The architecture, eruptive history, and evolution of the Table Rock Complex, Oregon: From a Surtseyan to an energetic maar eruption

Journal of Volcanology and Geothermal Research 180 (2009) 203–224

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Journal of Volcanology and Geothermal Research j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / j vo l g e o r e s

The architecture, eruptive history, and evolution of the Table Rock Complex, Oregon: From a Surtseyan to an energetic maar eruption Brittany D. Brand ⁎, Amanda B. Clarke 1 School of Earth and Space Exploration, Arizona State University, Box 871404, Tempe, AZ, 85287-1404, United States of America

a r t i c l e

i n f o

Available online 28 October 2008 Keywords: phreatomagmatic Surtseyan maar base surge Table Rock Complex

a b s t r a c t The Table Rock Complex (TRC; Pliocene–Pleistocene), first documented and described by Heiken [Heiken, G.H., 1971. Tuff rings; examples from the Fort Rock-Christmas Lake valley basin, south-central Oregon. J. Geophy. Res. 76, 5615-5626.], is a large and well-exposed mafic phreatomagmatic complex in the Fort Rock–Christmas Lake Valley Basin, south-central Oregon. It spans an area of approximately 40 km2, and consists of a large tuff cone in the south (TRC1), and a large tuff ring in the northeast (TRC2). At least seven additional, smaller explosion craters were formed along the flanks of the complex in the time between the two main eruptions. The first period of activity, TRC1, initiated with a Surtseyan-style eruption through a 60–70 m deep lake. The TRC1 deposits are dominated by multiple, 1-2 m thick, fining upward sequences of massive to diffusely-stratified lapilli tuff with intermittent zones of reverse grading, followed by a finely-laminated cap of fine-grained sediment. The massive deposits are interpreted as the result of eruption-fed, subaqueous turbidity current deposits; whereas, the finely laminated cap likely resulted from fallout of suspended fine-grained material through a water column. Other common features are erosive channel scour-and-fill deposits, massive tuff breccias, and abundant soft sediment deformation due to rapid sediment loading. Subaerial TRC1 deposits are exposed only proximal to the edifice, and consist of cross-stratified base-surge deposits. The eruption built a large tuff cone above the lake surface ending with an effusive stage, which produced a lava lake in the crater (365 m above the lake floor). A significant repose period occurred between the TRC1 and TRC2 eruptions, evidenced by up to 50 cm of diatomitic lake sediments at the contact between the two tuff sequences. The TRC2 eruption was the last and most energetic in the complex. General edifice morphology and a high percentage of accidental material suggest eruption through saturated TRC1 deposits and/or playa lake sediments. TRC2 deposits are dominated by three-dimensional dune features with wavelengths 200–500 m perpendicular to the flow, and 20–200 m parallel to the direction of flow depending on distance from source. Large U-shaped channels (10–32 m deep), run-up features over obstacles tens of meters high, and a large (13 m) chute-and-pool feature are also identified. The TRC2 deposits are interpreted as the products of multiple, erosive, highly-inflated pyroclastic surges resulting from collapse of an unusually high eruption column relative to previously documented mafic phreatomagmatic eruptions. © 2008 Elsevier B.V. All rights reserved.

1. Introduction Hydromagmatic eruptions occur when rising magma violently fragments after intersecting and mixing with shallow surface water or groundwater (Sheridan and Wohletz, 1983). Fragmentation in this style of volcanism is driven principally by the energetic interaction between magma and external water (Houghton and Wilson, 1989), although expansion of magmatic volatiles can occur depending on the volatile content of the magma, and may provide a secondary mechanism of fragmentation (i.e., Houghton and Wilson, 1989; Houghton et al., 1999; Cole et al., 2001; Brand et al., in press). The degree of fragmentation associated with magma–water interaction

⁎ Corresponding author. E-mail addresses: [email protected] (B.D. Brand), [email protected] (A.B. Clarke). 1 Tel.: +1 480 965 6590; fax: +1 480 965 8102. 0377-0273/$ – see front matter © 2008 Elsevier B.V. All rights reserved. doi:10.1016/j.jvolgeores.2008.10.011

has been experimentally and theoretically determined to be a function of melt composition, magma flux, water–melt mass ratio, confining pressure, magma viscosity, and the degree of turbulent mixing of magma with water, steam, or water sprays (Sheridan and Wohletz, 1983; Wohletz and McQueen, 1984; Büttner and Zimanowski, 1998; Zimanowski et al., 1991; Mastin, 2007). In the last couple of decades there have been many advances in understanding how the hydromagmatic deposits of tuff cones, tuff rings, and maars relate to the eruptive dynamics and depositional mechanisms that produced them (e.g., Fisher and Waters, 1970; Crowe and Fisher, 1973; Lorenz, 1974; Sheridan and Wohletz, 1983; Kokelaar, 1983; Fisher and Schmincke, 1984; Houghton and Hackett, 1984; Kokelaar, 1986; Sohn and Chough, 1989; Dellino et al., 1990; White, 1996; Houghton et al., 1999; White, 2001; Nemeth et al., 2001; Cole et al., 2001; Mastin et al., 2004; Brand and White, 2007; Brand et al., in press). To assess the relative influence of external water (water–magma ratio) on overall eruption dynamics, researchers look at deposits for evidence for liquid

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water at the time of deposition. Wet conditions are typically indicated by a combination of features such as accretionary lapilli, fine-grained vesiculated tuff, pervasive soft-sediment deformation, surge cross-strata with stoss-side accretion and evidence for plastering, mud cracks, debris-flow filled erosion channels, and lahar deposits (Waters and Fisher, 1971; Lorenz, 1974; Wohletz and Sheridan, 1983; Sohn and Chough,1989; Dellino et al.,1990; Nemeth et al., 2001; White, 2001; Cole et al., 2001; Brand and White, 2007; Brand et al., in press). Wetphreatomagmatic conditions are interpreted to occur when external water at the zone of magma–water interaction is not efficiently converted to steam (Sheridan and Wohletz, 1983; Wohletz and McQueen, 1984) and abundant liquid water is retained in the eruption column and density currents. Dry conditions are distinguished by their lack of features indicative of liquid water at the time of pyroclastic deposition. These deposits contain no evidence of accretionary lapilli, little to no soft-sediment deformation, low-angle cross strata with dominantly lee-side accretion (although antidunes are also common), and no evidence for muddy deposits plastering against obstacles or on the stoss side of dunes (e.g., Fisher and Waters, 1970; Sheridan and Wohletz, 1983; Sohn and Chough, 1989; Chough and Sohn, 1990; Doubik and Hill, 1999; Nemeth et al., 2001; Brand and White, 2007). Dry-phreatomagmatic conditions represent efficient conversion of external water to steam at the site of magma–water interaction, or subsequently in the eruption plume. As a consequence, little or no liquid water is retained in the eruption column or proximal density currents, although steam may condense in density currents and rising plumes with distance from the source (e.g., Sheridan and Wohletz, 1983; Wohletz and McQueen, 1984). These conditions are thought to represent highly efficient conversion of thermal energy to kinetic energy, yielding the highest degree of melt fragmentation and vapor production, and therefore the highest explosivity (Wohletz and McQueen, 1984; Büttner and Zimanowski, 1998). Many moderately-to well-exposed remnants of tuff cones, tuff rings, and maars exist in south-central Oregon (Peterson and Groh, 1961, 1963; Heiken, 1971; Heiken et al., 1981). The Table Rock Complex (TRC; Pliocene–Pleistocene), one of the best exposed examples of hydrovolcanism in the Fort Rock-Christmas Lake Valley Basin, was first mapped and described by Heiken (1971). We have revisited TRC with the goal of studying the stratigraphy in detail in order to reconstruct the volcanic evolution of the complex. Our results provide evidence for two major eruptions, TRC1 and TRC2, for which we constrain the temporal evolution, dominant depositional mechanisms, the influence of liquid water on deposit characteristics, and relative eruption energy. The objectives of our work are to build on the existing framework for hydromagmatic pyroclastic deposits by continuing to identify relationships between deposit characteristics and eruptive dynamics of mafic hydromagmatic eruptions. The main topics of this paper include (1) subaqueously emplaced density currents of a large Surtseyan-style eruption early in the complex's history (TRC1); and (2) The formation of multiple, largescale base surge deposits produced during a later, highly energetic maar-forming eruption (TRC2). Finally, a more general, broad-scale reason for studying tuff cones, tuff rings, and maars is that they offer an opportunity to observe, describe, and study the deposits of many pyroclastic processes on easily accessible vertical and lateral scales. 1.1. Geologic setting The Fort Rock–Christmas Valley basin is 64 km long by 40 km wide, and was the location of an extensive, ancient Pliocene to Pleistocene lake (lake boundary dashed in Fig. 1; Heiken, 1971). The basin is occupied by a variety of basaltic eruptive features which trend northwest to southeast. The ages of these features are poorly constrained, but due to the obvious interaction with external water, they likely formed during the time of the extensive basin lake.

