; maximum 450 m, and the ridge tapers in height to the northwest Edwards et al., 2006). Of particular interest to paleoclimate studies where it dives beneath talus and lava flows from the Edziza Formation are two of the larger volcanic complexes in the NCVP, Hoodoo to the west and the Big Raven Formation to the east (Fig. 3; Souther, Mountain (Edwards et al., 1999; Edwards et al., 2002; Edwards and 1992). Its present day volume is ∼1.8 km3 and its deposits cover Russell, 2002) and the Mount Edziza volcanic complex (Souther et al., ∼3km2. However, it is not certain how well these figures represent 1984; Souther and Hickson, 1984; Souther, 1992), because each has the actual eruption volumes because the base of the ridge is not preserved evidence of volcanic activity spanning tens of thousands to exposed unequivocally, and deposits from the ridge are buried millions of years. These long-lived centres are key locations for beneath overlying Edziza Formation lava flows to the northwest. establishing the presence and extent of former ice in western North Clearly some of the volume of the original deposits has been removed America during the Pleistocene. by subsequent erosion. The Mount Edziza volcanic complex (Fig. 1B), which comprises For convenience of description, several distinctive topographic predominantly alkaline basalt and trachyte, has had sporadic eruption features have been given informal names (Fig. 2B), and the ridge has activity since about 10 Ma (Souther et al., 1984; Souther, 1992). been divided into generally west-facing and generally east-facing slopes. Although Souther (1992) documented many examples of volcano–ice Informal geographic names used in the text will always be italicized to interaction at the MEVC, Pillow Ridge is one of the most distinctive distinguish them from formally recognized geographic designations. glaciovolcanic landforms at the complex. It formed along the north- The west-facing flank of the ridge is unvegetated and dominated by cliffs western flank of the remnants of “Ice Volcano”, a large central volcano and talus fans (Fig. 2A,B). It has obviously been eroded by past ice that erupted a broad range of lava compositions in subaerial and sub- advances (Fig. 2A,C); a tongue of ice fed from the main icecap still marks ice conditions around 1 Ma (Fig. 1B; Souther, 1992). Fission track ages the southwestern margin of the ridge (Fig. 2C,E). Erosion by thicker on apatite crystals from granitic xenoliths within Pillow Ridge lava precursors of the present day ice tongue has cut steep slopes along this give resetting ages of 0.9+/−0.3 Ma and 0.8+/−0.25 Ma, which margin, creating an asymmetric form to the ridge cross-section, with broadly constrain the formation of Pillow Ridge to the Mid Pleistocene much steeper slopes on the western side than on the eastern (Fig. 2C). A (Souther, 1992). narrow canyon occupied by a small stream marks the extreme

Fig. 2. Morphology of Pillow Ridge. A) View looking from the west at the western side of the ridge. B) Sketch of (A) showing approximate contacts between major units. The northern half of the ridge, from North Peak to the north, is underlain predominantly by pillow lava. Central Ridge comprises mostly tuff breccia and dykes. South Peak, the southern tip, comprises tuff breccia overlain by pillow lava. C) View looking to the southeast at the northern end of the ridge. Surface sloping on the northeastern side of the ridge appears to have undergone less glacial modification than the western side. D) View looking northwest along the central crest of the ridge. The symmetrical part of the ridgeline, in the middle of the image, is underlain by a dyke intruded parallel to the ridge crest. E) Clipped section of BC Airphoto BC5203-10 showing the shape of the entire ridge, outlined by black dotted line. Approximate view perspectives for images A, C, D are shown for reference, with ‘V’ opening in the viewing direction. 254 ..Ewrse l ora fVlaooyadGohra eerh15(09 251 (2009) 185 Research Geothermal and Volcanology of Journal / al. et Edwards B.R. – 275

Fig. 3. Hillshade and geology maps of Pillow Ridge. A) A 1:20 000 shaded relief map derived from 30 m Digital Elevation Model (DEM), showing station locations and illustrating the topography of Pillow Ridge. The maximum surficial extent of Pillow Ridge deposits is outlined for reference. Locations of pillows from which samples were collected for electron microprobe and H2O analyses are shown for reference (e.g., E06BE40). The extent of Pillow Canyon is also shown (white dashed oval). B) Detailed geological map of Pillow Ridge with boundaries modified from Souther (1992). Orientations of ovoids in pillow lava units are approximately parallel to measured orientations of pillow lobes in the field. Dyke locations are shown accurately, but dyke dimensions (length and width) are approximate. Dyke orientations and dimensions are also given in Table 2. B.R. Edwards et al. / Journal of Volcanology and Geothermal Research 185 (2009) 251–275 255 northwestern end of the steep southwestern side; the location of the (Fig. 3). To the north of North Peak,andthesouthofSouth Peak,unitsof canyon coincides with the contact between younger, Edziza Formation pillow lava are relatively broad and so the ridge crest is also broader. The trachyte lava flows to the west and the main mass of Pillow Ridge to the ridge-crest is inferred to be in a very similar position to that formed by east (Fig. 3). Within this narrow canyon, informally referred to as Pil- the last eruptions of pillow lavas at Pillow Ridge (Souther, 1992; Fig. 3), low Canyon, are excellent exposures of several different lithofacies, because pillows to the east of the ridge crest generally dip eastward, showing a cross-section through the northern end of the ridge. while those to the west generally dip westward. Central Ridge is very The east-facing flank appears to have experienced less erosion narrow at its apex (Fig. 2D), which is underlain by a dyke. than the western flank. Field evidence from pillow orientations is interpreted to indicate that the eastern flank approximates the 4. Descriptions of lithofacies depositional slope during pillow lava emplacement. Most of the crest of the ridge comprises coherent lithofacies: pillow Souther (1992) identified four lithologies at Pillow Ridge: (1) pillow lava at the southern and northern ends, and dykes in the central saddle lava, (2) dykes, (3) chaotic and bedded ‘aquagene’ tuff-breccia, and

Table 1 Codes, descriptions and interpretations of lithofacies at Pillow Ridge.

Facies Code Description Vol. %a Figs. Interpretation

Coherent Lpw1–5 Pillow lava: dark grey to black to orange grey; diameter of individual 40 4,5 Subaqueous extrusion of lava at moderate to low lobes ranges from 0.2 to N1 m; pillows typically have vitric rims up to eruption rates in ice-confined lake 1 cm thick all around exterior; vitric rims locally show corrugations and/or spreading cracks; vesicularity varies, locally with pillow cores having open cavities or highly vesiculated cores; concentric vesicle rings are present but pipe vesicles are uncommon; radial jointing commonly well-developed; locally accompanied by intertube breccia or ash; subdivided into 5 different stratigraphic units

LD1–2 Dykes: medium grey; massive and tabular; ranging in width from b1 6 Intrusions into TB1 over an indefinite time-span;

b1 to 3 m; locally bifurcating near top of ridge; commonly pervasively some emplaced while TB1 was still not fully lithified; fractured with joints parallel to dyke trend, with little to no well- variations in margins with elevations reflect vertical

developed fracture orthogonal to dyke trends; contacts with TB1 vary decreases in pressure and consolidation with elevation on ridge from straight, to curved, to pillowed; predominantly subvertical dip directions; locally with abundant gabbroic and granitic inclusions; subdivided into two units based on dominant orientations (see Table 2)

Volcaniclastic TB1 Tuff Breccia: grey to yellow; massive to crudely bedded, poorly sorted, 50 7A–D Near-vent subaqueous deposition of Surtseyan varying from matrix to clast-supported; comprising ash, lapillus tephra mostly by eruption-fed mass flows and and block size juvenile fragments; monomict with clasts of density-modified grain-flows predominantly angular, vesicular basalt; locally clasts are subrounded to fluidal; clasts commonly have wedge shapes, and locally vitric rims; matrix is various shades of yellow-orange

TB2 Tuff Breccia: grey; massive, clast supported, very poorly sorted, b5 7E–F Low energy, proximal talus avalanches of pillow monomict; comprising angular clasts of juvenile vesicular basalt; lava fragments, forming immediately downslope locally clasts have vitric rims and wedge shapes, or are large of leading edge of advancing pillow lobes. fragments of pillow lobes with well-developed jointing;

directly overlain locally by Lpw5, Lpw4, and T LT Lapilli Tuff: dark grey; massive, clast-supported, moderately ≪1 – Proximal airfall from subaerial surtseyan well-sorted, comprising equant, angular fragments of vesicular eruption column.

basalt; directly overlain by TB2 and T T Tuff: Orange to tan to dark grey; well-laminated with local oscillating ≪1 7G,H Deposits of low volume sediment gravity flows cross-stratification, comprising mainly dense to partly vesiculated platy and stream flows from site of pillow lava extrusion. fragments of sideromelane, with subordinate palagonite and crystals; Hyaloclastite sensu strictu. Cross-stratification well-sorted; varying in thickness from 10–14 cm; locally directly indicates current deposition.

overlain by LT, Lpw4 or interbedded with TB1 or S1 Sedimentary B Breccia: orange to grey; massive, well-sorted, clast-supported with b5 8A,B Lithified talus derived from weathering of jointed very angular, monomict clasts of vesicular basalt; clasts are blocky to pillow lava deposits. (differences between B and rectilinear in shape; most clasts covered with thin orange coating and TB1 need to be thought out carefully…) range from 4 to 9 cm in maximum dimension (very coarse pebble gravel to fine cobble gravel); apparent thickness on slope exposures N2m C Conglomerate: variegated; massive, poorly sorted, clast-supported ≪1 8C Fluvial channel fill. Clasts from a variety of different with subangular to subrounded, polymict clasts of volcanic and sources, including volcanic and plutonic. exposure plutonic rocks; clasts vary in size from b1cmtoN10 cm in maximum appears to be an erosional window hence dimension (medium pebble gravel to coarse cobble gravel); deposit geometry and dimensions unknown

S1 Sandstone: orange to tan-grey; massive, moderately sorted with b1 8D,E Material produced during Surtseyan eruptions and subrounded, oligomict grains; no apparent size grading or bedding; deposited by mass flows into an englacial lake or local cobble-sized, vesicular clasts with wedge shapes comprise ice-confined canyon from actively growing, upslope overall b5% by volume; minimum estimated deposit thickness is 5 m pillow lava emplacement; larger clasts are pillow fragments. Oligomict source.

S2 Sandstone: orange to tan-grey; massive, well sorted with subrounded, b1 8F,G Material deposited by debris avalanches into polymict grains, including black vitric grains and rounded, tan ice-confined canyon. Basaltic debris derived pumiceous grains; locally evidence of normal size-grading; minimum from Pillow Ridge from upslope pillow lava deposit thickness is 5–7 m; contains lenses of gravel to boulder-size lithofacies, other materials derived from melting of fragments that may define crude bedding; locally bedding is deformed confining ice on upslope side. Polymict source.

a Estimated relative volume percentage of Pillow Ridge total volume. 256 B.R. Edwards et al. / Journal of Volcanology and Geothermal Research 185 (2009) 251–275

(4) pillow breccia. Based on fieldwork in 2006 and 2007, we identified (3) sedimentary (rocks comprising a majority of clasts formed or fifteen different lithofacies at Pillow Ridge (Table 1). The lithofacies are reshaped during transport or depositional processes not directly related separated into three groups: (1) coherent (lava and subvolcanic to eruption processes). We use terminology advocated by White and intrusions), (2) volcaniclastic (rocks comprising fragments formed, Houghton (2006) to describe the volcaniclastic lithofacies and standard transported and deposited directly by eruption processes), and terminology to describe the sedimentary lithofacies.

Fig. 4. Field characteristics of pillow lava lithofacies (Lpw1–5) at Pillow Ridge. A) Photomosaic of the northwestern ridgeline as viewed from the west. B) Sketch of (A) showing approximate contacts and stratigraphic relationships between pillow lava (Lpw), tuff breccia (TB1–2), and dyke (LD2). C) Cliff-face on western side of ridge showing closely packed ‘entrail’ pillow lava. Uppermost pillow units have very little intra-lobe volcaniclastic material. The average diameter of individual pillows is 0.61 m (see Fig. 5). D) ‘Master’ pillow tube feeding three smaller, distributary pillows along the northern crest of ridge. E) Cross-sectional view of typical pillow from the western side of ridge, with radial jointing, oval vesicle bands, and a small, open, central cavity. B.R. Edwards et al. / Journal of Volcanology and Geothermal Research 185 (2009) 251–275 257

4.1. Coherent lithofacies units (Lpw4 and Lpw5) are distinctly darker in color and show less evidence of oxidation. They are also separated by up to 5 m of coarse tuff

4.1.1. Pillow lava (Lpw1–5) breccia, and have been intruded by dykes. Pillow lava, which comprises hundreds of individual pillows (Figs. 3B The pillows are commonly closely stacked, with little to no and 4), is the dominant lithofacies (Table 1) of the northern half of Pillow accompanying fragmental material between pillow tubes (Fig. 4C) Ridge, from the highest elevation along the ridge at North Peak except for local infillingsbetweenpillowsorcoatingsonpillowrimsof downslope to the northwestern end of the ridge at the mouth of Pil- orange ash-sized fragments (Lpw2 is exceptional in this respect). Locally, low Canyon. It also occurs in lesser abundance along the southern crest of larger feeder tubes branch into multiple smaller pillows (Fig. 4D). the ridge (Souther, 1992). Based on apparent breaks in slope and Patterns of vesiculation within individual pillows vary from well- weathering characteristics, we divided the pillow lava into five separate defined, concentric layers of vesicles (Fig. 4E) to disorganized highly units, designated Lpw1–5 (Fig. 4A,B). The lowest unit (Lpw1) is exposed vesicular masses filling the pillow core and grading to lower over a very limited area at the western base of the ridge along the bed of vesicularities at the pillow rim. Locally pillow cores have small voids a small stream. The pillows have sub-horizontal dips that appear to (Fig. 4E), but no systematic variations of vesicularity with height in a control the slope of the present streambed. The second unit (Lpw2)is specific pillow unit were observed. exposed locally downstream from Lpw1, and is distinct from the other Qualitative and quantitative field observations indicate several Lpw units because it is interbedded with coarse tuff breccia (TB2), and general trends in the pillow sequences. While the average measured appears to have some intrusive pillows. It is also intruded by vertical maximum cross-sectional dimension is 0.6 m, qualitative observations dykes (LD), which clearly cut through pillows, and horizontal, sill-like indicate a general trend of larger pillow dimensions in units Lpw1–3, bodies. The third unit (Lpw3) forms the base of a cliff-forming section of and smaller dimensions in with Lpw4–5 (Fig. 5A). Most pillows have the northwestern side of the ridge. It is orange colored in outcrop and is slightly ovoid shapes, with fewer showing deformation from overlying approximately 50 m thick. In this unit pillows locally have up to 50 vol.% pillows or molding around lower pillows. Quantitative measurements small gabbroic inclusions. The contact between Lpw3 and Lpw4 is locally of over 130 pillow lobe trend and plunge directions were measured for marked by a discontinuous, 3 m thick sequence lithofacies association, Lpw1–5, in order to identify possible centers of eruption for pillow which is described in detail below (Section 5.5). The upper two pillow units. The measurements show that the majority of the pillow lobes

Fig. 5. Field measurements of pillow dimensions and plunges at Pillow Ridge. A) Histogram showing the variation in pillow dimensions. Measurements were madetoreflect the largest dimension in view orthogonal to the length of the pillow where possible. Population shows one clear maximum in the 50–60 cm bin, and possibly a secondary maximum in the 90–100 cm bin. One measurement, of 400 cm, is not shown. B) Variation of the measured plunge directions as a function of pillow maximum dimension. No strong trends are apparent, although the largest pillow dimensions (greater than 150 cm) plunge less than 20°. 258 B.R. Edwards et al. / Journal of Volcanology and Geothermal Research 185 (2009) 251–275 plunge to the north, down the slopes and northern axis of the present (Table 2). Contacts between the dykes and the surrounded units vary ridge. Measured pillow maximum dimensions do not appear to be from straight (Fig. 6A) to undulating (Fig. 6B and C) to bifurcating clearly related to measured pillow plunges (Fig. 5B). At the very (Fig. 6D) to bulbous, with limited observation indicating that contact northern end of the ridge, the pillow trends shift to the northwest. character reflects elevation of exposure (straight=lower on ridge; undulating=higher on ridge). Most of the dykes intruded tuff breccia

4.1.2. Dykes (LD1–2) (TB2), although dykes also intrude Lpw2,4,5 (Table 2). Contact effects As noted by Souther (1992), dykes form a prominent lithofacies in are locally observed as discoloration in tuff breccia (Fig. 6A) and as the Central Ridge area (Table 1; Figs. 3 and 6). They are found on the thin glassy margins to the dykes (Fig. 6C). Internally the dykes are flanks as well as along the top of the ridge crest along Central Ridge heterogeneous in texture and vesicularity. Locally the interiors of (Fig. 2). The lithofacies is subdivided into two units based on domi- dykes are highly vesicular and contain broken fragments of chilled nant orientations, although most of the dykes share the same field margins (Fig. 6F); along the crest of Central Ridge dykes contain characteristics of dark gray colors and highly fractured interiors gabbroic and granitic inclusions. Frequently the dyke interiors are

Fig. 6. Field characteristics of dyke lithofacies at Pillow Ridge. A) Subvertical dyke with straight margins on lower western side of the ridge. Host tuff breccia (TB1) shows slight discoloration near the contact. Dyke is approximately 0.5 m wide. B) Subvertical dyke with irregular, wavy margins on the lower eastern edge of ridge intruding volcanic breccia. Dyke orientation is subparallel to axis of the ridge. The wavy margin has a thin, black, glassy selvage at contact. The rock hammer is 84 cm in length. C) Vesiculated interior of a dyke along eastern margin of the ridge. Platy fragments are disaggregated segments of the dyke margin. B.R. Edwards et al. / Journal of Volcanology and Geothermal Research 185 (2009) 251–275 259

Table 2 sorting by clast size is more evident towards the top of the ridge, with Locations and orientations of dykes. crude layers prominently defined by higher than normal concentra- Field Location Elevation Orientationa Unit Notes tions of juvenile lapilli to block size clasts (Fig. 7C). On the eastern side Station Northing Easting (m ASL) of the ridge the tuff breccia shows some size sorting regardless of Number (m) (m) position. In general clast shapes vary from angular (Fig. 7C) to ovoid

E07-04-01 6403695 0401515 2107 4 LD1 0.50 m wide, throughout the facies, and range up to 1 m in maximum dimension, (?) subvertical, intrudes although more commonly clast sizes vary between 10 and 40 cm. The TB1, vesicular core finer-grained parts of the lithofacies are well indurated and have lapilli E07-04-02 6403666 0401466 2114 306 LD 4 m wide, subvertical, 1 size clasts in an ash-size matrix that appears to be strongly intrudes TB1 E07-04-03 6403637 0401340 2195 81 LD2 ∼1 m wide, dips palagonitized (Fig. 7D). The smaller clasts are monomict, angular,

north, intrudes TB1 slightly vesicular, grey, and porphyritic, with a general basaltic and Lpw , glassy 4 appearance. On the lower, southwestern flank of the ridge TB1 is margin distinct in character in that it has a lighter grey color and is less well E07-04-04 6403577 0401352 2202 316 LD ∼1 m wide, dips 1 indurated than elsewhere. It is difficult to estimate the true thickness west, intrudes TB1

E07-04-09 6403500 0401900 2150 228 LD2 ∼1 m wide, dips for this lithofacies, but it is probably between 150 and 250 m. slightly north, A second unit of tuff breccia (TB2) is intimately associated with intrudes TB1 Lpw4,5 in at least two locations. Grey to black, poorly sorted tuff breccia E07-06-01a 6403701 0401594 2092 240 LD 1.5 m wide, 2 is found immediately underlying Lpw , near North Peak (Table 1; subvertical, 5

intrudes TB1 Figs. 3 and 7E), prominently exposed on its eastern and western sides. ∼ E07-06-01b 6403724 0401614 2071 45 LD2 1 m wide, dips The contacts between TB2 and Lpw4,5 are sharp but irregular. The north, intrudes TB1, deposit is clast-supported by lapilli-size, juvenile, vesicular basalt vesiculated core, fragments, which locally contain white, granitic inclusions. The pillowed margins E07-06-02 6403604 0401555 2142 320 LD 2 m wide, subvertical, thickness of the unit, less than 5 m, remains relatively constant over 1 b intrudes TB1, glassy the length of its exposure, and it dips gently ( 10°) to the north. The internal margins base of TB2 comprises clasts with large vesicles (greater than 2 cm) ∼ E07-06-04 6403373 0401479 2226 321 LD1 0.5 m wide, and fluidal shapes, which grade upwards into clasts with smaller subvertical, intrudes vesicles (less than 0.5 cm) and more angular shapes. Orange to grey TB1, tuff breccia, with very similar clast characteristics, is found along the E07-06-05 6403335 0401465 2210 113 LD1 b0.5 m wide, subvertical, northwestern flank of the ridge, interbedded with Lpw2 (Fig. 7F). intrudes TB1 ∼ ∼ E07-06-08 6403055 0401461 2169 45 LD2 3.5 m wide, 4.2.2. Lapilli tuff (LT) subvertical, intrudes Dark grey lapilli tuff (Table 1; Fig. 7G) occurs as three separate TB1, pillowed margins

E08-03-08 6403637 0401362 2215 302 LD1 ∼1 m wide, layers within a complex section of interbedded lithofacies discontinu- subvertical, intrudes ously separating Lpw4 and Lpw3, on the northwestern flank of the TB1, TB2, and Lpw4, ridge. The upper unit varies in thickness up to 60 cm, the middle is glassy margins much thinner (∼5 cm), while only the uppermost 20 cm of the lower E08-03-09 6403533 0401342 2222 312 LD1 ∼05. m wide, subvertical, intrudes is exposed. All three units are similar in appearance and are

TB2 and Lpw5, moderately well-sorted with the majority of clasts ranging in size pillowed margins from fine to medium lapilli. All of the clasts are juvenile, moderately a Orientations are with respect to True North, with a magnetic declination of 22° east vesicular, and show little to no evidence of palagonitization. The of north. characteristics of this lithofacies are very similar to TB2 except that LT lacks the larger clast sizes and shows better size sorting. pervasively fractured, with dominant fracture directions parallel to dyke margins (Fig. 6G). 4.2.3. Tuff (T) Tuff comprising predominantly ash size juvenile grains occurs in 4.2. Volcaniclastic lithofacies multiple locations along the ridge (Table 1). Ash that is distinctly

orange in color is found locally throughout Lpw3–5, with the greatest 4.2.1. Tuff breccia (TB1–2) concentrations at the base of the pillows but minor amounts Souther (1992) mapped the Central Ridge area as comprising interstitial to individual pillows and locally, inside hollow pillow predominantly ‘sideromelane tuff-breccia’ (p.165). He noted that the cores. Thicker accumulations, up to 14 cm, show well developed unit in general was ‘chaotic’, with a wide variation in the development laminae with cross stratification (Fig. 7G) that persists for 4–6cm. and orientations of bedding, clast size, and clast morphology. We have Locally the upper surface of the unit appears to be slightly deformed subdivided this unit into two lithofacies (Table 1; Fig. 7A–D). by the mass of overlying pillows. Most of the grains are fine to very

The most abundant tuff breccia facies (TB1) is yellow-orange in fine ash, with rare very coarse to fine lapilli size clasts. The grains are color, generally poorly sorted (Fig. 7A and B), and accounts for the juvenile, vitric, angular and non- to slightly vesicular. bulk of the southern half of the ridge (Table 1; Fig. 3). As noted by Dark grey, finely laminated ash that lacks conspicuous cross Souther (1992), the development of bedding is difficult to identify at stratification (Fig. 7G) is only known from one field location, although the base of the ridge, but near the ridge crest it can be well developed it occurs as four stratigraphically separate units. The units varying in (Fig. 7A and B). In general the bedding dips less steeply on the western thickness from 11 to ∼2 cm; units are laterally continuous for ∼5 m along flank of the ridge (Fig. 7A; subhorizontal to 15° west) than on the the length of the outcrop and appear to be subhorizontal. The grains are eastern flank (Fig. 7B; up to 30° east). Bed thicknesses vary from very fine ash size, juvenile, vitric, angular and non- to slightly vesicular. b5 cm to a maximum of 3.5 m, averaging about 50 cm. However, beds Light tan to dark brown ash forms very thin (less than 10 cm) are laterally impersistent and bed amalgamation is common. Bed discontinuous layers within TB1 on the eastern side of Central Ridge.It contacts are typically indistinct with local erosional scours. Traction also occurs at the northwestern end of the ridge, downstream from current structures are very rare, with rare, poorly developed cross- the mouth of Pillow Canyon, as a discontinuous layer in S1. It is less stratification. No bomb sags were recognized. On the western flank, than 5 cm thick, and locally persistent for no more than a few meters. 260 B.R. Edwards et al. / Journal of Volcanology and Geothermal Research 185 (2009) 251–275

Fig. 7. Field characteristics of volcaniclastic lithofacies at Pillow Ridge. A) Bedding and attitude characteristics of TB1 on western side of ridge. Bedding is sub-horizontal; west is to the left. Person for scale is ∼1.5 m tall. B) Large scale characteristics of TB1 on eastern side of ridge. Bedding is dipping to the east, away from the ridge but less steeply than the slope. Person for scale is ∼1.5 m tall. C) Lapilli size clast derived from pillow lava (centre). Thin outer rim of glass is preserved covering most of lapillus. The head of the rock hammer is

20 cm long. D) View of angular lapilli in palagonitized, fine ash matrix. E) Tuff breccia (TB2) with isolated pillow exposed in oblique section. Tuff breccia is overlain by pillow lava.