Towards the center of the basin, the basaltic features are dominantly tuff cones and tuff rings; whereas, towards the boundaries the features are dominated by maars and cinder cones (Heiken, 1971). The oldest formation identified in the area is the Picture Rock basalt, a 230 m thick sequence of 10 m thick basalt flows interbedded with various sandstones, conglomerates, and tuffaceous mudstones, which are interpreted as flood plain and/or shallow lake deposits (Walker et al., 1967). The Silver Lake graben, located immediately south of TRC, is the beginning of a 25 km wide structural arch that forms the southwest boundary of the lake basin (Heiken, 1971). Driller's logs show that ∼220 m of flat-lying lacustrine sediments and interbedded tuffs overlie the Picture Rock basalt formation at the center of the basin, and that these sediments thin to 0 m towards the basin boundary (Hampton, 1964). At the center of the basin (beneath TRC), these sediments consist primarily of diatomites, whereas closer to the basin margins they consist of coarse clastic sediments, lava flows, and volcanic breccias (Heiken, 1971). The underlying stratigraphy exposed on the west side of TRC consists of a 5–10 m thick basaltic lava flow, followed by ∼ 8 m of interbedded volcanic litharenites, lithic arkoses, and diatomaceous siltstone and mudstones (Heiken, 1971). In contrast, the non-volcanic stratigraphy beneath the eastern portion of TRC is dominated by wellbedded diatomites (Heiken, 1971). The interbedded, non-volcanic units exposed on the west side of the edifice are interpreted as outwash apron deposits from the Connley Hills to the northwest, a 6.4 km wide by 19-km-long volcanic feature consisting of a basaltic shield and intermediate to silicic domes, which likely formed an island in the Pliocene–Pleistocene basin lake (Heiken, 1971). 1.2. Table Rock Complex TRC is located ∼ 14 km east of Silver Lake, Oregon, on the shore of present day Silver Lake, which is likely the small remnant of the much larger lake that occupied the Fort Rock–Christmas Valley Basin in Pleistocene time (Heiken, 1971). TRC has an elongated-oval shape, trends NNW (along strike with the other phreatomagmatic complexes in the Fort Rock–Christmas Valley region), and covers an area approximately 8 km by 5 km (Heiken, 1971). Two large phreatomagmatic edifices; a large southern tuff cone with a capping solidified lava lake at 395 m above the basin floor (TRC1), and a low, broad tuff ring in the northeast (TRC2; Fig. 2), make up the complex. Additionally, seven smaller tuff rings and vents were identified along the flanks of the complex, yielding a complicated network of tuff ring–tuff cone deposits. For the purposes of this paper, only the two largest and most significant eruptions of TRC1 and TRC2 will be discussed in detail. The flank vents will be discussed briefly in terms of size, location, and cross-cutting relationships with other erupted tuffs. 2. Data Detailed geologic mapping and twenty-three stratigraphic sections were completed to determine the eruptive history of TRC (Fig. 2), resulting in the identification of 42 lithofacies based on variations in grain size, composition, and sedimentary structures (following Fisher, 1961; Schmid, 1981; Table 1). Detailed descriptions of each lithofacies are available as online supplementary data. These lithofacies were grouped into six Facies Associations (FA) according to common bedding styles, juvenile fragment morphology, and the percentage and type of accidental components (Table 2). Five of these FAs correspond to the TRC1 eruption, and one corresponds to the TRC2 eruption. Inferred eruptive conditions and depositional mechanism are discussed below in the text. Grain percentage was estimated in the field and supported by subsequent petrographic thin section analysis of each FA from multiple locations around the complex. In the field, juvenile clasts were distinguished from accidental basalt by their fresh appearance,

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Fig. 1. Study location (modified from Heiken, 1971). Upper right image is the state of Oregon (USA), and the gray shaded area is enlarged as the lower left image. The enlarged image shows the location of the Table Rock Complex, as well as many of the other hydromagmatic edifices and cinder cones within the basin.

irregular shapes, or twisted-fluidal shapes, greater vesicularity, and quenched glassy rinds with vesicular cores. Thirty thin sections, sampled from multiple lithofacies of the TRC1, TRC2, and flank vent eruptions, were analyzed and point counted to better constrain the composition and percentage of accidental versus juvenile clasts, especially in the matrix. For each thin section, 300 points were counted with ∼ 1 mm spacing between traverses and between points. Details of these analyses are available as online supplementary data, and a summary is presented below in Sections 2.2 and 2.4. 2.1. TRC1 Facies Associations: descriptions and interpretations The TRC1 deposits originated from the eruption that formed the large tuff cone with a capping solidified lava lake in the south-central part of the complex (Fig. 2). This was the first eruption in the sequence of tuffs, as it directly overlies the pre-existing lake and lacustrine sediments described above. The deposits are well exposed along the south, east, and north sides of the edifice, and intermittently exposed on the west side as they were eroded away or covered by later eruptions. The sequence is consistent around the vent with minor variations in depositional characteristics such as grain size and bedding thickness due to distance from source. The grains within TRC1 consist dominantly of juvenile basalt, accidental basalt, and white mudstone–siltstone lake sediments, which will be referred to as mudstone from here onward. Most of the deposits also contain abundant fine-to-medium ash matrix. Petrographic analysis of the matrix reveals that ∼98% of the matrix and grains within the thin sections are of juvenile origin (a detailed thin section analysis is presented in Section 2.3 below).

2.1.1. PH1: massive to stratified tuff and lapilli tuffs PH1 consists of alternating, palagonitized lithofacies, T1, LT1, and LT2 (Tables 1 and 2; Fig. 3). Early beds alternate from grain supported, graded, and generally stratified lapilli tuff (LT1, 0.1–1 m thick), and a massive, matrix supported (95–100% fine-medium grained ash) deposit with occasional imbricated pebble stringers of mudstone and rare accidental basalt fragments (T1, 0.5 to 2 m in thickness; Fig. 3a, c). Stratigraphically higher beds are well-stratified, wavy-planar bedded, and show reverse grading from coarse ash up to coarse lapilli-to-blocks (32 cm down to b6 cm thick with distance from source). The coarse lapilli to block-sized juvenile grains consist of scoriaceous, twisted, fluidal, and often flattened juvenile grains with quenched rinds (found both in LT2, and at random, concentrated horizons within LT1). These juvenile spatter-bomb clasts are not found above the PH1 horizon. Soft sediment deformation is prevalent throughout PH1. 2.1.2. PH1 interpretation T1 and LT1 are likely a combination of tephra fallout and subaqueous sediment gravity flows. The angular grains and moderate sorting within LT1 may be indicative of fallout, possibly through a density current (Valentine and Giannetti, 1995), whereas the lenticular interbeds, erosive contacts, and reverse grading often found within LT1 are consistent with a deposition by lateral movement of grains (White, 1996, 2000). The fine-grained, massive, poorly sorted, and non-graded deposits of T1 suggest deposition from a concentrated suspension density current with little tractional transport, which inhibits the development of grading (Sparks, 1976; Sparks et al., 1978; Chough and Sohn, 1990; Freundt and Bursik, 1998). The reverse-graded, coarse ash to fine block deposits of LT2 are interpreted

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Fig. 2. Geologic map (modified from Heiken, 1971) overlain on topographic map of TRC.

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Table 1 Lithofacies of both the TRC1 and TRC2 eruptive products, as well as the non-volcanic products (NV), classified on the basis of grain size/abundance, sedimentary features, and dominance of juvenile fragments [table modeled after Nemeth et al. (2001) and Nemeth and White (2003)]. Detailed descriptions of individual lithofacies are available as online supplementary data Lithofacies

Tuff Breccia (TB)

Lapilli Tuff (LT)

Tuff (T)

Magmatic (M)

Pre-TRC1 eruption TRC1 eruption Clast supported Massive Massive-to-diffuse stratification Stratified Stratified, reverse grading Scour-fill massive Scour-fill bedded Cross-stratified Matrix supported Massive-to-diffuse stratification Diffusely stratified Stratified, laminated Scour-fill bedded Planar-to-wavy beds Strombolian deposits, lava lake, and radial dikes

NV1

LT4, LT6 LT3, LT5 LT1 LT2 TB1 LT8, LT9 LT18 T1 T18 T2, T3 LT7 T17 M1

Post-TRC1 eruption TRC2 eruption Clast supported Massive, non-graded Massive, reverse grading Massive-to-diffusely stratified bed in a well-stratified deposit sequence Stratified, reverse grading Scour-fill massive Planar-to-wavy stratification, occasional lenticular deposits Wavy-to-cross-stratified Matrix supported Massive Massive-to-diffusely stratified bed in a well-stratified deposit sequence Stratified — pinch and swell Stratified Stratified, laminated Scour-fill bedded Wavy-to-cross stratified

Non-volcanic (NV)

NV2

TB2

as a combination of fallout and density current deposits. The coarse deposits of LT2 are interpreted as having an initial fallout origin that, after landing, transitioned into thin, laterally moving, traction carpet dominated, subaqueous sediment gravity flows. The finely-laminated, well-sorted, normal-graded ash beds that overlie the coarser deposits are interpreted to have been deposited by ash-fallout settling through a column of water. These beds are typically highly deformed under the weight of the overlying material, further suggesting that they were unconsolidated and water-saturated during emplacement of subsequent units (Nichols et al., 1994; White, 2000). The field and thin section analyses show that N98% of the grains are juvenile and typically angular to subrounded, indicating minor to no abrasion (see Section 2.2 below; also see online supplementary data). The angular, blocky glass shards indicate magma–water interaction as the dominant fragmentation mechanism. However, the fluidal, scoriaceous juvenile coarse lapilli-to-fine blocks suggest that, in the early stages of the eruption, both hydromagmatic and magmatic processes were occurring simultaneously (Houghton and Wilson, 1989; Cole, 1991; Houghton et al., 1999). 2.1.3. PH2: normal graded, massive to diffusely stratified sequences The PH2 facies association consist of repeating packages of lithofacies (i.e., LT6–LT5–T3 and LT4–LT3–T2; Tables 1 and 2). Each package begins with a coarse-grained (angular to subangular grains of fine-to-coarse lapilli), massive to diffusely-stratified, poorly-to-moderately well sorted, grain-supported deposit with variable thickness (0.25 up to 1.5 m thick; LT6, LT4). Both normal and reverse grading is common within the same bed (Fig. 4), however the package as a whole gradually fines upwards (LT3,

LT17 LT10 LT14, LT15, LT16 LT12 LT11

T13, T15 T21

T20 T10 T9, T16 T4, T5 T12 T11, T14, T6, T7, T8 T19

LT5), and is capped by a planar-to-cross laminated, medium-to-coarse ash (grains subangular to subrounded) bed that commonly displays soft sediment deformation (20–40 cm thick, T2, T3; Fig. 4). Individual lithofacies pinch and swell laterally, but the sequences overall are laterally continuous for 10's up to 100's of meters away from source. Field and petrographic analysis show that the grains in all of these deposits are dominantly juvenile. Accidental grains of mudstone and dense, angular basalt compose b5% of the deposits. However, many of the block-sized ballistic clasts with underlying sags are composed of accidental basalt. 2.1.4. PH2 interpretation The repeated, fining upward packages of PH2 are interpreted to correspond to the traction carpet and suspension sedimentation stages of high-density, turbulent, subaqueous sediment gravity flows (Nemec et al., 1980; Lowe, 1982; Postma, 1986). Lithofacies LT6–LT5 and LT4–LT3 are thick, diffusely stratified deposits with alternating coarsening- and fining-upward sequences. They have no distinct bedding surfaces, and have an overall fining upward trend. Changes in grain size throughout a given deposit reflect the grain size variation of supplied sediment with time, and/or intermittent waxing and waning flow conditions (Sohn, 1997). The overall fining upward sequence is interpreted to be a consequence of waning flow conditions and loss of sediment supply over the duration of the current. The fine-grained, stratified lithofacies (T2, T3), which overlie the coarser, thicker deposits, represent the last stage of sedimentation from the collapsing turbulent, suspended-load region. As with the deposits of PH1, grains from PH2 are subangular to subrounded, and consist dominantly of juvenile volcanic clasts. This

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Table 2 The lithofacies have been grouped into Facies Associations (FA) according to common occurrences, depositional features, and compositional similarities. These are discussed in detail in the text. PH = Phreatomagmatic and M = Magmatic Facies Associations description (TRC1)

Facies Associations description (TRC2)

PH1: Massive to stratified tuff and lapilli tuffs

PH5: Strata within large dune-forms.

T1, LT1, LT2

PH5a — Repeating and interbedded lithofacies found just above the contact with the TRC1 deposits on the north side of the complex.

PH2: Normal graded, massive-to-diffusely stratified sequences with occasional zones of inverse grading

T4, T5, LT10, LT11

LT4–LT3–T2, LT6–LT5–T3

PH5b — Repeating lithofacies found on the east side of the crater.

PH3: Scour and fill deposits

T9, T10, T11, T12, T13, LT13

TB1, LT7, LT8, LT9

PH5c — Interbedded, planar-to-cross stratified lithofacies that dip into the TRC2 crater on east side.

PH4: Cross-bedded deposits T14, T15, LT14, LT15, LT16 T17, T18, LT18 M1: Spatter, lava lake deposit, and radially intruding dikes within tuff cone walls of TRC1.

PH5d — These are interbedded and alternating beds found in the southern part of the complex. T19, T20, T21, T22

M1 PH5e — Found filling in large scours— likely closely related to first 20 m of the PH5a deposits. LT12, T6, T7, T8 PH5f — Massive tuffs, lapilli tuffs, and tuff breccias found within the large dune forms throughout the east and southeast quadrants TB2, T16, LT17

observation, combined with the lack of highly abraded and interbedded non-volcanic clasts and sediments, suggests that the density currents were eruption-fed, and that grain size and morphology of the clasts reflect eruptive conditions at the vent (White, 2000). Additional coarse tephra dispersed through the water column, either by fallout from tephra jets or from dilute density currents traveling across the surface of the water, may have caused ash and lapilli to be continuously rained into the subaqueous flows (White, 2000). This process may have suppressed the formation of distinct thin layering (Lowe, 1988; Arnott and Hand, 1989; White, 2000), and could further explain the thick massive nature of LT6–LT5 and LT4–LT3 (Fig. 4). In several areas N1 km from the vent, up to 50 cm of smectitic lake sediments overlie the deposits of PH2 (representing the end of the TRC1 eruptive sequence distal from source). This observation, combined with the lack of subaerial features such as base surge deposits and well-sorted strata from fallout, further supports the subaqueous depositional environment, and suggests the lake existed well after the eruption ceased. 2.1.5. PH3: scour and fill deposits PH3 consists of channel-shaped, scour-and-fill deposits within the FAs of PH1 and PH2. The lower contact is invariably erosive into underlying substrate, and is filled with a variety of lithofacies including massive, poorly sorted tuff breccias (TB1), and various lapilli tuffs that fill in large scour features (LT7, LT8, and LT9; see Table 1 for description of these lithofacies; also see online supplementary data). Pervasive soft sediment deformation such as sags and flame structures are common, and fine-grained beds on the channel walls are often deformed into small convolute folds towards the axis of the channel. 2.1.6. PH3 interpretation The scour and fill features of PH3 are interpreted to be the result of erosive subaqueous sediment gravity flows, and are primarily found N1 km from the vent on the flanks of the growing tuff cone platform. The highly deformed and occasionally folded beds

that fill the large scours suggest that the sediment gravity flows were saturated, and therefore likely emplaced subaqueously and deformed by the weight of the subsequently deposited overlying material. The sediment gravity flows could have originated from one of two sources: (1) remobilization of tephra avalanching down the steepening slopes of the outer tuff cone as large quantities of new sediment were added to the system from the erupting vent; or (2) high-concentration, eruption-fed density currents derived from column collapse. The fine strata that fill the scour features are likely the deposits of the waning sediment gravity flows, or fill from subsequent currents and fallout. 2.1.7. PH4: cross-bedded deposits PH4 was found only proximal to the TRC1 tuff cone, above 1451 m. PH4 deposits more distal from the tuff cone are not exposed, and either were eroded, or never deposited. PH4 deposits are dominated by wavy- to cross-stratified tuffs and lapilli-tuff beds (T17, T18, and LT18; Tables 1 and 2; also see online supplementary data). Dunes are symmetrical, have wavelengths from 10–20 m parallel to the direction of flow, and amplitudes of 0.5–1 m. Individual beds within a single dune consist of one of the aforementioned lithofacies (e.g., a dune with a 1 m amplitude may be composed of multiple, 10–20 cm layers of T17). The PH4 FA beds dip inwards towards the vent at roughly 5–8°, similar to the dip of the underlying PH2 and PH1 FAs. The contact between PH2 and PH4 is not exposed, but is constrained to within 10 vertical meters. 2.1.8. PH4 interpretation Based on the cross-stratified nature of these deposits, PH4 is interpreted to be the result of dilute density currents flowing across a subaerial, gently-sloping platform. These were determined to originate from the TRC1 eruption based on location and dip inwards towards the TRC1 vent, and consistency in dip with the underlying TRC1 deposits. The presence of preserved base surge deposits suggests that in this part of the stratigraphy, tephra was being deposited above the level of the lake. The stratigraphy therefore indicates that the level

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Fig. 3. a) Generalized stratigraphic column representing PH1 (composed of lithofacies T1, LT1, and LT2). The dark blobs represent the spatter-bomb clasts which range from medium lapilli to fine blocks depending on distance from source (as described in text); b) LT1; c) T1 with lenses of LT1; d) LT2 (photographs a–c taken from section W4, 2.4 km from source, pencil for scale); e) LT2 (photograph taken in section SW1, 0.6 km from source).

of the lake at the time of the TRC1 eruption was approximately 1450 m above sea level. 2.1.9. M1: spatter deposits and solidified lava lake The last FA identified in this sequence is that of black and red scoriaceous-to-spatter deposits overlain by a cap of flat-lying, gray, aphanitic, high-alumina basalt (Heiken, 1971). M1 (M for magmatic) is found at the highest point in the complex, at 395 m above the basin (Fig. 2). Two dikes, one trending north and one south–southeast, extend from this point.