F) Transition in TB2 showing fluidal clasts with large vesicles and angular clasts with small vesicles. Lens cap for scale is 6 cm. G) Cross-bedding in tuff (T) immediately beneath Lpw4. Grains are predominantly platy sideromelane and are interpreted as hyaloclastite senso stricto. H) Two planar, laminated beds of lapilli tuff (LT; bottom and middle) interbedded with tuff (T2). The tuff units are each approximately 10 cm thick.

4.3. Sedimentary lithofacies Ridge. We classify these lithofacies as sedimentary because the majority of their clasts have been generated or have undergone Four different sedimentary lithofacies are present at Pillow Ridge, subsequent modification (i.e. rounding) by processes not related to which are most extensively exposured along the northwestern edge of active volcanism. Two of the four lithofacies (S2 and C) also contain the ridge in Pillow Canyon (Fig. 2B). The units comprise predominantly materials that are not basaltic and hence not derived from Pillow volcanigenic clasts derived from the primary volcanic units of Pillow Ridge. Given the steep topography and limited bed thicknesses, three B.R. Edwards et al. / Journal of Volcanology and Geothermal Research 185 (2009) 251–275 261 of the four lithofacies are not mappable at 1:20,000 as separate units eastern side of Pillow Canyon. It appears to directly overlie Lpw. Clasts but are lumped together as ‘Sedimentary Lithofacies’ (Fig. 3B); how- range from 4 to 9 cm in maximum dimension (very coarse pebble ever, lithofacies (S1) is more aerially extensive (Fig. 3B). gravel to fine cobble gravel), are subprismoidal to equant in shape, and are commonly covered with a thin orange coating (Fig. 8A and B). The 4.3.1. Breccia (B) lithofacies is oligomict, massive, well-sorted, clast-supported with Orangish grey breccia occurs as a very distinctive lithofacies along very angular clasts of vesicular basalt. The bounding surfaces of many the northwestern edge of the ridge, exposed predominantly along the clasts have the rectilinear appearance of polygonal jointing found in

Fig. 8. Field characteristics of sedimentary lithofacies. A) Outcrop scale photograph showing gross-bedding in outcrop of breccia (B). B) Close-up view of clast-supported breccia interpreted to be paleo-talus. C) Conglomerate (C) interpreted as fluvial lag deposit. D) Volcanic sandstone (S1) with slightly larger vitric clast. E) Boulder clast in S1 with jointing and rounded exterior edge interpreted as fragment of pillow lobe. Polylithic, well-sorted volcanic sandstone. F) Moderately well-sorted volcanic sandstone (S2) with locally abundant vesicular basalt clasts. G) Closer view of S2; lighter grains are pumiceous. 262 B.R. Edwards et al. / Journal of Volcanology and Geothermal Research 185 (2009) 251–275 pillows. It is typically open-framework with local lenses of lapilli- to near the ridge crest and 5 to 7 m in Pillow Canyon, where it locally ash-sized, less well sorted vitric-rich tuff. The apparent thickness on contains lenses of gravel to boulder-size fragments (Fig. 8F) that define slope exposures is N2m. crude bedding. Slump structures and areas of contorted bedding are well developed in the middle section of Pillow Canyon. 4.3.2. Conglomerate (C) Grey to orange conglomerate is exposed in a small lense at the 5. Interpretation of lithofacies associations northern end of Pillow Canyon (Fig. 8C). It is totally encapsulated within lithofacies S1, and comprises clasts that vary in size from b1cm Each of the fifteen lithofacies described at Pillow Ridge is to N10 cm in maximum dimension (medium pebble gravel to coarse interpreted as having formed via processes strongly influenced by cobble gravel). Clasts vary from angular to subrounded and are the presence of ice or water (Table 1). The predominance of a water- polymict, including granite (sensu lato), trachyte, and basalt. It is rich environment is also broadly reflected in the five different massive, poorly sorted, and clast-supported. The exposed part of the lithofacies associations recognized in the field (Table 3). deposit is limited to ∼2 m in horizontal dimension and ∼1.5 m in vertical dimension, and is roughly lensoidal. 5.1. Lithofacies Association A

4.3.3. Sandstone (S1–2) This association comprises Lpw2–5, TB2, and LD1–2 (Table 3); Two different sandstone lithofacies are found at Pillow Ridge, in although pillow lava is volumetrically predominant (Table 1), locally

Pillow Canyon and near the ridge crest. One comprises orange to the abundance of TB2 is significant with respect to adjacent pillow lava tan-grey sandstone (S1) that directly overlies Lpw and B lithofacies in (e.g. North Peak; Fig. 10). Generally the tuff breccia occurs as lenses Pillow Canyon. It appears to be more extensive north of the main intercalated within the pillow lava, as is the case for two of the lower ridgeline than to the south. It is well lithified and forms smoothed pillow lava lithofacies (Lpw2–3) at Pillow Ridge, where relatively thin, outcrops along the small stream at the mouth of Pillow Canyon.Itis discontinuous lenses of TB2 with a palagonitized matrix are inter- massive, moderately sorted with subrounded, oligomict grains of sand bedded with much thicker sequences of Lpw2–3. Locally dykes and size (Fig. 8D,E). The grains are dominantly blocky, partly to small, sill-like intrusions also crosscut TB2, which varies from matrix- moderately vesiculated sideromelane. It is matrix supported but supported to clast-supported. However, the uppermost pillow lava locally has discontinuous lenses of cobble size clasts of vesicular basalt units (Lpw4–5) show a unique structural relationship with TB2 (Fig. with wedge shapes (Fig. 8E) that comprise less than 5 vol.% of the 10). Exposures near North Peak show a layer of TB2 that comprises deposit. S1 shows no apparent size grading, cross stratification or pillow fragments, is of relatively uniform thickness (∼1.6 m; Fig. 10C), bedding, though some slump structures are present locally. The and is laterally persistent for tens of meters beneath the overlying minimum estimated deposit thickness is 5 m. pillow lavas (Fig. 10A,C). The overlying pillow lava rests conformably

The second sandstone lithofacies (S2) comprises orange to dark grey, on an irregular top of the underlying tuff breccia (Fig. 10B), which dips polymict sandstone (Fig. 8F,G). It outcrops in two places: upstream from to the north at ∼5–10° roughly parallel to the ridge axis (Fig. 10A,C).

S1 in the middle section of Pillow Canyon, at the base of the western flank Dykes (LD1) cut through Lpw4 at North Peak and at least into the of the main ridge, and within a stratigraphic sequence directly base of Lpw5. The dykes have irregular, bulbous to pillowed, glassy underlying Lpw4, near the northern crest of the ridge. It is massive and margins where they cut TB2. We interpret them as the feeders to well sorted with subrounded, polymict grains, including black vitric Lpw4,5. grains and rounded, tan pumiceous grains. Locally it shows evidence of Formation of pillow lava is generally thought to imply relatively normal size grading. The minimum deposit thickness is less than 10 cm low rates of lava extrusion (e.g. Griffiths and Fink, 1992) and relatively

Table 3 Summary of distribution, descriptions and interpretations of lithofacies associations at Pillow Ridge.

Association and Stratigraphic relationships Distribution Fig. Interpretation lithofacies present

(A) Lpw2–5,TB2, Lpw2 is interbedded with TB2; Lpw4–5 Northern half of ridge, best exposed on 10 Lpw2–3 predominates but is interbedded irregular lenses of TB2 that

LD1–2 directly overlie TB2 western flank and immediately north probably formed by sporadic gravitational collapse of growing

and east of North Peak pillow pile; TB2 that directly underlies Lpw4–5 is much thicker and more laterally persistent; possibly formed by much larger collapses of oversteepened or temporarily confined pillow pile, after which pillow lobes continued to advance

(B) TB1,LD1–2 LD1–2 cross-cuts TB1 in several locations inCentral to southern half of ridge 6 LD possibly represents the progressive growth of the feeder system

centre of the ridge; no cross-cutting to a Surtseyan eruption, during which TB1 was formed; strait contacts

relationships between LD1 and LD2 with obvious discoloration of TB1 indicates intrusion after surrounding

were noted TB1 was partly lithified; bulbous to pillowed contacts indicate contemporaneous intrusions into the actively forming cone structure

(C) C, S1, TB2, Lpw S1 conformably overlies TB2 and Lpw; Northernmost end of ridge in 11 As pillow emplacement continued at the northern end of the ridge,

C contained within S1 Pillow Canyon axis of pillow pile became more westerly, possibly following sub-ice drainage system, forming a partial barrier to downslope transport of coarse sediments. Finer-grained sediments transported down and over pillow ridge by debris flows into ice-confined canyon, forming a broad apron to the north

(D) S2,TB2,B Predominantly S2 with discontinuous Along northwestern side of ridge, 12 Grains are dominantly subrounded to subangular basaltic material

lenses of TB2, with subordinate B towards the upper, southern end presumably derived from proximal Pillow Ridge units; well-sorted of Pillow Canyon B is formed by gravity-driven grain flow; presence of minor felsic pumice may represent material melted out of overlying ice

(E) Lpw4, TB2,LT, T Lpw4 conformably on top of T, TB2, LT, T In two locations immediately below 13 Materials deposited in differing subaqeuous conditions; bottom

Lpw4 on western side of ridge layers of T in standing water/low energy; TB2 and LT by subaqueous mass flow deposits; lower units of T by unidirectional current to produce cross stratification. Light grey pumice either from syn-eruptive, unknown felsic vent or ice meltout B.R. Edwards et al. / Journal of Volcanology and Geothermal Research 185 (2009) 251–275 263

high confining pressures (Guðmundsson et al., 2002a,b; Tuffen, 2007). emergent ‘Surtseyan’ tephra cone, with the features shown by TB1 The angular shapes of clasts in the tuff breccia are thought to indicate resulting from fragmentation, proximal transport and deposition of that it is a proximal deposit, and many of the larger clasts are whole high concentration mass flows (e.g. Group II deposits of White, 2000; pillows (Fig. 7E) or fragments of pillows (Fig. 7F). All occurrences of Fig. 9). Such eruptions and their resultant deposits in glaciovolcanic

TB2 are interpreted as proximal pillow breccia derived from local, environments are relatively common and have been describe else- gravity driven collapse of pillow lava (Fig. 9). However, the North Peak where in detail (e.g. Skilling, 1994; Smellie and Hole, 1997). Extensive sections of TB2 seem to require relatively larger collapse events, as palagonitization is consistent with deposition in the presence of hot they are anomalously thick and continuous. It may be that they water. The absence of bomb sags is consistent with predominantly formed beneath advancing complexes of pillow lava, in a manner subaqueous deposition. Variations in contacts at dyke margins, from analogous to the development of lava flow foot breccias in subaerial straight and sharp to bulbous, are consistent with continued dyke lava flows (e.g. Jones, 1970). Alternatively, they may have formed from intrusion during the course of the eruption. The abundance of dykes collapse or explosive disruption of larger pillow mounds by episodes and the predominant orientations of the longest dykes subparallel to of more rapid emplacement, or by dyke emplacement into the base of the axis of the ridge (Table 2; Fig. 3B) are consistent with the overlying ice (e.g. Wilson and Head, 2004). Many workers have formation of the ridge mainly via fissure eruptions. described lava deltas, where pillow lava breccia forms in response to subaerial lava flowing into water (cf. Skilling, 2002). However, water 5.3. Lithofacies Association C concentrations greater than the one atmosphere solubility (∼0.1 wt.%) in pillow rim glasses are not consistent with formation of pillows by At the northern end of Pillow Canyon, the contact relationships subaerial lava flowing into water, nor is the relatively shallow dip of between Lpw3(?), S1, TB2 and C are well-exposed (Table 3; Fig. 11A–C). TB2 where observed. The axis of Pillow Ridge, as defined by the main mass of pillow lavas, changes strike from a northwesterly direction to a westerly direction 5.2. Lithofacies Association B close to where Pillow Canyon cuts through the ridge, effectively creating a cross-section almost perpendicular to the ridge axis (Fig. 11D).