2.1.10. M1 interpretation Over the duration of the TRC1 eruption a large, symmetrical tuff cone was built above the surface of the lake water. This cone as seen today is ∼1530 m in diameter at the base, and ∼360 m in diameter at the top. M1 is interpreted to represent a final, magmatic stage of the TRC1 eruption, in which the vent of the growing tuff cone was gradually isolated from interaction with external water, resulting in a fire-fountaining, Strombolian stage (cinders and spatter deposits), and a final effusive stage which formed a crater filling lava lake. The dikes appear to radiate out

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Fig. 4. (a) Generalized stratigraphic column representing PH2; (b) LT4–LT3–T2 sequence, ∼2.8 km from source; (c) LT6–LT5–T3 sequence, 3.6 km from source (section N4, person 1.5 m for scale); (d) Closer view of LT6–LT5–T3 sequence, 1.6 km from source (section W4, hammer for scale).

from the lake and likely intruded the unconsolidated tuff cone walls, which have now been eroded away. 2.2. TRC1 thin section Petrographic analyses of samples from the TRC1 eruption (including all FAs) show that the grains are composed dominantly of palagonitized glass shards of varying sizes. Some glass has elongated, swirly, and fluidal textures, but most are angular, fractured, and blocky with little evidence of rounding and abrasion. Many of the larger glass shards (N300 µm) have tiny, round, thick-walled vesicles, which are on average 20 to 80 µm in diameter, but can be found up to 240 µm in diameter. These vesicles compose 1% to 26% of the glass grain; however, coalesced bubbles are rare. Subhedral to euhedral phenocrysts of plagioclase, orthopyroxene, and olivine are common within the glass fragments, and are also found broken and occa-

sionally abraded within the fine ash matrix. The altered, glassy, fineash matrix contains features similar to mud cracks (on the micron scale). Accidental material consists of (1) subrounded to rounded, crystalline basalt lava flow clasts with a plagioclase-rich groundmass in which the elongated microlites and phenocrysts are aligned; (2) rounded, highly weathered mudstone clasts, which in some locations contain siltsized particles of altered glass in the matrix; and (3) silicic pumice with elongated, stretched, thin vesicles walls. The silicic pumice are identical in composition and micro-scale texture to samples of the NV1 deposits just below PH1 of TRC1 (Table 1). On average, samples from TRC1 contain less than 5% accidental clasts (detailed petrographic analysis available as online supplementary data). The subaerial deposits (PH4) contain 1–4% armored and accretionary lapilli, but overall the components do not vary significantly from the beginning to the end of the pyroclastic sequence. Details of the spatter and lava flow deposit can be found as online

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Table 3 This table shows the range of dune wavelengths (λ) and average maximum dips of PH5 both parallel and perpendicular to the direction of flow around the TRC2 tuff ring. The maximum dip refers to measurements taken just below the crest of a dune W–NW

N–NE

E

SE

Orientation relative to flow direction

λ (m)

Average dip (°)

λ (m)

Average dip (°)

λ (m)

Average dip (°)

λ (m)

Average dip (°)

Perpendicular Parallel

80–300 m 80–150 m

19° 30°

80–200 m 80–120 m

10° 33°

300–500 m 20–200 m

19° 27°

− 20–100 m

– 28°

supplementary data, and compositional data can be found in Heiken et al. (1981). 2.3. TRC2 Facies Associations: descriptions and interpretations 2.3.1. PH5: long-wavelength dune bedforms The TRC2 deposits originated from the eruption that formed the large depression and tuff ring in the north–northeast part of the complex (Fig. 2). The TRC2 eruption has only one FA, PH5, which

contains 26 lithofacies (Tables 1 and 2; also see online supplementary data). There are two scales on which the deposits of TRC2 must be described. The first focuses on the large features (tens to hundreds of meters), and the second focuses on individual strata (centimeters to decimeters). On the large scale, the TRC2 deposits consist of long-wavelength, three-dimensional, symmetrical dune structures with wavelengths that vary from 20 to 200 m in the direction parallel to flow and from 80 to 500 m perpendicular to flow. Table 3 presents the range of dune

Fig. 5. (a) TRC2 dune form above section W4 in the western TRC. Flow direction is oblique to the plane of the photograph; (b) Large lobate feature in the TRC2 surge deposits. The photo was taken from an adjacent ridge to the south, and the flow direction was from right to left. Note that the strata dip to the west (left), shallow in a trough, and then dip to the east (right). Person 1.5 m circled for scale in (a) and (b); (c) TRC2 dune in the eastern sector of TRC. The crest of this dune is on the right side of the photograph. To the right and left of the crest, the strata dip away from the crest. This dune is parallel to flow direction (flow direction from left to right), and is close to 200 m in length; (d) trace of dune in (c).

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Fig. 6. Small scale PH5 features from the TRC2 eruption (a) massive sandy beds with gas (steam) escape pipes; (b) alternating lapilli and tuff beds, dark clasts are juvenile; (c) antidune; (d) laterally continuous beds that slightly pinch and swell; (e) laterally continuous beds with some cross stratification at base; (f) low angled, low amplitude dune bed. Hammer for scale in (a)–(c), (e), (f); 10 cm tall notebook for scale in (d).

wavelengths and average dips just below the crest for four regions around the TRC2 tuff ring (west–northwest, north–northeast, east, and southeast). Wavelengths in the east tend to exceed those in other regions. Dips parallel to flow tend to exceed dips perpendicular to flow (30° vs. 10°–20°), with similar values in all regions. Dune wavelengths in the direction perpendicular to flow exceed wavelengths in the direction parallel to flow, consistent with the dip data. The amplitude of the dunes are consistently approximately 1/10th of the wavelength parallel to the direction of flow (Fig. 5). The dips shallow towards the middle and tops of the waveforms to b1°, often forming topographic saddles and flat crests. Similar features were noted at El Chichon and described as transverse, sinuous ridges up to 100 m in length with horizontal form index (breadth/wavelength, or in our case perpendicular wavelength/parallel wavelength) ranging from 1 to 10 (Sigurdsson et al., 1987). The horizontal form index for the TRC2 dunes ranges from 1 to 5, with the longest perpendicular flow direction wavelength equal to 500 m. The difference between the TRC2

and El Chichon dunes is that the TRC2 dunes are symmetric with dips on the lee and stoss sides of 27–33°, whereas the El Chichon dunes are asymmetrical with steep stoss and gently sloping (20° dip) lee sides. The deposits of PH5 also include features such as (1) large U-shaped channels (10–32 m deep; Heiken, 1971; Heiken et al., 1981) where the steeply dipping deposits along the channel walls are plastically deformed, slumped and folded towards the channel axis (PH5e; Tables 1 and 2, also see online supplementary data); (2) Largescale chute-and-pool features (up to 13 m tall); and (3) deposits plastered up and around pre-existing obstacles. Also, each set of dunes truncate pre-existing deposits and/or are truncated by later deposits, which results in large-scale hummocky cross-stratification. These features will be discussed in more detail in Section 3.3 below. On the small scale, each dune consists of a range of lithofacies, the most dominant being centimeter to decimeter thick, laterally continuous, wavy-planar strata that are internally massive, poorly to moderately well-sorted tuff and lapilli tuff beds. The large dunes also

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Fig. 7. Measured stratigraphic sections, from right (proximal to vent) to left (distal from vent), along with FAs interpreted from the section.

commonly contain smaller, isolated, sporadic, meter-scale antidunes (Fig. 6). In general PH5 deposits are dominated by fine- to medium-grained ash, and contain abundant accretionary and armored lapilli, vesiculated tuff, and pervasive soft sediment deformation. The grains consist dominantly of juvenile scoria, but also contain a higher percentage of accidental clasts than those of TRC1 (35–55%). Accidental lithics are dominated by platy, aphanitic basalt, and to a lesser degree by angular, vesicular basalt and mudstone clasts (same as found in deposits of TRC1). Grain types includes 10–40% accidental basalt, 5–15% mudstone, and 45–85% juvenile clasts and matrix. Where exposed, the base of PH5 has a sharp and erosive contact with underlying TRC1 deposits. However, in a few areas up to 50 cm of diatomitic lake sediments exist between TRC1 and TRC2 deposits, at the contact between FAs PH2 and PH5. Facies Association PH5 consists of several subgroups (PH5a–PH5f; Tables 1 and 2, online supplementary data). 2.3.2. PH5 interpretation PH5 is interpreted to be the result of multiple, high velocity, column collapse-induced, dilute pyroclastic density currents (e.g., Fisher and Waters, 1970; Schmincke et al., 1973). These currents, also known as pyroclastic base surges, are unsteady, density-stratified gas-particulate currents where turbulence is the dominant particle transport mechanism (Fisher and Waters, 1970; Crowe and Fisher, 1973; Sigurdsson et al.,

1987; Valentine, 1987; Druitt, 1992; Wohletz, 1998). The variation in depositional characteristics from one bed to the next within the large dune features (i.e., fines rich, fines poor, massive, stratified, inverse graded, non-graded), and lack of distinct or discernible vertical bedding patterns between various internal strata, reflect spatial and temporal variations in bed load flow dynamics as the deposits aggraded vertically. Deposits consisting of laterally continuous beds with grain alignment likely reflect laminar flow conditions in the depositional region of the current (Druitt, 1992), whereas deposits consisting of massive deposits indicate inertial grain flow (Wohletz and Sheridan, 1979). Deposits consisting of thin inversely graded strata suggest traction carpet transport in the depositing layer of the bed load (Lowe, 1982; Sohn, 1997), whereas cross-stratified dunes and antidunes indicate turbulent flow during sedimentation (Valentine, 1987; Druitt, 1992), or energetic, turbulent sweeps through a basal granular fluid, as suggested by Brown et al. (2007). Therefore, variations in small-scale bedform features are interpreted as a consequence of temporal unsteadiness in the flow, similar to the interpretation for isolated dune features within the distal Mt St Helens blast deposits (Druitt, 1992). Where present, the small-scale dune features within the larger TRC2 bedforms are dominated by antidunes, which suggest that the flow velocity exceeded internal wave speed in the bed load at the time of deposition, and that standing internal waves developed within the stratified flow (Crowe and Fisher, 1973; Hand, 1974; Allen, 1982; Valentine, 1987). The same argument could be made for the larger,

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Fig. 8. (a) Deformed clastic and lacustrine lake sediments (NV1) beneath the TRC1 pyroclastic deposits; (b) NV1 (base) and overlying, massive, sandy T1 deposits of TRC1. Note the rectangular rip-up clasts of lacustrine substrate in the tuff deposits (NV1; Brunton for scale); (c) Large clastic dike intruding the TRC1 deposits. Person 1.5 m for scale (a) and (c).

20–200 m wavelength dune forms, as their symmetrical nature also implies standing waves within the broader context of the overall bed load, and therefore possibly supercritical flow conditions. In this regard, we interpret the larger dune forms (20–200 m wavelengths) in the TRC2 deposits to reflect internal waves in the thick bed load region of the greater flow, and the smaller-scale features as the result of layerby-layer deposition and variations in internal waves and flow dynamics within the lowermost depositional bed load region. The average clast size most proximal to the vent (b1 km) is less than 0.5 cm. However, the particle sizes increase to an average grain size of 1–5 cm with distance from source, suggesting maintained high flow velocities and current competence with distance from source up to 1.5 km, and thereafter declining current competence up to 3 km distance where the average grain size again decreases to b1 cm. Abraded, block-sized clasts are found intermittently at bedding horizons, and are interpreted as ballistic clasts that were entrained into the flow after impact, as it is unlikely that block-sized clasts remained in suspension over the distance of the current. Their abraded nature also suggests that the block-sized clasts were tumbled or bounced along via saltation at the base of the flow. Other than the large-scale dune forms, additional features which attest to the high velocity of the base surges include deep U-shaped channels in the northeastern part of the complex, the 13 m tall chuteand-pool feature observed in TRC2, and evidence for currents surmounting and depositing across 21 to 45 m tall obstacles N2 km from source (discussed further in Section 3.3). The presence of accretionary lapilli, vesiculated tuff, steam-escape structures, plastered beds, and pervasive soft sediment deformation suggests that

these currents contained a significant amount of liquid water during transport and deposition (Wohletz and Sheridan, 1979; Allen, 1982; Wohletz, 1998). 2.4. TRC2 thin section analysis Twelve samples were taken from various strata throughout the PH5 sequences (detailed descriptions of these samples are available as online supplementary data). Petrographic analysis shows that grain morphology within the matrix is consistent in each of the samples, and varies only in grain size depending on the coarseness of the deposits they were collected from. Grains consist of subangular to rounded, clear-brown to dark and sub-opaque (altered) brown glass with 2–25% rounded vesicles. The glass shards are blocky and have fine fractures running across their surfaces. Vesicles within the glass shards are rounded, have thick bubble walls, and range from 20–100 µm in size. Plagioclase, orthopyroxene, and olivine are also present as individual grains in the matrix. They are also angular to subrounded, except when found as phenocrysts within the glass grains where they are subhedral to euhedral. The only new grain within these deposits (i.e., not found in the TRC1 samples) is a dark, dense, most commonly irregularly shaped, but occasionally subrounded juvenile grain. It differs from the glass grains in that it contains 10–18% plagioclase needles that are dispersed throughout the sample, rather than found as radiating clusters. The dense, altered matrix of the new juvenile grain, which contains 2–5%, and rarely up to 25% rounded vesicles, composes the rest of the juvenile

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clast. Accidental material within PH5 consists of the same subrounded to rounded crystalline basalt, rounded, highly-weathered mudstone, and silicic pumice that were found in the TRC1 deposits. Most grains, especially those larger than 300 µm, have a thin fine ash coating. The grains within these samples have very similar textures and morphology to the TRC1 eruptive products; therefore, it is not possible to tell if the glass shards in these samples were derived from juvenile material from the TRC2 eruption, or entrained from the deposits of TRC1.

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3. Distribution of Facies Associations The geologic map (Fig. 2) shows the distribution of the TRC1, TRC2, and flank vent deposits around the complex. The dashed circles with speckled fill represent the approximate locations of the various vents. These are surrounded by dotted lines which represent the approximate location of the inferred crater rims. Numerous stratigraphic sections were measured in four sectors around the vent (West, North, East, and Southern arm) in order to reconstruct the temporal evolution of the eruption as well as visualize facies variations with distance from source.

Fig. 9. Measured stratigraphic sections of N4 and NE1 (Fig. 1 and 10) and associated FAs. N4, the western-most measured section is on the right, and the eastern-most on the left.

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3.1. Western TRC exposures Four vertical sections were measured in the western sector of the volcano; SW1 at 0.6 km from TRC1 source, W4 at 2.4 km from TRC1 source, W8 at 2.6 km from TRC1 source, and W10 at 3.4 km from TRC1 source (Fig. 7d, c, b, and a, respectively). The sections reveal a progression from less than 10 m of PH1 (first set of eruption-fed turbidity current deposits) to a few tens of meters of PH2 (second set of eruption-fed turbidity current deposits). PH3 (TRC1 scour and fill deposits) appears in only one section, near the base of section W8, but was traced to more distal locations in the field. PH4 (TRC1 surge deposits) is found only proximal to source, and was likely eroded away in other locations, replaced by the later deposits of TRC2-PH5, or never deposited in the first place. The stratigraphic sequence (other than in areas proximal to the TRC1 tuff cone) is overlain by up to 45 m of PH5 (TRC2 surge deposits). Given the location of the subaqueous– subaerial depositional transition within the TRC1 eruptive products, and the basal contact of the TRC1 pyroclastic deposits and older lake sediments, the paleo-lake at the site of the TRC1 vent was roughly 60–70 m deep. The contact between the lake sediments and pyroclastic deposits is exposed in sections W8 and W10 (Fig. 7a, b). The sediments dip inwards beneath the pyroclastic section at 30–40° (Fig. 8a), and form clastic dikes and flame structures into the overlying pyroclastic sediments of TRC1 (Fig. 8c). Thin (b3 cm), interbedded, lenses of NV1 lake sediments or large (up to 40 cm long), intact pieces of the NV1 substrate (Fig. 8b) are commonly found at the base of the pyroclastic section. 3.2. Northern TRC exposures Two sections, N4, 3.6 km from the TRC1 source, and NE1, 3.5 km from the TRC1 source (Figs. 2 and 9), were measured on the northern side of the complex. Both sections begin much lower in elevation than the western sections, suggesting a deepening of the lake to the north. The two most interesting features are the increased number of ballistics in the N4 section, and the finer grain size (average mediumto-coarse ash) in the NE1 section. The ballistic clasts in the N4 section are on average 0.3–0.5 m in diameter, but occasionally up to ∼ 1 m, and all have deep bomb sags (e.g., at base of the stratigraphic column, a 97 × 66 cm accidental clast of basalt deforms the underlying strata N1.5 m, Fig. 9b). Ballistic clasts comprise 5–7% of clasts, are found randomly scattered throughout the deposits and are present in much higher proportions than in more proximal sections. Dewatering features such as squeeze-up and flame structures (0.1–1 m) and discontinuous offsets and faults (usually N1 m in length) are also common in this section. A 40 cm thickness of NV2, the non-volcanic lithofacies between the TRC1 and TRC2 deposits, was found about 80 m southwest of the N4 outcrop. Fig. 10 is a schematic fence diagram for the north face of the complex over a cross section shown in Fig. 2. Stars indicate where the cross section bends (Fig. 10). The fence diagram extends from east (left) to west (right; Fig. 10a). On the east side of the complex, the contact between NV1 and PH1 is exposed at lower elevations (approximately 1326 m a.s.l.), suggesting that the lake deepened to the east. The TRC1 deposits are grouped in Fig. 10 to illustrate their lateral extent and relationship with the overlying TRC2 deposits. The contact between the TRC2 and TRC1 deposits is irregular and varies in elevation, which we attribute mostly to scouring by TRC2 surges. However, NV2 lake sediments were found in several locations in the west (40 cm thick near section N4), and in the east (east of section NE1), indicating that the TRC1–TRC2 contact is not erosive in all locations. Three smaller flank vents, FV 1, FV 2, and FV 3 (Fig. 10b and c; first recognized by Heiken, 1971), cut through the TRC1 deposits but are overlain by the deposits of TRC2. This indicates that the flank eruptions occurred after the TRC1 event, but before the TRC2 event. The crater of FV 1 is ∼ 180 m in diameter with a surrounding tuff ring