This association comprises TB1 and LD1–2 (Table 3); the tuff breccia Lithofacies TB2 forms a partial carapace directly overlying the mound- is volumetrically predominant (Table 1). The tuff breccia is interpreted shaped ridge of Lpw3(?) (Fig.11A) and is directly, conformably overlain by as palagonitized proximal to medial phreatomagmatic (Surtseyan) S1, which locally contains lenses of coarser fragments (Fig.11B) as well as tephra (Fig. 9). The angular, basaltic clasts within the tuff breccia a single, irregularly-shaped pod of C (Fig. 11B). To the north and appear to mainly be derived from jointed pillow lava. The clasts were downstream, S1 continues for a few hundred meters. The carapace of TB2 likely sourced from three areas: (1) Surtseyan eruptions through the is never more than 3–5 m in thickness and most prominent immediately pre-existing pillow lava pile (Fig. 7C), (2) pillow lava talus collapsing overlying the mass of pillow lava. To the south, on the upstream side of onto a Surtseyan tephra slope, and (3) fragmentation of dykes and the ridge, very little to no S1 is present. Assigning the pillow lava (Lpw3(?)) pillowed dyke margins by subsequent eruptions. The relatively poor at the northernmost end of the ridge to a stratigraphic unit is tentative; sorting, coarse clasts, and crude parallel to massive bedding imply although this association includes the lowest elevation exposures of rapid proximal deposition. This is also consistent with the apparent pillow lava at Pillow Ridge, the orientation of pillows for the upper units lack of traction current structures and obvious size grading. Primary (Lpw3–5) indicates that they all flowed down the ridge axis to the north. bedding up to 20–25° is interpreted to indicate deposition by grain This lithofacies association is interpreted largely as a distal and avalanches. Discontinuous lenses of clasts might also be generated by down drainage equivalent of association (A), with the larger volumes periodic dyke intrusion and dissagradation into the growing cone. of finer-grained clastic materials having been derived by high Souther (1992) noted an apparent increase in angular clasts that he concentration turbidity currents from the upslope axis of the ridge, interpreted, based on vesicle and shape characteristics, as fragments while TB2 was derived from local pillow collapse (Fig. 9). The of pillow lava near the ridge crest. mounded shape of the pillow lava, with overlying clastic materials, We interpret this association as recording the formation of a is consistent with growth of the pillow ridge into an ice-confined relatively small (∼1.5 km long by 300 m high) subaqueous to space, possibly similar to the scenario described at Mt. Pinafore by

Fig. 9. Cartoon showing possible depositional environments for Pillow Ridge lithosfacies. Not to scale. 264 B.R. Edwards et al. / Journal of Volcanology and Geothermal Research 185 (2009) 251–275

Fig. 10. Field relationships of Lithofacies Association A. A) Photomosaic showing northern sections of pillow lava units (Lpw4–5) immediately underlying North Peak looking from the west. Approximate boundaries between lithofacies are also shown. B) View showing contact between Lpw5 (above) and TB2 (below). C) View looking to the north showing TB2 underlying Lpw5. D) View of vertical cooling joints in lobe of Lpw5 immediately overlying TB2.

Smellie and Skilling (1994). The apron of S1 extending to the north, as defined bedding planes evident in vertical exposures (Fig. 12A–C), well as the conglomerate lense it contains, is interpreted as indicating and the presence of non-basaltic clasts, including quartzofeldspathic that this lithofacies formed in an environment where sub-ice drainage sandstone (Fig. 12E) and light grey pumice (Fig. 12F). played a role in shaping the deposit characteristics (Fig. 9). Lack of Lithofacies B looks strikingly like present-day, active talus slopes cross stratification or other indications of unidirectional currents is seen along the steep western flank of Pillow Ridge, except that it is consistent with deposition as debris flows in a lacustrine environment partly consolidated (Fig. 8A). Most of the cobble-sized clasts are from actively growing, upslope pillow lava emplacement. The western vesicular and bounded by polygonal surfaces. Lithofacies B is bend in the axis of the pillow-dominated ridge might have acted as a interpreted as lithified talus derived from in situ weathering and barrier to coarser debris, preventing it from being carried further disaggregation of proximal jointed pillow lava by gravity-driven grain north (Fig. 11D). The smaller, sand-size debris was carried over the flow over an indeterminate period of time (McPhie et al., 1993; van ridge crest to form a wedge-shaped deposit that thins to the north. Steijn et al., 2002). Periodic re-exposure of B or similar deposits would provide a steady source for the clasts seen in many deposits at Pillow

5.4. Lithofacies Association D Ridge (e.g. TB1–2, S1–2) long after volcanic activity at the edifice ceased. Poorly defined bedding in S1 has been deformed over a broad area Most of Pillow Canyon appears to have formed by erosion at the of exposure (tens of metres; Figs. 9–12A–C). The lithofacies is margin of the pillow-dominated units. Exposed in the canyon walls is presently lithified, but the lack of evidence for deep burial or brittle a mixture of lithofacies (Table 3; Fig. 12A–C), dominated by deformation is consistent with the deformation having occurred sedimentary deposits (S2, B) with only minor volumes of TB2 or its before S1 was completely consolidated. In an ice-rich environment, derivatives. The deposits in this association record a variety of several possible events could be responsible for the deformation, two characteristics not commonly preserved or observed elsewhere at of which seem to be the most likely. One possibility is that the Pillow Ridge. The most intriguing are the preservation of well-sorted, lithofacies was deposited in an ice-free cavity along the edge of Pillow pillow fragment breccia (Fig. 8A,B), the deformation of crudely Ridge either immediately after an eruption had melted space in the B.R. Edwards et al. / Journal of Volcanology and Geothermal Research 185 (2009) 251–275 265

Fig. 11. Field relationships of Lithofacies Association C, exposed at the distal, northwestern end of Pillow Ridge. A) Northern-most exposures of pillow lava tubes, overlain by coarse volcanic sandstone (S1) containing lenses of tuff breccia (TB2). B) Contact between S1 and massive unit of TB2. C) Partly exposured section of cobble-filled channel in S1. D) Schematic cross-section parallel to the ridge axis showing asymmetric distribution of finer-grained materials (S1) to the north (downstream), and thicker coarse-materials (TB2) to the south (upstream). surrounding ice, or during a subsequent interglacial period. Immedi- the base of the slope. However, the timing of formation for these ately after eruption, ice would have flowed back into the cavity formed deposits is difficult to constrain with respect to the eruption history of by melting (e.g. Guðmundsson et al., 2002a,b; Tuffen, 2007), Pillow Ridge. deforming the poorly consolidated S1. The local flow of ice would be The abundance of non-basaltic grains and clasts in S2 is also inwards towards the ridge axis, and it could exert a small compressive consistent with its deposition at the ice-contact margin of the ridge force against the partly consolidated S1. A second possibility is that the (Fig. 9). The most conspicuous component not obviously derived from sediments were originally emplaced on top of snow or ice, essentially basaltic Pillow Ridge materials is light grey pumice, present as sand- to forming an ice-cored alluvial fan. When the underlying ice melted, the gravel-sized grains (Fig. 12F). The grains are subrounded, but would deposits would begin sag into the cavity left by the ice. Minor slope not likely have survived extensive transport. We have identified three failures could have produced the observed bedding distortions as the possible sources for the pumice. The first is a deposit at the base of beds were deformed by compression while slowly moving downslope the southwestern end of the ridge of poorly consolidated, lapilli- or by buckling once the leading edge of the block of material reached dominated tephra (Fig. 3B). The deposit has pumiceous lapilli and 266 B.R. Edwards et al. / Journal of Volcanology and Geothermal Research 185 (2009) 251–275

Fig. 12. Field relationships of Lithofacies Association D. A) Photomosaic of outcrop in upper part of Pillow Canyon showing heterogeneous nature of S2 as well as folded/deformed lenses of coarser fragments. Person for scale. B) Sketch of (A) to highlight distorted lenses. C) More severely contorted bedding on upper part of slope above outcrop shown in (A).

D) Close-up view of lense of cobble-size vesicular basalt clasts in S2. E) Disc-shaped, quartz-rich metasedimentary clast in S2. F) Light colored pumiceous clasts in S2.

round, fluidal clasts, both of which could be vent-proximal airfall. It is (1992) suggested that Pillow Ridge formed largely between eruptive directly overlain by TB2. A second possible source is well-illustrated by episodes of Ice Peak Formation and Edziza Formation, both of which Holocene deposits on the western flank of the MEVC. Modern ice at comprise in part trachyte pyroclasts. Mount Edziza contains conspicuous lenses of pumice from the Sheeptrack eruption, thought to have occurred more recently than 5.5. Lithofacies Association E 10 ka (Souther,1992). Similar deposits in the ice extant during eruptions at Pillow Ridge could also have provided a source for the pumice (Fig. 9). A relatively inconspicuous but complex sequence of lithofacies

Based on the probably age of the Sheeptrack pumice, the pumice present occurs along the contact between Lpw3 and Lpw4 (Fig.13), immediately in S2 could be at least 10 ky older than the Pillow Ridge eruptions. The south of North Saddle; the sequence of volcaniclastic and sedimentary presence of pumice could also indicate an eruption of more evolved units outcrops at two different locations. At the largest exposure, five compositions approximately contemporaneously with the Pillow Ridge different lithofacies (TB2, LT, T, S2, Lpw4) comprising eleven distinct eruptions, although this seems the least likely possible source. Souther units are found in stratigraphic succession (Fig. 13A, B, D). Lithofacies B.R. Edwards et al. / Journal of Volcanology and Geothermal Research 185 (2009) 251–275 267

Fig. 13. Field relationships of Lithofacies Association E, exposed in between Lpw3–4, immediately south of North Saddle (Fig. 2B). A) Photographic cross-section of outcrop. Hammer for scale is 1 m long. B) Sketch of (A) with labels for individual lithofacies labeled to show relative stratigraphic positions. C) View looking towards section in (A) showing its position immediately underlying thick sequence of pillow lava (Lpw4). D) Detailed lithostratigraphic log through lithofacies showing relative average grain sizes.

LT occurs at two different horizons: at the base, where the exposed Two thin (less than 10 cm) units of volcanic sandstone (S2) overlie thickness is at least 10 cm, and in the middle of the sequence, where it the lower LT unit. The lithofacies occurs as distinct horizons in the occurs as a discontinuous lense. The lense appears to be made of large field, and comprises grains of poorly vesiculated sideromelane and blocks of LT that have a jigsaw fit, as if the entire lense was one large pumice. The lithofacies is inferred to have a sedimentary origin based block that had been fractured but not disrupted. Both occurrences of LT on the presence of pumice grains, although the majority of the grains are moderately well sorted but grains show little evidence of rounding. appear to be locally derived basaltic material from Pillow Ridge. 268 B.R. Edwards et al. / Journal of Volcanology and Geothermal Research 185 (2009) 251–275

Table 4 Major element geochemistry whole rock (WR) and glass analyses from Pillow Ridge samples.

Oxides C2a PR59a E06BE46b E6BE46c E6BE41c E6BE40c E6BE44c E6BE48c WR WR WR Glass Glass Glass Glass Glass

SiO2 50.62 47.20 48.49 48.66 51.69 49.69 50.91 49.72

TiO2 2.45 1.90 2.41 3.08 2.74 2.95 2.62 2.90

Al2O3 17.02 16.20 16.19 14.27 15.41 14.85 14.95 15.09

Fe2O3 3.20 3.10 –––––– FeO 5.38 8.40 11.87 12.23 11.28 11.54 11.25 11.85 MnO 0.10 0.18 0.16 0.19 0.20 0.19 0.18 0.18 MgO 3.93 7.53 4.14 4.03 3.79 3.58 3.66 4.31 CaO 9.55 10.70 9.64 9.01 8.16 7.85 8.12 8.79

Na2O 3.73 2.80 3.41 3.58 4.01 3.29 3.96 3.79

K2O 1.76 0.66 1.44 1.85 2.19 2.37 2.27 1.84

P2O5 0.40 0.24 0.40 ––––– S ––– 0.20 0.21 0.23 0.22 0.20 Cl− ––– 0.04 0.05 0.05 0.05 0.05 d H2O ––– 0.542 0.655 0.631 0.81 0.542 LOI 2.52 0.20 2.21 Total 100.57 99.73 100.48 97.15 99.74 96.6 98.18 98.71 Mg/Mg+Fe 39.0 39.0 39.0 36.3 37.4 34.8 36.3 39.0

a From Souther (1992). b Bulk rock analysis from Geochemical Labs, McGill University. c Quantitative analyses from University of Chicago Electron Microprobe Analyzer; data for each sample are averages of at least three spot analyses. d Concentrations of H2O in glasses not included in ‘Total’ as they were determined by FTIR at University of Wisconsin-Milwaukee.

Two variations of lithofacies T are present in this section. The locally with glassy selvages and wedge-shaped jointing; these are grains in both have similar characteristics and are mixtures of larger, interpreted as fragments of pillow lava. Towards the northern end of vesiculated sideromelane granules and finer, blocky to elongate, the exposure, it appears as though two separate units of TB2 are poorly vesiculated sideromelane granules. Much of this material is present, separated by a relatively thick (up to 60 cm) unit of LT. interpreted as having formed by spallation of vitric rinds on pillow However, towards the south end of the exposure the upper and lower lava. The lowermost three layers of T lack obvious bedding structures; horizons of TB2 seem to merge around the block of LT. This poorly however the uppermost unit appears to be distinctly cross-stratified sorted sequence, including the upper block of LT is interpreted as (Fig. 7F). Such structures imply deposition from a current, and could resulting from a subaqueous debris flow. be an indication of shallow water deposition. Alternatively, such Our interpretation of this sequence is that it formed during a features are also consistent with deceleration of a deeper water period of cessation of pillow-forming eruptions between the empla- density current. The unit is directly overlain by Lpw4, and some of the cement of Lpw3 and Lpw4. All of the deposits could have formed by tuff appears to have been squeezed up into gaps between overlying sedimentation from turbidity currents triggered from upslope by pillow lava. It is also possible that the apparent cross-laminations seismic activity, talus avalanches or ice collapse. The thick package of formed by disruption of bedding when the overlying pillows were TB2 may have been transported as a block during a subaqueous emplaced. avalanche; its matrix appears to be slightly more consolidated and

Lithofacies TB2 directly overlies the middle unit of T, and locally the palagonitized than the bounding lithofacies. Much of the material in contact is erosional. Most of the clasts in TB2 are vesiculated basalt, the sequence could represent more distal equivalents of TB2, which

Table 5

Analyses of H2O, calculated liquidus temperatures, and estimated confining pressures, water depths, and ice thicknesses for samples of glass from pillow rims from Pillow Ridge.

Sample Sample elevation Wt. % Terupt Wt. % Methodc Pressure Water depth Ice thick. Min. ice elevation Max. ice elevation a b d e f number m ASL H2O (°C) CO2 (MPa) (m) (m) (m) (m) E06BE46 2100 0.542 1119 25 N&L 8.4 857 934 2396 3034 0 N&L 2.9 296 322 0 GM 4.5 459 501 E06BE41 2023 0.655 1108 25 N&L 9.5 967 1055 2429 3078 0 N&L 4.0 406 443 GM 6.2 633 690 E06BE40 1923 0.631 1109 25 N&L 9.1 929 1013 2289 2936 0 N&L 3.6 366 401 0 GM 5.8 592 645 E06BE44 1894 0.810 1106 25 N&L 11.7 1190 1297 2522 3191 0 N&L 6.2 628 685 0 GM 8.9 908 990 E06BE48 ∼1700 0.542 1118 25 N&L 8.4 857 934 1996 2634 0 N&L 2.9 296 322 0 GM 4.5 459 501

a Determined by FTIR at University of Wisconsin-Milwaukee. b Calculated using MELTS (Ghiorso and Sack, 1995; Asimow and Ghiorso, 1998) from glass compositions and water contents (see Table 2). c Calculated using VOLATILECALC (Newman and Lowenstern, 2002) and assuming default SiO2 weight percent value of 49 [denoted by ‘N&L’], and H2OSOLXI, which does not explicitly account for the presence of CO2 (Moore et al., 1998; G. Moore, pers. comm. 2008) [denoted ‘GM’]. d Calculated from t=ρ⁎g⁎P. e Assuming a base elevation of 1600 m, corresponding approximately to the plateau elevation underlain by the Ice Peak Formation, and using the minimum water depth, assuming that for water to be impounded ice must be at least as thick as minimum water depth. f Assuming a base elevation of 1600 m and the maximum ice thickness for the sample. B.R. Edwards et al. / Journal of Volcanology and Geothermal Research 185 (2009) 251–275 269

underlies Lpw4 upslope at North Peak (Fig. 10A). The grey pumice 50.08 to 51.85 wt.%, and values of S and Cl also showing little variation from S2 horizons likely had one of the same three origins as the (Table 4). However, molar Mg/Mg+Fe shows more variation, with pumice clasts in Lithofacies Association D. values for glasses ranging from 39.3 to 35.6 (Table 4). Few other analyses of glasses from glaciovolcanic edifices have been published 6. Geochemistry for samples from the NCVP, with two exceptions: work by Moore et al. (1995) on samples from Ash Mountain, South Tuya and Tuya Butte, Materials from lithofacies Lpw were analyzed for major element and by Dixon et al. (2002) on samples from Tanzilla Butte. Both works compositions and for H2O concentrations (Fig. 3A; Tables 4 and 5). The reported glass compositions from glaciovolcanic deposits ∼100 km material erupted at Pillow Ridge is hawaiite to alkaline basalt in north of Pillow Ridge in the Tuya volcanic field. The glass compositions composition based on major element analyses of whole rock and glass reported from the two studies have ranges in SiO2 (48.94–51.56 wt.%) samples (Table 4, Fig. 14). Mineral compositions for olivine, plagio- and Cl (0.014–0.053 wt.%) broadly similar those for samples from clase and clinopyroxene reported by Souther (1992) and Lloyd (2007) Pillow Ridge, but with lower values of S (0.04–0.11 wt.%). are consistent with alkaline basaltic compositions. All of the samples from which vitric material was analyzed are from pillow lavas that 6.2. Water contents in pillow rim glass were vesiculated, which is indicative of volatile saturation in the melt phase (Höskuldsson et al., 2006). Larger fresh glass chips were selected for FTIR analysis. The chips were made into doubly polished wafers approximately 100 μm thick.

6.1. Major element compositions of glass samples Dissolved total H2O concentrations were measured at the University of Wisconsin-Milwaukee using a Nicolet Nexus 470 FTIR with a Major elements, S and Cl concentrations of glasses from pillow Continuum Analytical-IR microscope attachment using KBr beam rims were measured using a Cameca SX-50 electron microprobe at the splitter and a liquid nitrogen-cooled MCT-A detector. A 100 μm University of Chicago. An acceleration voltage of 15 keV, a beam diameter square aperature was used during the analysis, although the current of 25 nA, and a beam diameter of ∼20 μm were used for all size and shape of the aperture was modified to avoid vesicles and measurements. Counting times for S and Cl were 90 and 60 s, microlites present in some of the glass samples. Each spectrum respectively. A barite and chloro-apatite were used as mineral consisted of 512 scans between 650 and 4000 cm− 1, taken at 4 cm− 1 standards for S and Cl, respectively. Three Smithsonian glass standards resolution. In basaltic glasses, total H2O is calculated using the (Indian Ocean basaltic glass USNM 113716, VG-A99 basaltic glass vibration band at 3550 cm− 1. The molar absorptivity coefficient USNM 113498/1, VG-2 basaltic glass USNM 111240/52) were also used for this wave number is 63 l mol-cm− 1 (Dixon et al., 1988). used to monitor accuracy and precision. An alpha scan for the S peak Density of basaltic glass was estimated to be 2800 kg m− 3. Glass wafer on the VG-2 standard necessitated a slight shift in the S peak from the thickness was determined by a Mitutoyo digital micrometer, micro- mineral standard. VG-2 from the Juan de Fuca ridge was run nine scope focusing, and by FTIR interference fringes in reflectance mode. times during the microprobe session with a mean S and Cl content of Estimations of wafer thickness using a digital displacement gauge are 0.147±0.005 and 0.028±0.005 wt.%. Our results on VG-2 were very no better than +/−3 μm. The greatest uncertainty related to similar to Gurenko et al. (2005). thickness determinations using interference fringes in reflectance Major element compositions of glass samples show some variation, mode is the refractive index of the basaltic glass, for which we used a particularly with respect to whole rock analyses. The glasses have value of 1.546 (Kumagi and Kaneoka, 2003). concentrations of SiO2, TiO2 and K2O higher than whole rock analyses, The measured H2O concentrations from Pillow Ridge samples vary and lower Al2O3 and MgO concentrations. Variations between the from 0.54 to 0.81 wt.%. These values are very similar to those reported glass samples are smaller, with normalized values of SiO2 varying from by Dixon et al. (2002) for Tanzilla Butte (0.49–0.92 wt.%), but generally higher than the limited measurements for Tuya Butte

glasses from Moore et al. (1995;0.19–0.56 wt.%). Reported H2O concentrations in the rims of glaciovolcanic pillow lavas from Iceland are generally lower than values from British Columbia (Nicholls et al., 2002; Schopka et al., 2006; Höskuldsson et al., 2006). Saturation pressures were estimated by two different methods

(Table 5). Measured H2O concentrations were input into VolatileCalc (Newman and Lowenstern, 2002) assuming a basaltic composition of

49 wt.% SiO2, eruption temperatures calculated using MELTS (Ghiorso and Sack, 1995; Asimow and Ghiorso, 1998), and estimated concen-

trations of CO2 of either 0 or 25 ppm. Calculations for 0 ppm CO2 gave eruption pressures range from 2.9 to 6.2 MPa. Given the assumption that the pillow lava formed by eruption in water, minimum and

maximum ice thicknesses can be calculated for both CO2 values. If the eruption melted through the overlying ice very rapidly, most of the pillow lava would have been emplaced in an englacial lake; thus the confining pressure would be that generated by the mass of overlying water. If the eruption produced a water-filled cavity within the ice, than the overlying pressure could be mainly that of the overlying ice. The minimum ice thickness would be that necessary to confine a lake whose depth is equal to the calculated confining pressure. Assuming a − 3 Fig. 14. Total alkalies versus silica classification diagram (Le Bas et al., 1986) comparing water density of 1000 kg m the estimated minimum ice thicknesses data from analyses of samples from Pillow Ridge (Souther, 1992; this work) to those for for the Pillow Ridge samples are between 296 and 628 m (Table 5). the rest of the MEVC (Souther, 1992), and the NCVP (cf. Edwards and Russell, 2000). In However, the ice thickness estimate is strongly dependent on the CO2 general samples of mafic rocks from the NCVP are mildly alkaline, while more evolved content of the glass. The absence of carbonate peaks during the FTIR compositions are moderately to highly alkaline. Sample from the MEVC span most of the compositional range for the entire NCVP. All samples from Pillow Ridge plot above analysis suggests that the glasses contain less than 30 ppm, the the alkaline–subalkaline division of Irvine and Baragar (1971). accepted detection limit for FTIR. Based on manometry measurements 270 B.R. Edwards et al. / Journal of Volcanology and Geothermal Research 185 (2009) 251–275

of CO2 on glasses from nearby Tennena Cone, an alkaline basalt, the ridge, and always above Souther's ‘aquagene tuff-breccia’. It also glaciovolcanic cone on the southwestern flank of Mount Edziza implies the presence of obvious subaerial tephra and/or lava, the

∼10 km south of Pillow Ridge, we have assumed a CO2 concentration existence of a stable englacial lake throughout the duration of the of 25 ppm for the Pillow Ridge samples to estimate the maximum ice eruption, and fully degassed pillow lavas, as they formed essentially as thickness for Pillow Ridge samples (Table 5). The maximum ice a lava delta from initially subaerial lava flows. We will address each of thickness was calculated by estimating the highest likely concentra- these predictions below to test Souther's model and determine what, tion of CO2 and assuming that the resulting confining pressure is due if any, refinements need to be made to explain the observations and solely to a mass of overlying ice. The resulting confining pressures data presented in this paper. range from 8.7 to 11.7 MPa, corresponding to ice thickness estimates of Our fieldwork shows that pillow lavas can be found at the base of 722 to 1302 m. the ridge in at least two places; on the western side, along a stream In order to corroborate the VolatileCalc estimates, we repeated the demarcating the contact between the ridge and moraine-covered ice

0 ppm CO2 calculations using H2OSOLXI (Moore et al., 1998; Moore, and upstream from the start of Pillow Canyon, and at the extreme pers. comm., 2008). It gives results for minimum ice thicknesses northwestern end of the ridge, at the mouth of Pillow Canyon. In the ranging from 459 to 908 m, which are intermediate between the upstream section, the pillow lavas are subhorizontal, are the lowest minimum and maximum ranges from VolatileCalc (Table 5). However, exposures of Pillow Ridge material on the western side of the ridge,