∼250 m in diameter; and the crater of FV 2 is ∼190 m in diameter with a surrounding tuff ring ∼ 440 m in diameter. The lateral tuff ring deposits have been eroded away, and the exposed outer tuff ring deposits are highly weathered and form inaccessible cliffs making detailed stratigraphic sections impossible. FV 3, previously named vent 8 by Heiken (1971), is also located in the northern section. It is a small flank vent, only 100 m at the base and 200 m in diameter at the top, and is well exposed in the cliffs of the northern face. Three large U-shaped channels were identified in the northeastern side of the complex (location shown by the three arrows in Fig. 2; also first recognized by Heiken, 1971). The channels, which scour the TRC1 deposits, are located close to the measured section of NE1. The upstream side of the westernmost channel is 7.5 m tall, 6.75 m wide at the base, and 13 m wide at the top (Fig. 11a, b). The downstream side, which is roughly 30 m further from source, is 13 m tall, 10 m wide at the base, and 35 m wide at the top (Fig. 11d). Similar increases in channel dimensions with distance from source have been noted at Koko crater, HI (Fisher, 1977) and Barcena Volcano, Mexico (Richards, 1959). Near the west channel, parallel to the flow direction the strata are observed to plaster up and over the pre-existing TRC1 deposits (Fig. 11c), and drape the obstruction perpendicular to direction of flow (Fig. 11a). 3.3. Eastern TRC complex-PH5 The eastern flank of the complex is dominated by the surge deposits of TRC2. The TRC1 deposits are poorly exposed in one small area in the east, and another in the southeast (Fig. 2). TRC1 deposits were either eroded by the base surges of the TRC2 eruption, or were eroded by non-volcanic processes prior to the TRC2 eruption. The hummocky topography created by the PH5 deposits is most obvious in the eastern side of the complex. A well exposed outcrop in the east, beginning with the diamond labeled E4 on Fig. 2, and extending 500 m to the south, exposes ∼ 40 vertical meters of section in the direction perpendicular to flow (Fig. 12a). The corresponding 500 m long dune truncates pre-existing TRC2 deposits on the north (right) side (Fig. 12c); and is truncated by another large dune on the south (left) side (Fig. 12b). Fig. 12d represents the direction parallel to flow, and shows two 40–50 m long dunes that are part of the larger feature. These dunes dip ∼13° to the southeast, consistent with the dip shown in Fig. 12c (crest of dune on right side of photograph). These photographs illustrate the three-dimensionality of the dune forms. What is interpreted as a large chute-and-pool structure, first mentioned in Section 2.3.1 above, is found at location E5 (Fig. 2). This exposure is 1.6 km from the source, extends 40 m parallel to the direction of flow, and is 13 m tall (Fig. 13). The deposits parallel to the flow direction contain wavy-planar and horizontal strata for the first ∼ 35 m of the outcrop. These strata range from 10–30 cm thick, and are composed of alternating beds of matrix supported (up to 100% fine ash beds) and poorly sorted, grain supported beds. The strata begin to bend upwards at angles of 33 to 44°, where the beds abruptly thin and fine to an average 2–7 cm thick (Fig. 13). The strata continue to steepen and thin towards the downstream flow direction to the full outcrop height of 13 m before they begin bending back towards horizontal (Fig.13). The rest of the downstream side of this feature has been eroded away. The exposed flat-lying strata are 4 m thick on the left side of the feature (but are probably closer to 6 m in thickness given the exposed strata a few meters further to the east), and are overlain by slightly thicker and somewhat more diffusely stratified deposits which also cover the steeply dipping strata to the right. Although poorly exposed, the upward bending strata can be traced N100 m north of this feature, at the same elevation and distance from vent. Chute-and-pool features are commonly observed in base-surge deposits, and represent pffiffiffiffiffiffi a hydraulic jump where supercritical flow (chute, Fr N1,Fr = V= gh, where V=velocity, g=gravitational acceleration, and h=depth of flow) abruptly changes to subcritical flow

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Fig. 10. (a) North fence diagram. The location of sections N4 and NE1 are designated by long rectangles; (b) Flank vent 3. The vent margins are outlined with a thick white line, the strata with thin white lines. The top of FV 3 is 200 m across, and the base 100 m; (c) Flanks vents 1 and 2. Again, the margins are outlined with a thick white line. The remnant tuff ring around FV 1 is ∼ 250 m in diameter, and the remnant tuff ring around FV 2 is ∼440 m.

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Fig. 11. (a) Western channel from the south side of the outcrop (most proximal to vent); (b) Closer view of the western channel, person 1.5 m for scale(flow direction into page for a and b); (c) Plastered TRC2 surge deposits riding up and over a pre-existing obstacle (TRC1 tuffs, flow direction from right to left); (d) North side of outcrop (more distal from vent). Person on right ∼2 m tall for scale (flow direction out of page).

(pool, Fr b1; e.g., Schmincke et al., 1973). However, they are typically much smaller in scale (meters rather than 10s of meters). Similar large chute-and-pool features were identified in the 1991 Pinatubo deposits and at the base of Mt St Helens in the 1980 pyroclastic flow deposits, but have since been eroded away (Steve Self, personal communication). However, structures as large as this have never before been identified in basaltic hydromagmatic eruptions. An alternative interpretation for this feature is a surge current that surmounted a pre-existing obstacle that is no longer exposed, although

there is no evidence of any obstruction in this region (i.e., no remnant tuff rings or older deposits). Furthermore, given that the morphology of the feature is the same as previously identified chute-and-pool deposits (Jopling and Richardson,1966; Fisher and Waters,1970; Schmincke et al., 1973) and that the ratio of deposit thickness (downstream to upstream side of the jump, 2.2) is similar to other well documented chute-andpool features (Schmincke et al.,1973; Weirich, 1988), the chute-and-pool interpretation is preferred. This indicates supercritical flow conditions up to 1.6 km from source.

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Fig. 12. (a) 500-m-long dune form, exposed in the direction perpendicular to flow; (b) and (c) Closer view of the crests of the feature with the strata outlined; (d) Direction parallel to flow, located on the north side of the larger feature (see c for location marked (d) illustrating the three-dimensionality of this dune form; (e), (f), and (g) Closer views of the strata at the crest of the dune in (d) (location E4 in Fig. 2).

3.4. Southern arm of TRC complex The long arm that extends towards the south–southeast of the complex is composed primarily of bedded tephras from one or more of the surrounding flank vents, and is overlain by the deposits of PH5. At least four obvious flank vents exist in this area (FV 4–7, Fig. 2). Flank vents 4 and 5 are located in the west–southwest part of the complex (Fig. 2) and represent two highly eroded inner craters, which are partially nested within each other.

Strike and dip data suggest that another sizeable crater may have existed in the southern part of the complex (FV 6; Fig. 2), but the deposits have since been eroded away beyond confident recognition. Based on distinct differences in accidental components and depositional characteristics, the first 15 m of strata along the southern arm were determined to have originated from vents other than TRC1 or TRC2, and are likely the products of either FV 6 or 7. The upper ∼60 m of bedded tephra deposits in the southern arm consist of strata from FA PH5 of the TRC2 eruption, and contains dunes 20 m to 80 m in

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Fig. 13. This series of photographs were taken from location E5 (Fig. 2). (a) Large chute-and-pool feature. Outcrop is 40 m in length, and 13 m tall (person for scale); (b) closer view of the strata steeply bending upwards; (c) Feature with strata outlined. The six-meter height within the chute regions was extrapolated from deposits exposed to the east (right). Everything further east of the outcrop has been eroded away.

wavelength parallel to the direction of flow. At the southern tip of the arm, the surge deposits ramp up and over a pre-existing high in the topography (likely the remnant of an older crater), and have nearly vertical dips as they plaster against the 21 m high obstacle.

the time of the TRC1 eruption, and deformed under the weight of the rapidly accumulating pyroclastic deposits (Heiken, 1971). The clastic dikes likely represent dewatering features that occurred both due to the overlying load and volcanic seismicity (Nichols et al., 1994).

4. Discussion 4.1. TRC1 TRC1, the first eruption to take place at the Table Rock Complex, initiated with a Surtseyan-style eruption though a 60–70 m deep, fresh water lake, as suggested by the distance between the freshwater lake sediments and the subaqueous–subaerial pyroclastic contact. Where exposed, the NV1 lake sediments dip toward the vent beneath the pyroclastic section at 10–30°, soft sediment deformation structures are pervasive, and clastic dikes varying from decimeters up to 6 m in length and 1–100 cm in width are common. These features suggest that the NV1 sediments were unconsolidated and saturated at

Table 4 Lithofacies thickness at various distances from the vent (LT1 is not included as it was rarely completely exposed, and accurate thicknesses were not obtained) Facies Association

Lithofacies

0.6 km (SWI) max thickness (cm)

2.4 km (W4) max thickness (cm)

2.7 km (WI0) max thickness (cm)

3.5 km (NEI, N4) max thickness (cm)

PHI PH2 PH2 PH2 PH2 PH2

LT2 LT4 LT3 T2 LT6 LT5

32 25 25 10 0 0

10 70 50 10 100 80

10 40 30 10 80 40

6 50 50 20 0 0

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Fig. 14. Wavelength versus distance from source for several well-documented surge deposits (Mt St Helens Blast, Druitt, 1992; El Chichon, Sigurdsson et al., 1987; Ubehebe, Crowe and Fisher, 1973; Taal, Moore, 1967; Sinker Butte, Brand and White, 2007; Narbona Pass, Brand et al., in press). Note: Only the measured wavelengths with distance from source are shown on this plot, thus the total runout distance is not represented for each example.