H2OSOLXI cannot be used to explicitly account for variations in CO2 and appear to be directly overlain by the thickest part of the ridge. concentrations. Much of the western side of the ridge appears to have been excavated by ice; this is also consistent with the lowest exposed pillow lava lobes 7. Discussion having been originally under the crest of the ridge. We interpret the presence of basal pillow lava to indicate that the initiation of the 7.1. Two contrasting models for the formation of Pillow Ridge eruptions effusively. This is an important observation as it has implications for the pressure regime, and hence ice/water thickness, Souther (1992) proposed a three stage model for the development at the start of the eruption. The downstream exposures of pillow lava of Pillow Ridge (Fig. 15A). He suggested that the eruption started appear to be the northern-most, downslope extension of Lpw3, which explosively within an ice-dammed lake, leading to the formation of a partly covers the tuff breccia lithofacies. subaerial tephra cone intruded by dykes. The final stages of the Souther (1992) proposed the existence of a significant subaerial eruption produced lava flows that traveled into the lake to form pillow tephra cone, in addition to subaerial lava flows capping the southern lavas. Several predictions can be made from this model to test its part of the ridge. We were unable to confirm either of these findings in efficacy for explaining field and laboratory observations. Firstly, this the field. Our brief reconnaissance at the southern end of the ridge model implies that pillow lavas should only be found along the top of suggested that it comprised predominantly pillow lava and tuff

Fig. 15. Comparison of two models for the evolution of Pillow Ridge tindar. A) Schematic cross-section after Souther (1992) with summary of pertinent events and problems. B) Schematic cross-section based on this work, with summary. B.R. Edwards et al. / Journal of Volcanology and Geothermal Research 185 (2009) 251–275 271

Fig. 16. Constraints on paleo-ice thicknesses/water depths with relationships to important elevations. Minimum (light grey) and maximum (dark grey) water depth/ice thickness measurements based on H2O contents of pillow rind glass are shown as are the elevations for Pillow Ridge, Mount Edziza, and the inferred elevations for Ice Peak volcano and the Cordilleran ice sheet.

breccia, but no obvious subaerial lava flows. The main mass of TB1 in The next phase of the eruption was explosive and phreatomag- the central part of the ridge does not appear to comprise a subaerial matic. It is difficult to determine whether this change in eruption style tephra cone. Field and preliminary microscopic analysis is consistent is due to: a change in the nature of the magma being erupted (e.g. with much of TB1 having formed by Surtseyan eruptions, which might becoming more volatile rich, or simply having lower viscosity have breached the surface of an ice-confined lake, but not with resulting from a decrease in the volume of inclusions), a change in extensive subaerial eruptions. eruption rate (higher rates favor more fragmentation and less pillow Finally, Souther's model implies that the pillow lavas formed by formation; e.g. Griffiths and Fink, 1992), or a change in confining flow of subaerial lava into an englacial lake. If this were true, the pressure. A drop in confining pressure could result from upwards pillow lavas should be largely degassed. However, FTIR measurements growth of the pillow pile, from a drainage event (e.g., Höskuldsson of H2O in the glassy rims of individual pillows shows that the lavas et al., 2006), or breaching of the ice surface. Recent eruptions in were not totally degassed; all analyses imply confining pressures Iceland show that it takes less than 36 h from the start of an eruption equivalent to a water depth of at least ∼290 m. As well, field to melt through ice at least 600 m thick (Guðmundsson et al., 2004). observations show that individual pillow lava lobes are commonly Whatever the cause, at this stage the eruption became more energetic moderately vesicular (e.g. Fig. 4E), and locally the interiors are highly and produced a subaqueous cone at least 300 m in height. Most of the vesicular, which would be somewhat unexpected if the lavas were tuff breccia in the central part of the ridge is poorly to moderately degassed. sorted, and most of the clasts lack oxidized, scoriaceous surfaces as We propose an alternative model for the formation of Pillow Ridge would be expected for a predominantly subaerial eruptions. We (Fig. 15B), which is broadly consistent with Souther's field observa- suggest that most of this part of the ridge formed by subaqueous vent tions as well as our own. The initiation of the basaltic eruption at proximal deposition during Surtseyan eruptions in an englacial lake.

Pillow Ridge began with emplacement of pillow lava in a sub-ice The presence and abundance of dykes cutting through TB1 is meltwater lens. Confining pressures derived from H2O measurements consistent with pulses of magma being supplied to the system over indicate minimum ice thicknesses of ∼300–600 m (Table 5; Fig. 16), a period of at least weeks to months, and possibly much longer which is consistent with non-explosive effusion of pillow lava. (years). The variations in the nature of the LD–TB1 contacts showing Although the contacts are largely covered by talus, the central mass that some were intruding relatively coherent material (straight of TB1 appears to directly overlie at least the basal three pillow lava contacts) while others intruded relatively unconsolidated TB1 (bul- units (Lpw1–3). These pillow lobes are distinctive from subsequent bous contacts). The majority of the TB1 clasts could have been derived pillow units in two ways. Firstly, they appear to be on average slightly from pillow lava lobes based on their morphologies. It is likely that the larger in diameter, although we do not have enough field measure- change to more energetic eruptive activity disrupted parts of the ments to quantify this change accurately. Secondly, individual pillow earlier formed pillow units and dykes. lava lobes locally are packed with up to 50 vol.% gabbroic inclusions. The upper two pillow units clearly cover the upper part of TB1; Most of the inclusions are only a few tens of cm long, but at least one however estimated confining pressures (Table 5) indicate that the found in the field was N30 cm long. Gabbroic and granitic inclusions pillows, which were all from intact lava lobes and showed no evidence are found higher in the stratigraphy as well, but are much less for having been transported after their formation, formed beneath at abundant. The presence of the inclusions is important for two reasons. least 300–600 m of water and/or ice. The techniques, measurements They indicate that the erupted lava was either stored in a magma and calculations used to infer water/ice thickness are still actively reservoir for some time before eruption, or that the magmas passed being developed and refined, so that the absolute thickness estimates through pre-existing gabbroic intrusions before eruption. The abun- have a significant uncertainty (Dixon et al., 2002; Schopka et al., dance of inclusions would also have caused a significant increase in 2006); however we have no reason to believe that the relative the effective viscosity of the lava, and is consistent with the lower differences in confining pressures are not real. We interpret the units having pillow lobes with larger diameters (e.g. Walker, 1992). variations in pressure estimates as recording changes in the Our field measurements are also broadly consistent with a viscosity paleohydrology of the ice-confined lake, possibly coupled with control on pillow sizes (Fig. 5B), as pillow maximum dimensions do slowdown of the rate of eruption. Our constraints from H2O not appear to correlate with plunge of the emplacement slope. measurements show that confining pressures seem to have varied 272 B.R. Edwards et al. / Journal of Volcanology and Geothermal Research 185 (2009) 251–275 over the course of pillow emplacement (Fig. 16). These changes may (2002) for Tanzilla Butte (300–900 m) to the north. Given that the indicate changes in the level of a progressively enlarging and filling elevation difference between the assumed basal of Pillow Ridge englacial lake. It is possible, if not likely, that the lowermost and (1660 m) and the high point of the MEVC (ca. 2800 m) is approximately uppermost pillow units did not form within the same eruptive event, 1150 m, the minimum ice depths could be accommodated easily by a and that they formed under different ice sheets with different thick local ice cap centered on the MEVC (Fig. 16). However, if the ice thicknesses. The 1996 Gjálp eruption in Iceland formed in the same thicknesses were closer to the maximum estimates (Fig. 16), they location as a pre-existing glaciovolcanic ridge thought to have formed indicate very thick ice that could be consistent with eruption into a during an eruption in 1938 (Guðmundsson et al., 2002a,b). Thus regional ice-sheet. Although few physical constraints exist for mini- glaciovolcanic ridges can build progressively over time-periods that mum thicknesses for the CIS, geophysical models (e.g. Marshall et al., span decades if not longer. 2002) suggest that the CIS had a maximum thickness of between 2 and At least one episode of relative quiescence is documented between 3 km. Given that the present day elevation of the land surface under the