The beginning of the pyroclastic section is dominated by eruptionfed aqueous density currents, some of which probably transformed after initial fallout. These deposits were interpreted to have been eruption-fed rather than reworked due to the dominance of juvenile material (N95%), general consistency in the bedding style, and lack of significant grain abrasion and rounding. Furthermore, the repeating packages of lithofacies (i.e., LT1–T1, LT2, LT4–LT3–T2) indicate that this was a pulsating eruption, and each repeated lithofacies or set of lithofacies (i.e., LT6–LT5–T3) represents one explosive pulse. Individual, fine-ash glass shards within the matrix are fractured, blocky, and angular, indicating that the dominant mechanism of fragmentation was magma–water interaction (Heiken and Wohletz, 1985). White (2000) argues that although subaqueous eruption-fed flows commonly involve water-supported transport, the transport and depositional processes are controlled by the nature of the eruption and its interaction with the surrounding water. The contrast in the bedding style and thickness, the juvenile grain morphology between individual lithofacies, and the considerable differences between FAs PH1 and PH2 demonstrate this well; both FAs were determined to have been emplaced subaqueously, however, we attribute the differences in deposit characteristics to differences in eruptive style. PH1 begins with thick, alternating sequences of tuff and lapilli tuff that often contain either lenses of the underlying lake sediment, or in some cases large, elongated, and intact slabs of the substrate (section W10; Fig. 8b), suggesting that the first deposits were dominated by lateral erosive transport. The sharp transition to LT2 deposits, which are repeating 10–30 cm thick beds of reverse-graded coarse ash to lapilli, followed by finely laminated fine ash, are interpreted as a combination of fallout-initiated, thin density currents, and watersettled, fine ash tuff. While the initial 5–8 m of PH1 deposits represent thick, concentrated flows, the more thinly bedded deposits of LT2 suggest that the eruption eventually attained a higher frequency of jetting or explosivity. PH1 contains coarse lapilli to block-sized juvenile scoria with quenched rinds. These coarse juvenile grains are found either randomly dispersed throughout the T1 and LT1 lithofacies and supported in a fine-grained matrix, or in concentrated horizons in the reverse-graded strata of lithofacies LT2. While it has been determined based on ash morphology and texture that magma– water interaction was the dominant fragmentation mechanism, the fluidally-shaped, scoriaceous juvenile coarse lapilli-to-fine blocks suggest a concurrent fire-fountaining stage at the beginning of the eruption (Houghton and Schmincke, 1986; Mueller and White, 1992). As all evidence points to magma interacting with an abundant source of lake water, the early phase of fire-fountaining may suggest an

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initially high magma mass flow rate at the onset of the eruption that could have isolated some of the erupting magma from the external water, and led to less efficient mixing (Wohletz and McQueen, 1984; Büttner and Zimanowski, 1998; Mastin, 2007). The lack of welding textures in the matrix surrounding the fluidal juvenile clasts suggests that the flows which produced these deposits contained water as the continuous intergranular phase rather than hot gas, which is consistent for deposits of Surtseyan-style explosive eruptions (White, 1996; Sohn, 1997; White, 2000). In contrast, hot gas is speculated to be the interstitial fluid in high-temperature pyroclastic flows into subaqueous environments when the flows are derived from sustained explosive eruptions with well-developed gas thrust regions (Sparks et al., 1980; Cas and Wright, 1991; Kokelaar and Busby, 1992; Schneider et al., 1992; White and McPhie, 1997; White, 2000). Therefore, PH1 deposits likely originated from eruption-fed aqueous density currents including combinations of low- and highconcentration turbidity currents (Lowe, 1982; Postma et al., 1988; Kneller and Branney, 1995), cohesionless sediment gravity flows or grain flows (Lowe, 1976; Postma, 1986; Nemec, 1990), and a concurrent fallout component. As only the coarsest juvenile clasts show a fluidal nature, we suggest that the larger clasts remained insulated by self-generated steam jackets within a flow in which water was the interstitial fluid phase (White, 2000). The PH2 sequence is interpreted to be the result of multiple, highdensity, eruption-fed turbidity currents, where the overall fining upward nature and progressive development of different bedding structures (i.e., normal and reverse grading, diffuse stratification, finegrained interbeds) in the deposits represents waxing and waning flow conditions over the duration of deposition (Bouma, 1962; Nemec et al., 1980; Lowe, 1982; Postma, 1986). This repeating set of lithofacies elucidates the pulsating nature of the eruption, and the variation in grain size and depositional characteristics from one package to another represents variations in initial particle concentration, grain size, and velocity of each flow. The subaqueous deposits thicken away from vent, consistent with a growing platform with shallow slopes (1–7°) into deeper water. Table 4 shows the variation of individual lithofacies thickness with distance from vent. The coarsest deposits, LT6 and LT5, are not recognized in either proximal or distal locations from source. It is interpreted that the initial high-density turbidity currents carrying the coarser-grained loads of LT6–LT5 had a higher initial capacity and competence to carry the coarse clasts N0.6 km away from the vent before sedimentation occurred. The lack of these lithofacies at the most distal exposures suggests that the coarsest grains settled out at medial locations. This also suggests that, while we have distinguished LT6–LT5–T3 and LT4–LT3–T2 stratigraphically, the deposits of LT4–LT3 in the distal regions may represent a lateral facies change from the coarser-grained LT6–LT5 currents. It is also interesting to note that the deposits of LT4–LT3 thicken with distance from source, with the thickest deposits occurring at a distance of ∼2.4 km. These deposits subsequently thin to a distance of 2.7 km. This likely represents a low initial rate of deposition, followed by an increasing rate of sedimentation with distance from source, which finally tapers off with decreasing current load at distal locations. Note also that the most distal deposits of LT4–LT3 thicken again, and then follow the thinning trend of LT6–LT5 (Table 4), which also supports a lateral facies change from LT6–LT5 to LT4–LT3. Significant soft sediment deformation (SSD) was found throughout PH1 and PH2 deposits. Some SSD occurs in the steep walls of scour and fill channels, where it is interpreted that thinly bedded, fine-grained fill deposits slipped and formed convolute folds towards the axis of the channel. Other SSD in the form of flame features and localized folded strata appear to be due to the weight of overlying and likely quickly deposited pyroclastic deposits. SSD is also found beneath ballistic blocks. All of these SSD features attest to the subaqueous and saturated conditions at the time of deposition. In several distal locations around

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the TRC1 remnants, lake sediments of NV2 were identified at the contact of the TRC1 and TRC2 deposits, suggesting that the lake level was well above the height of the distal TRC1 deposits long after the end of the TRC1 eruption, and that much of the distal deposits may have been eroded below the wave base. Additionally, this suggests a significant repose period between the TRC1 and TRC2 eruptions. Given the distance from the subaqueous–subaerial contact to the solidified lava lake, the subaerial portion of the TRC1 tuff cone grew at least 285 m above the lake level. The most proximal, final pyroclastic deposits recognized before the transition to spatter at the top of TRC1 contain similar accidental clasts as in the lower tuff cone, and also have similar juvenile glass morphologies including bubble number density and size. This suggests that as the eruption evolved, the location of magma–water interaction did not excavate into the subsurface to entrain deeper, or a higher percentage of, country rock fragments. Additionally, these deposits are highly palagonitized and contain a high proportion of accretionary and armored lapilli, which supports the interpretation that abundant liquid water was present at the time of deposition. This implies that a relatively high water– magma ratio existed just before the magmatic stage. The transition from phreatomagmatic to Strombolian deposits is evidenced by a rapid increase in progressively coarser-grained scoria (1–3 cm) over two vertical meters just below the Strombolian-spatter deposits. This observation suggests that the vent was quickly sealed off from the influence of external water to produce the Strombolianspatter and a ponded lava lake at the close of the eruption. 4.2. TRC2 TRC2 was the last and most energetic eruption to occur in the TRC. The crater rim is ∼ 2.7 km in diameter, and the deposits, which form a broad tuff ring surrounding the crater, are dominated by largescale pyroclastic base surge deposits. The TRC2 eruption occurred a significant time after the TRC1 eruption, as evidenced by the 25–50 cm of lake sediments found at the TRC2–TRC1 contact. Additionally, there appears to have been a significant drop in the regional lake level prior to the TRC2 eruption. Primary depositional characteristics of surge currents are found as low as 1325 m, which suggests the lake, if present at the time of the TRC2 eruption, was no deeper than 15 m. Wave cut terraces within the TRC2 deposits are present at 1356 m elevation, which indicates that at some point after the eruption the lake water rose to at least 46 m deep. The TRC2 surge deposits drape features to the east–southeast and northwest of the TRC1 tuff cone, which are likely well below the maximum height of the original tuff cone walls (Fig. 2). This suggests either that much of the TRC1 tuff cone was eroded away prior to deposition of the TRC2 deposits, or it was eroded by the TRC2 eruption itself. A difficult question to answer is whether or not the TRC2 eruption occurred due to magma–water interaction at depth to create a maar and surrounding tuff ring, or if it erupted in a near surface or playalake setting to create a tuff ring with the crater floor above the level of paleotopography. While the accidental clast component is much higher in the TRC2 deposits (between 30 and 55%) than in the TRC1 deposits (b5%), it is not quite as high as proportions recognized for other maar deposits where the accidental component composes up to 80% of the total tuff deposits (Cas and Wright, 1987). However, the true accidental component cannot be quantified with certainty as it is impossible to distinguish between the accidental clasts derived from the TRC1 deposits and the juvenile clasts of the TRC2 deposits. Thus, 30–55% is a minimum estimate for the total amount of accidental material in the TRC2 deposits. The crater floor at 1356 m is well above the floor of the coeval lake (1310 m). However, one could argue that the eruption occurred through the flank of the remnant TRC1 tuff cone, and therefore may have interacted with groundwater within the tuff cone deposits and