Lpw3 and Lpw4 by Lithofacies Association E. presumed northern point of accumulation for the CIS is ca. 1 km, an ice- The final events identified in the evolution of the ridge were the sheet of 2–3 km thickness would have a relative surface elevation up to emplacement of the uppermost two units of pillow lava, Lpw4 and 1 km higher than the present-day elevation of the MEVC, assuming that Lpw5. Both of these units are underlain by TB2, which is somewhat present day topography generally mimics paleotopography beneath the anomalous for its lack of palagonitization of finer vitric grains in CIS at 1 Ma. To our knowledge this is the only potential constraint for comparison with TB1 and stratigraphically lower occurrences of TB2.It regional ice thicknesses of the CIS at this time. is possible that these last two units could be significantly younger In order to assess the potential impact of the Pillow Ridge than the rest of the ridge, although we did not observe any obvious eruptions on the overlying ice sheet, we have examined the potential unconformities in the field that would be consistent with long periods thermal output from the eruption and estimated how much ice could of erosion preceding their formation. have been melted during the eruption. Many previous workers have In contrast to Souther's model (Fig. 15A), we have found evidence examined the heat transfer between glaciovolcanic eruptions and for a complex evolution of Pillow Ridge, including multiple episodes of overlying ice (Allen, 1980; Höskuldsson and Sparks, 1997; Guðmunds- pillow emplacement and likely changes in water/ice thickness over son, 2003; Kelman et al., 2002; Guðmundsson et al., 2004; Wilson and the course of the eruption. Radiometric datings of samples from Pillow Head, 2004; Tuffen, 2007). The most general conclusions resulting Ridge are in progress and will, in the near future, help to resolve from these studies is that basaltic lava can melt substantial amounts of questions about the duration of formation of the ridge. Future work is ice (greater than 10 times the volume of lava erupted for pillow lava if also planned to determine the eruption age of the light grey pumice heat exchange efficiency is high; Höskuldsson and Sparks, 1997), and from S1, which may be an indication of previously undocumented, that explosive fragmentation is a more efficient way to rapidly release contemporaneous felsic volcanism. the heat than effusion of lava (heat efficiency as high as 70–80%; Guðmundsson, 2003; Guðmundsson et al., 2004). 7.2. Implications for local and regional ice-sheets Because Pillow Ridge has significant volumes of fragmental material and coherent pillow lava, we use slightly different approaches The long-term objective of our research at the MEVC is to use and heat efficiency values for each in order to estimate the maximum glaciovolcanic deposits to constrain paleo-ice conditions for local ice and minimum amounts of ice that could have been melted during the sheets and for the CIS in the northern Cordillera and examine the formation of the ridge. Using the appropriate thermodynamic values, impact of glaciovolcanic eruptions on the ice sheets. Our work at Equation 2 of Guðmundsson et al. (2004), and the assumption of no Pillow Ridge is an example of the type of detailed paleo-environ- significant heating of water above 0 °C, we estimate that complete mental information that can be preserved in glaciovolcanic deposits. crystallization of 1 m3 of Pillow Ridge basalt could melt at most 17 m3 Historically, most of the glaciovolcanic deposits that might relate to of ice, assuming 100% efficiency of heat transfer, and at a minimum changes in the CIS have been attributed to the last glacial maximum 1.7 m3, assuming 10% efficiency (e.g., Guðmundsson (2003) estimated (Frasier/Wisconsinan; ∼20 ka; Mathews, 1947; Hickson, 2000; Dixon an efficiency range for effusive basaltic eruptions of 10–45%). Given et al., 2002; Kelman et al., 2002). However, Souther (1992) reported that the estimated volume of pillow lava erupted at Pillow Ridge is K–Ar geochronology indicating that Pillow Ridge formed ∼900 ka. ∼0.7 km3 (40% of 1.8 km3), eruption of pillows alone could melt a More recent work at Edziza (Spooner et al., 1995) and in the nearby minimum of ∼1.2 km3 of ice. Assuming melting was concentrated Stikine River canyon (Spooner et al., 1996)hasdocumented immediately above the volcanic ridge and that all of the pillows were glaciogenic sediments deposited by regional ice circa 1 Ma and 340 ka. emplaced in one episode, heat released from the pillow lava could melt Minimum ice thicknesses extant during glaciovolcanic eruptions through ice less than 300 m thick (assuming an ice cavity that has can estimated in at least two ways: based on the maximum elevation horizontal dimensions similar to those of the ridge, e.g. 4 km long by of pillow lavas or other clearly subaqueous deposits (e.g. Smellie, 1 km wide). Similar calculations for crystal free cooling of 1 m3 of vitric 2000), and by measuring the volatile contents of undegassed samples Pillow Ridge basalt show that it could melt at most 13.8 m3 of ice, (Moore and Calk, 1990; Moore et al., 1995; Dixon et al., 2002; assuming 100% efficiency of heat transfer, and at a minimum 6.9 m3, Höskuldsson et al., 2006; Schopka et al., 2006). The highest elevation assuming 50% efficiency (minimum value given by Guðmundsson et al. pillow lavas at Pillow Ridge are ca. 2100 m, which is ∼500 m above the (2004) for the 1996 Gjálp eruption, which was dominated by 3 elevation of the surrounding plateau; we interpret this to indicate that fragmentation). Given that the estimated volume of TB1 is ∼0.9 km when the uppermost pillow lavas were erupted, the enclosing ice (50% of 1.8 km3), the Surtseyan stage of the eruption could melt a must have been more than 500 m thick. This is consistent with the minimum of ∼6.2 km3 of ice. Assuming melting was concentrated formation of Pillow Ridge during a period of relatively ‘thick’ ice, immediately above the volcanic ridge, heat released from TB1 could either flowing from the local MEVC ice-cap or potentially from a major melt a cavity in ice less than ∼1500 m thick (assuming an ice cavity that advance of the CIS. The water contents of undegassed samples from was 4 km long by 1 km wide). Even if the Pillow Ridge eruption took pillow rim glass can also be used to estimate the confining pressure of place beneath the thickest estimates for the CIS, the combined effects overlying ice/water during their emplacement, given a number of of effusive and fragmental eruptions could release enough thermal assumptions that have been discussed in detail by Schopka et al. energy to melt completely through extremely thick ice. However, these (2006). Estimated ice thicknesses from undegassed pillow rims calculations have several caveats. If Pillow Ridge formed during the (Table 5) give minimum values of 300–600 m, and maximum values course of several eruptions widely-spaced in time, which is possible, of 900–1300 m, which are similar to those estimated by Dixon et al. individual units, especially from pillow eruptions, would have been B.R. Edwards et al. / Journal of Volcanology and Geothermal Research 185 (2009) 251–275 273 able to melt through ice only ∼2 times the thickness of the eruption facies and five lithofacies associations were recognized. A deposits. Modeling of transitions between effusive and explosive revised model for the formation of Pillow Ridge better explains eruptions by Tuffen (2007) indicates that low discharge rates, which the evolution of the ridge, including multiple episodes of may be required for pillow forming eruptions (e.g. Griffiths and Fink, subaqueous emplacement of pillow lavas, and more detailed 1992), can lead to intrusive style eruptions for a range of reasonable evidence for specific depositional processes. The lithofacies sub-ice cavity underpressures due to ductile ice flow during the course associations provide detailed constraints on eruption and of the eruption. Tuffen (2007) suggested that ductile flow may be depositional processes extent ca. 900 ka at the northern end important on the timescale of individual eruptions. This may explain of the Mount Edziza volcanic complex. the apparent presence of intrusive pillow lava in Lpw2. (2) Subtle variations in glass compositions and water concentra- tions are consistent with multiple distinct periods of pillow 7.3. Implications for the formation of tindars emplacement under variable water/ice conditions. Estimates of minimum ice thicknesses are consistent with eruption beneath Tindars have the potential to preserve important constraints on a local ice sheet (300–600 m), whereas estimates of maximum eruption and magma dynamics, paleotectonics, and paleoclimate. ice thicknesses are most consistent with emplacement beneath Their presence documents the existence of ice at the time of eruption. a regional ice sheet (900–1300 m). The wide variation in the relative volumes of volcaniclastic (e.g. (3) The abundance of dykes, predominantly oriented parallel to the Schopka et al., 2006) versus coherent lithofacies (e.g. Höskuldsson et axis of the ridge, is interpreted as physical evidence that the al., 2006) is undoubtedly controlled by a number of factors acting in eruptions responsible for the formation of the ridge were isolation or together, including initial volatile contents of pre-eruptive predominantly fissure-fed. Contacts between dykes and enclos- magmas, effusion rates, and confining pressures. Tindars dominated ing tuff breccia vary from sharp, nearer to the base of the ridge, by volcaniclastic lithofacies are potential recorders of variations in ice to pillowed and wavy higher up on the ridge. We interpret this paleohydrology, and when mapped in detail can be used to determine as showing that the exposed dykes were emplaced syn- directions of meltwater drainage (e.g., Schopka et al., 2006). Those eruption and that formation of the breccia resulted in part dominated by coherent facies have the potential to record ice from fragmentation of the previously emplaced units of pillow thicknesses and may preserve evidence for depressurization events lava and/or dykes. caused by meltwater drainage (e.g., Höskuldsson et al., 2006). Tindars (4) Regardless of ice thickness, the total heat evolved during the like Pillow Ridge and Kalfstindar (Jones, 1970) that fall between the evolution of Pillow Ridge was sufficient to melt through the two end members may provide the best opportunities to more overlying icesheet, given realistic assumptions about the completely document dynamic events in the co-evolution of the efficiency of heat transfer. enclosing ice and the volcanic vents. (5) Tindars are very important volcanic features as they cannot Tindars may form during a single eruption, or may be the result of only be used to constrain volcanic, magmatic, and tectonic multiple eruptions from the same fissure system. For example, the processes, but most importantly can provide critical constraints 1996 Gjálp eruption occurred partly along the side of a ridge produced for paleoclimate studies focused on characterizing the presence by an eruption in 1938 (Guðmundsson et al., 2002a,b). Glass and thickness of terrestrial ice deposits. compositions from Pillow Ridge pillows are consistent with liquids derived by small amounts of fractional crystallization from an alkaline Acknowledgements basalt parent. The whole rock composition is significantly higher in MgO, consistent with the presence of phenocrysts of olivine. However, Funding for the project comes from NSF-EAR 0439707 to BRE, NSF- the glasses do not show a systematic decrease in Mg Number with EAR 0439699 to IPS, and the Research and Development committee at elevation (Fig. 16); they seem to show a repetition of cycling that Dickinson College. We thank Chira Endress and Kristen LaMoreaux for could indicate eruptions from a source that has been replenished/ assistance in the field in 2006, and Curtis Brett for assistance in 2008. recharged multiple times. We also thank Janice Joseph, Maxine Graydon, and Chris Price, all at BC Although tindars vastly outnumber tuyas in Iceland (Chapman et al., Parks, for logistical and planning help during fieldwork, as well as Jim 2000), in British Columbia tindars are subordinate to tuyas. In Iceland and Nancy Reed at the Pacific Western Helicopters for help with Jones (1969) recognized that the linear forms of the tindars were transport logistics. Reviews by H. Tuffen, J. Stevenson and J. Dixon consistent with the dominant tectonic stress regime of Iceland, which is were insightful and helped to improve the manuscript. parallel to the extensional plate boundary that divides the island. The Pleistocene tectonic setting of northern British Columbia is less well References known, but thought to be dominated by transtension along the Queen Charlotte fault system (e.g., Souther, 1992; Edwards and Russell, 1999), Allen, C.C., 1980. Icelandic subglacial volcanism: thermal and physical studies. Journal of which is the boundary between the Pacific and North American plate off Geology 88, 108–117. Allen, C.C., Jercinovic, M.J., Allen, J.S.B., 1982. Subglacial volcanism in north-central shore of the western coast of British Columbia. It may be that tindars at British Columbia and Iceland. Journal of Geology 90, 699–715. MEVC preserve evidence for that tectonic regime around 900 ka. Pillow Asimow, P.D., Ghiorso, M.S., 1998. Algorithmic modifications extending MELTS to Ridge, trending at ∼315°, is approximately parallel to the Queen calculate subsolidus phase relations. American Mineralogist 83, 1127–1131. Barendegt, RW, Irving, E, 1998. Changes in the extent of North American ice sheets Charlotte fault system. Limited measurements of dyke orientations during the Late Cenozoic. Canadian Journal of Earth Science 35, 504–509. show two different trends (Table 2). It is possible that the orientation of Benson, LV, May, HM, Antweiler, RC, Brinton, TI, Kashgarin, M, Smoot, JP, Lund, SP, 1998. Pillow Ridge and its two dyke sets reflect local stress field associated Continuous lake-sediment records of glaciation in the Sierra Nevada between 14 – with the MEVC Edziza complex (crustal warping from added weight). 52,600 and 12500 C yr BP. Quaternary Research 50, 113 127. Bye, A., Edwards, B.R., Hickson, C.H., 2000. Preliminary field, petrographic and Changes in the dyke orientations might also reflect changes in the mass geochemical analysis of volcanism at the Watts Point volcanic centre, southwestern of a large, overlying ice-sheet during the course of the formation of the British Columbia: a subglacial, dacite lava dome? Current Research, Part A; ridge, especially if the eruptions were widely spaced in time. Geological Survey of Canada (GSC) Paper2000-A20. 9 pp. Chapman, M.G., Allen, C.C., Guðmundsson, M.T., Gulick, V.C., Jakobsson, S.P., Lucitta, B.K., Skilling, I.P., Waitt, R.B. 2000. Volcanism and ice interactions on Earth and Mars. In 8. Conclusions Zimbelman, J.R. & Gregg, T.K.P. (eds.) Environmental effects on volcanic eruptions. Kluwer Academic/Plenum Publishers, New York, 39–74. Clague, J.J., 1989. Cordilleran ice sheet. In: Fulton, R.J. (Ed.), Quaternary Geology of (1) The evolution of Pillow Ridge tindar was complex and involved Canada and Greenland, Geology of North America. Geological Society of America, significant changes in eruption styles. Fifteen types of litho- vol. K-1, pp. 40–42. 274 B.R. Edwards et al. / Journal of Volcanology and Geothermal Research 185 (2009) 251–275

Clague, J.J., James, T.S., 2002. History and isostatic effects of the last ice sheet in southern Jackson, LE, Barendregt, RW, Baker, J, Irving, E, 1996. Glacial history of the central Yukon: British Columbia. Quaternary Science Reviews 21, 71–87. new evidence from field studies and paleomagnetism. Canadian Journal of Earth Dixon, J.E., Stolper, E.M., Delaney, J.R., 1988. Infrared spectroscopic measurements of Science 33, 904–916. CO2 and H2O in Juan de Fuca Ridge basaltic glasses. Earth and Planetary Science Jakobsson, S, Guðmundsson, MT, 2008. Subglacial and intraglacial volcanic formations Letters 90, 87–104. in Iceland. Jokull 58, 179–196. Dixon, JE, Filiberto, JR, Moore, JG, Hickson, CJ, 2002. Volatiles in basaltic glasses from a James, TS, Clague, J, Wang, K, Hutchinson, I, 2000. Rapid deglaciation of the subglacial volcano in northern British Columbia (Canada): implications for ice Cordilleran ice sheet and crustal rebound. European Geophysical Society sheet thickness and mantle volatiles. In: Smellie, J.L., Chapman, M.G. (Eds.), Meeting, Nice, France. Volcano–ice Interaction on Earth and Mars. Special Publications, vol. 202. Jones, J.G.,1969. Intraglacial volcanoes of the Laugarvatn region, southwest Iceland I. Quaterley Geological Society, London, pp. 255–271. Journal of the Geological Society of London 124, 197–211. doi:10.1144/gsjgs.124.1.0197. Edwards, B.R., Russell, J.K., 1999. Northern Cordilleran volcanic province: a northern Jones, J.G., 1970. Intraglacial volcanoes of the Laugarvatn region, southwest Iceland, II. Basin and Range? Geology 27, 243–246. Journal of Geology 78, 127–140. Edwards, B.R., Russell, J.K., 2000. The distribution, nature and origin of Neogene– Kelman, MC, Russell, JK, Hickson, CH, 2002. Effusive intermediate glaciovolcanism in the Quaternary magmatism in the Northern Cordilleran Volcanic Province, northern Garibaldi volcanic belt, southwestern British Columbia, Canada. In: Smellie, J.L., Canadian Cordillera. Geological Society of America Bulletin 112 (8), 1280–1295. Chapman, M.G. (Eds.), Volcano–ice Interaction on Earth and Mars. Special Edwards, B.R., Russell, J.K., 2002. Glacial influence on the morphology and eruption Publications, vol. 202. Geological Society, London, pp. 195–212. products of Hoodoo Mountain volcano, northwestern British Columbia. In: Smellie, Kumagi, H., Kaneoka, I., 2003. Relationship between submarine MORB glass textures J., Chapman, M. (Eds.), Volcano–ice Interaction on Earth and Mars. Geological and atmospheric component of MORBs. Chemical Geology 200, 1–24. Society of London Special Publication, vol. 202, pp. 179–194. Kutzbach, J.E., 1987. Model simulations of the climatic patterns during the deglaciation Edwards, B.R., Edwards, G., Russell, J.K., 1995. Revised stratigraphy for the Hoodoo of North America. In: Ruddiman, W.F., Wright, H.E. (Eds.), North America and Mountain volcanic center, northwestern British Columbia. Current Research, part A; Adjacent Ocean during the Last Deglaciation. Geology of North America, vol. K-3. GSC Paper 95-1A, pp. 105–115. Geological Soceity of America, Boulder, USA. Edwards, B.R., Anderson, R.G., Russell, J.K., Hastings, N.L. & Guo, Y.T. 1999. Geology of the Le Bas, M.J., Le Maitre, R.W., Streckeisen, A., Zanettin, B., 1986. A chemical classification Quaternary Hoodoo Mountain Volcanic Complex and adjacent Paleozoic and of volcanic rocks based on the total alkali–silica diagram. Journal of Petrology 27, Mesozoic basement rocks; parts of Hoodoo Mountain (NTS 104B/14) and Craig 745–750. River (NTS 104B/11) map areas, northwestern British Columbia. Geological Survey Lloyd, A. 2007. Plagioclase growth rates in pillow basalts: a study in thermal modeling of Canada, Open File Report, scale 1:20 000. and crystal size distribution. Unpublished Honors Thesis, Dickinson College, 54p. Edwards, B.R., Russell, J.K., Anderson, R.G., 2002. Subglacial, phonolitic volcanism at Marshall, SJ, James, TS, Clark, GKC, 2002. North American Ice Sheet reconstructions at Hoodoo Mountain volcano, northwestern Canadian Cordillera. Bulletin of Volca- the Last Glacial Maximum. Quaternary Science Reviews 21, 175–192. nology 64, 254–272. doi:10.1007/s00445-002-0202-9. Mathews, W.H., 1947. “Tuyas”, flat-topped volcanoes in northern British Columbia. Edwards, B.R., Evenchick, C.A., McNicoll, V., Nogier, M., Wetherell, K., 2006. Overview of American Journal of Science 245, 560–570. the volcanology of the Bell–Irving volcanic district, northwestern Bowser Basin, McPhie, J., Doyle, M., Allen, R., 1993. Volcanic textures: a guide to the interpretation of British Columbia: new examples of alpine glaciovolcanism from the northern textures in volcanic rocks. Hobart, CODES Key Centre, University of Tasmania. 196 pp. Cordilleran volcanic province. Current Research, Geological Survey of Canada (GSC) Moore, J.G., Calk, L., 1990. Degassing and differentiation in subglacial volcanoes, Iceland. Paper2006-A3. 12 pp. Journal of Volcanology and Geothermal Research 46, 157–180. Fulton, R.J., 1984. Quaternary glaciation, Canadian Cordillera. In: Fulton, R.J. (Ed.), Moore, J.G., Hickson, C.J., Calk, L., 1995. Tholeiitic–alkalic transition at subglacial Quaternary Stratigraphy of Canada, a Canadian contribution to IGCP Project 24. volcanoes, Tuya region, British Columbia. Journal of Geophysical Research 100, Geological Survey of Canada paper vol. 84-10, pp. 39–48. 24,577–24,592. Fulton, R.J., 1992. Stratigraphy and paleomagnetism of Brunhes and Matuyama Moore, G.M., Vennemann, T., Carmichael, I.S.E., 1998. An empirical model for the (N790 ka) Quaternary deposits at Merritt, British Columbia. Canadian Journal of solubility of H2O in magmas to 3 kb. American Mineralogist 83, 36–42. Earth Sciences 29, 76–92. Newman, S., Lowenstern, J.B., 2002. VolatileCalc: a silicate melt-H2O–CO2 solution Fulton, RJ, Ryder, JM, Tsang, S, 2003. The Last Glaciation of British Columbia, Canada. XVI model written in Visual Basic. Computers & Geosciences 28, 597–604. INQUA Congress, Reno, USA. Nicholls, A.R.L., Carroll, M.R., Höskuldsson, A., 2002. Is the Iceland hot spot also wet? Ghiorso, MS, Sack, RO, 1995. Chemical mass transfer in magmatic processes. IV. A revised Evidence from the water contents of undegassed submarine and subglacial pillow and internally consistent thermodynamic model for the interpolation and extrapola- lavas. Earth and Planetary Science Letters 202, 77–87. tion of liquid–solid equilibria in magmatic systems at elevated temperatures and Raymo, ME, 1992. Global climate change: a three million year persepective. In: Kukla, pressures. Contributions to Mineralogy and Petrology 119, 197–212. Went (Eds.), Start of a Glacial, NATO ASI Series, Series 1. Global environmental Griffiths, RW, Fink, JH, 1992. Solidification and morphology of submarine lavas: a change, vol. 3, pp. 207–223. dependence on extrusion rate. Journal of Geophysical Research 97, 19729–19737. Rittman, A., 1962. Volcanoes and Their Activity. Interscience Publishers, New York. Guðmundsson, M.T., 2003. Melting of ice by magma–water interactions during 320pp. subglacial eruptions as an indicator of heat transfer in subaqueous eruptions. Ryder, JM, Maynard, D, 1991. The Cordilleran Ice Sheet in Northern British Columbia. Explosive Subaqueous Volcanism. Geophysical Monograph, vol. 140, pp. 61–72. Geographie Physique et Quaternaire 45, 355–363. Guðmundsson, M.T., Palsson, F., Björnsson, H., Hognadottir, T.H., 2002a. The hyaloclas- Schopka, H., Guðmundsson, M.T., Tuffen, H., 2006. The formation of Helgafell, tite ridge formed in the subglacial 1996 eruption at Gjálp, Vatnajökull, Iceland: southwest Iceland, a monogenetic subglacial hyaloclastite ridge: sedimentology, present day shape and future preservation. In: Smellie, J.L., Chapman, M.G. (Eds.), hydrology and volcano–ice interaction. Journal of Volcanology and Geothermal Volcano–ice Interaction on Earth and Mars. Special Publications, vol. 202. Research 152, 359–377. doi:10.1016/j.jvolgeores.2005.11.010. Geological Society, London, pp. 319–335. Simpson, K., 1996. The geology, geochemistry and geomorphology of Mathews tuya: a Guðmundsson, M.T., Sigmundsson, F., Björnsson, H., Hognadottir, T.H., 2002b. The 1996 subglacial volcano in northwestern British Columbia [B.Sc. Honours thesis]. eruption at Gjálp, Vatnajökull ice cap, Iceland: course of events, efficiency of heat University of British Columbia, Vancouver, British Columbia, 97 p. transfer, ice deformation and subglacial water pressure. Bulletin of Volcanology 66, Skilling, I.P.,1994. Evolution of an englacial volcano: Browns Bluff, Antarctica. Bulletin of 46–65. Volcanology 56, 573–591. Guðmundsson,M.T.,Sigmundsson,F.,Björnsson,H.,Hognadottir,T.H.,2004.The Skilling, I.P., 2002. Basaltic pahoehoe lava-fed deltas: large-scale characteristics, clast 1996 eruption at Gjalp, Vatnajokull ice cap, Iceland: efficiency of heat transfer, generation, emplacement processes and environmental discrimination. In: Smellie, ice deformation and subglacial water pressure. Bulletin of Volcanology 66, J.L., Chapman, M.G. (Eds.), Volcano–ice Interaction on Earth and Mars. Geological 46–65. Society, London, Special Publications, vol. 202, pp. 91–113. Gurenko, A.A., Belousov, A.B., Trumbull, R.B., Sobolev, A.V., 2005. Explosive basaltic Smellie, J.L., 2000. Subglacial eruptions. In: Sigurdsson, H. (Ed.), Encyclopedia of volcanism of the Chikurachki Volcano (Kurile arc, Russia): insights on pre-eruptive Volcanoes. Elsevier, New York, pp. 403–419. magmatic conditions and volatile budget revealed from phenocryst-hosted melt Smellie, J.L., 2006. The relative importance of supraglacial versus subglacial meltwater inclusions and groundmass glasses. Journal of Volcanology and Geothermal 147, escape in basaltic tuya eruptions: an important unresolved conundrum. Earth- 203–232. Science Reviews 74, 241–268. Harder, M., Russell, J.K., 2005. Basanite glaciovolcanism at Llangorse Mountain, Smellie, J.L., Hole, M.J., 1997. Products and processes in Pliocene–Recent, subaqueous to northern British Columbia, Canada. Bulletin of Volcanology 69 (3), 329–340. emergent volcanism in the Antarctic Pennisula: examples of englacial Surtseyan doi:10.1007/s00445-006-0078-1. volcano construction. Bulletin of Volcanology 58, 628–646. Hickson, C.J., 2000. Physical controls and resulting morphological forms of Quaternary Smellie, J.L., Skilling, I.P., 1994. Products of subglacial volcanic eruptions under different ice-contact volcanoes in western Canada. Geomorphology 32, 239–261. ice thicknesses—2 examples from Antarctica. Sedimentary Geology 91, 115–129. Honnorez, J., Kirst, P., 1975. Submarine basaltic volcanism: morphometric parameters Souther, J.G., 1992. The late Cenozoic Mount Edziza volcanic complex. Geol. Soc. Can. for discriminating hyaloclastite from hyalotuff. Bulletin of Volcanology 39, Mem., vol. 420. 320 pp. 441–465. Souther, J.G., Hickson, C.J., 1984. Crystal fractionation of the basalt comendite series of Höskuldsson, A., Sparks, R.S.J., 1997. Thermodynamics and fluid dynamics of effusive the Mount Edziza volcanic complex, British Columbia: major and trace elements. subglacial eruptions. Bulletin of Volcanology 59, 219–230. Journal of Volcanology and Geothermal Research 21, 79–106. Höskuldsson, A., Sparks, R.S.J., Carroll, M.R., 2006. Constraints on the dynamics of Souther, J.G., Armstrong, R.L., Harakal, J., 1984. Chronology of the peralkaline, late subglacial basalt eruptions from geological and geochemical observations at Cenozoic Mount Edziza volcanic complex, northern British Columbia, Canada. Kverkfjoll, NE-Iceland. Bulletin of Volcanology 68, 689–701. doi:10.1007/s00445- Geological Society of America Bulletin 95, 337–349. 005-0043-4. Spooner, I.S., Osborn, G.D., Barendregt, R., Irving, E., 1995. A record of early Pleistocene Irvine, T, Baragar, WRA, 1971. A guide to the chemical classification of the common glaciation on the Mount Edziza Plateau, northwestern British Columbia. Canadian volcanic rocks. Canadian Journal of Earth Science 8, 523–548. Journal of Earth Science 32, 2046–2056. B.R. Edwards et al. / Journal of Volcanology and Geothermal Research 185 (2009) 251–275 275

Spooner, I.S., Osborn, G.D., Barendregt, R., Irving, E., 1996. A middle Pleistocene (Isotope Walker, G.P.L., 1992. Morphometric study of pillow-size spectrum among pillow lavas. Stage 10) glacial sequence in the Stikine River valley, British Columbia. Canadian Bulletin of Volcanology 54, 4590474. Journal of Earth Science 33, 1428–1438. White, J.D.L., 2000. Subaqueous eruption-fed density currents and their deposits. Stumpf, A.J., Broster, B.E., Levson, V.M., 2000. Multiphase flow of the late Wisconsinan Processes in Physical Volcanology and Volcaniclastic Sedimentation: Modern and Cordilleran ice sheet in Western Canada. Geological Society of American Bulletin Ancient Precambrian Res., vol. 101, pp. 87–109. 112, 1850–1863. White, J.D.L., Houghton, B.F., 2006. Primary volcaniclastic rocks. Geology 34, 677–680. Tuffen, H., 2007. Models of ice melting and edifice growth at the onset of basaltic Wilson, L., Head, J.W., 2004. Heat transfer and melting in subglacial basaltic volcanic subglacial eruptions. Journal of Geophysical Research 112, B03203. doi:10.1029/ eruptions: implications for volcanic deposit morphology and meltwater volumes. 2006JB004523. In: Smellie, J.L., Chapman, M.G. (Eds.), Volcano–ice Interaction on Earth and Mars. van Steijn, H., Boelhouwers, J., Harris, S., Hetu, B., 2002. Recent research on the nature, Special Publications, vol. 202. Geological Society, London, pp. 5–26. origin and climatic relations of blocky and stratified slope deposits. Progress in Physical Geography 26, 551–575.