underlying country rock. Based on the accidental clast composition, and the level of the TRC2 crater floor, we hypothesize that the magma–water interaction occurred dominantly within the TRC1 deposits, and that the source of the water was related to a shallow playa lake, saturated, unconsolidated or very poorly lithified TRC1 deposits, and/or near-surface, water-saturated playa-lake sediments. Therefore the TRC2 feature could be considered to be a shallow maar in the sense that magma–water interaction occurred below the preexisting surface, and accidental clasts derived from the underlying country rock are common, although subordinate to the juvenile clasts. The rest of the TRC2 discussion will focus on the large scale surge bedforms that comprise the tuff ring, and the insight they provide into eruption dynamics. The pyroclastic surge deposits radiate axisymmetrically from the center of the crater. Vertical cross-sections through the TRC2 dune features reveal large-scale hummocky crossstratification. The fact that each set of dunes truncate the previous ones, and form the same three-dimensional, hummocky features suggests that the deposits were not simply mantling a pre-existing, hummocky topography. Rather the currents were carving, depositing, and re-creating the hummocky topography observed around the edifice with the passing of each surge current. These features demonstrate the pulsating nature of the explosive eruption. The dune forms at TRC2 are more than an order of magnitude larger than dune features recognized in other basaltic hydrovolcanic deposits (e.g., Fisher and Waters,1970; Waters and Fisher,1971; Crowe and Fisher, 1973; Schmincke et al.,1973; Sohn and Chough,1989; Dellino et al.,1990; Chough and Sohn, 1990; Brand and White, 2007; Brand et al., in press), and their unusually long wavelengths extend to great distances away from the vent (e.g., 20-m wavelengths recorded 4.7 km from source). The flows were highly erosive and scoured into the TRC1 deposits on all sides of the tuff ring, truncate earlier surge deposits from the same eruption, and form large U-shaped channels in the north. The channels were interpreted by Heiken (1971) to have been carved out by an extremely erosive surge, similar to the U-shaped channels described at Koko Crater, Hawaii (Fisher, 1977). Most U-shaped channels have been recognized on the steep slopes of tuff cones rather than the shallow slopes of tuff rings, and it has been suggested that large U-shaped channels similar to those at TRC2 indicate much larger and faster surges than those required to form similar features in steeper tuff cones (Fisher, 1977). When we plot the TRC2 data (wavelength vs. distance from source) along with data from the literature, we see that wavelengths of the TRC2 bedforms are much longer than those recorded at other basaltic hydrovolcanoes, and appear to follow the trends of larger scale eruptions such as El Chichon and Mt St Helens (Fig. 14). At El Chichon, similar three-dimensional features with shorter wavelengths were observed at distances greater than 4 km, and the authors suggest that if the dune wavelengths at greater distances were extrapolated toward the source, they should have wavelengths of 100 to 200 m (Sigurdsson et al., 1987), consistent with those at TRC2. The TRC2 surges are unusual in the sense that wavelengths of this scale are typically found on volcanic flanks with much higher aspect ratios and slopes, consistent with high column collapse or energetic blast origins (Sigurdsson et al., 1987; Druitt, 1992), or in large ignimbrite ash flow tuffs (e.g., Ohakuri-ignimbrite forming eruption in the central Taupo Volcanic Zone; Gravley et al., 2007). The unusually long wavelengths of the dune features proximal to the vent are indicative of highly inflated currents at the onset of the flows. The most likely way to obtain highly inflated flows proximal to source is to entrain air during rise and collapse of an eruption column (Sigurdsson et al., 1987). Therefore, the base surge currents are interpreted to be generated by collapse from an unusually high column. This is further supported by the runout distance, which, based on the 20 m wavelengths found at the distal exposures, was likely much greater than the exposed 4.7 km. By comparison, eruption column collapses at Ambae Island, Vanuatu, on the order of a couple hundred meters, produced surges to distances of 300 m (Nemeth et al., 2006), and the

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phreatomagmatic eruption in Karymskoye Lake, Kamtchaka, collapsed from 1 km and produced surges to distances of b2 km (Belousov and Belousova, 2001). A general relationship can be made from these data and weapons test-induced surge currents (e.g., Sedan weapons test; Rohrer, 1965); surge runout distance is approximately 1.3 to 1.5 times the column collapse height. The TRC2 surge deposits extend at least 4.7 km from source implying column collapse heights of 3–4 km, which is unusual for basaltic hydrovolcanic eruptions. On the fine-scale, the deposits within TRC2 are dominated by blocky, angular shards of fine ash, which suggests highly efficient hydromagmatic fragmentation. The juvenile glass shards have thickwalled bubbles, and low bubble fractions (0.01 to 0.25), implying that magmatic volatiles did not play an important role in magma fragmentation. The abundant fine-ash fraction combined with the large-scale, long runout surge features, are evidence of efficient conversion of thermal to kinetic energy and a highly energetic eruption, respectively. However, the deposits contain ample evidence for abundant liquid water in the surges during transport and emplacement, which is contradictory, as these features are often attributed to water–magma ratios above that which flashes all water to steam, and thus are associated with lower explosivity (Sheridan and Wohletz, 1983; Wohletz and McQueen, 1984). Based on all observed features, it is interpreted that the eruption was highly energetic and likely had an efficient conversion of thermal to mechanical energy, vaporizing most if not all external water at the site of magma water interaction. It is then hypothesized that the water vapor subsequently condensed due to entrainment-induced cooling during ascent and collapse from the eruption column to produce the “wet” deposits. The small-scale, centimeter to decimeter-thick strata within the larger dune features are interpreted to represent temporal variation in the bed load of a density stratified current due to flow unsteadiness, resulting in layer-by-layer variations during vertical aggradation (i.e., Druitt, 1992; Brown et al., 2007). The multiple scales of dune wavelengths, from the largest that create the hummocky topography around the tuff ring, to the medium dunes (up to 50 m) that appear nested within the larger wavelength dunes, to the smallest antidunes recorded in the individual 1–2 m thick packages of sediment (3–6 m in wavelength), support the Valentine (1987) model of deposition within a density stratified current. The various scales of dunes likely represent internal gravity waves on different length scales within the larger flow, as discussed in Section 2.3.2 above. Extensive base surge runout distances have been recognized for other phreatomagmatic eruptions. The Taal Volcano produced base surges with a runout of up to 6 km (Moore et al., 1966; Moore, 1967), the Monte Guardia volcano, Lipari with a runout of up to 7 km (Colella and Hiscott, 1997), and the Glaramara tuff in Scafell caldera, English Lake District, UK with runout up to 8 km (Brown et al., 2007). The latter two examples were associated with higher silica magmas, and based on the presence of pumiceous clasts in their deposits, likely had significant influence from magmatic volatiles (Colella and Hiscott, 1997; Brown et al., 2007). The explosive, basaltic eruption at Taal volcano matches TRC2 closest in terms of runout distance, but the bedform wavelengths at Taal were more than an order of magnitude smaller (Fig. 14; Moore, 1967). Thus, the TRC2 dune wavelengths expand the range possible for basaltic phreatomagmatic eruptions. 5. Conclusions TRC is a basaltic, polygenetic volcano that formed during several styles of hydrovolcanism, varying from Surtseyan to maar-style eruption. The first eruption to occur in the complex is represented by a large tuff cone (TRC1), which is the tallest feature in the complex. It began erupting through a 60–70 m deep lake, initially depositing multiple eruption-fed turbidity currents until a gently-dipping platform was built up above the surface of the lake. The hydromagmatic

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activity subsequently constructed a steep tuff cone with a crater floor ∼365 m above the base of the lake. The beginning of the eruption is interpreted to have had both a hydromagmatic and a simultaneous fire-fountaining stage, however, the eruption changed to an entirely hydromagmatic phase after the initial activity. The rest of the subaqueous deposits contain a mixture of fine ash and lapilli, suggesting moderately efficient hydromagmatic mixing. The subaerial deposits of TRC1 are not well preserved due to posteruptive erosion, but where present consist of base surge and fall deposits. The focus of explosivity was shallow and did not core into the pre-volcanic substrate, and the final eruptive stages are marked by both efficient hydromagmatic fragmentation and evidence for liquid water at the time of deposition, suggesting relatively high water-to-magma ratios. The eruption ended with a transition from hydromagmatic to Strombolian and effusive activity and produced a crater-filling lava lake that partially intruded the unconsolidated tuff cone walls. Once the eruption ceased, much of the medial to distal tuff cone was below wave base allowing portions of it to be eroded away, and allowing the TRC1 deposits to remain saturated long after the eruption. A long repose period occurred before later volcanic activity resumed. The seven flank vents around the complex erupted some time during or after the TRC1 activity, but before the TRC2 eruption. TRC2 is represented by a 2.7 km diameter crater in the northeast of the complex. It was the most energetic eruption in the complex, and likely occurred when magma interacted with a shallow lake, playa-lake, and/or saturated TRC1 tuff sediments. The TRC2 eruption produced multiple, highly inflated, erosive pyroclastic surges that radiated out from the base of a collapsing column. Dunes in the TRC2 surge deposits have longer wavelengths than other basaltic hydromagmatic deposits, and are instead more consistent with dune bedforms produced by larger eruptions such as the 1980 blast at Mt St Helens, and the column collapse surges at El Chichon in 1982. Such large dune wavelengths are rare for low-aspect ratio tuff rings, and suggest highly inflated flows produced by an unusually high eruption column. Acknowledgements The authors would like to thank Grant Heiken for suggesting this field area and spending several days in the field with B.B. We are especially grateful for the assistance of Mike, Jean, and Josh Bandfield, and Adam Frus in the field during summer 2006. This work greatly benefited from the thorough and insightful reviews of John Smellie and Corina Risso, and guest editor Károly Németh. Funding for this research was provided by an Arizona NASA Space Grant Consortium Student Fellowship, and the National Science Foundation, USA (EAR 0538125). Appendix A. Supplementary data Supplementary data associated with this article can be found, in the online version, at doi:10.1016/j.jvolgeores.2008.10.011. References Allen, J.R.L., 1982. Sedimentary Structures: Their Character and Physical Basis. Development in Sedimentology, vol. 30A & B. Elsevier, Amsterdam. 663 pp. Arnott, R.W.C., Hand, B.M., 1989. Bedforms, primary structures and grain fabric in the presence of suspended sediment rain. J. Sed. Petrol. 59, 61062–61069. Belousov, A., Belousova, M., 2001. Eruptive process, effects and deposits of the 1996 and the ancient basaltic phreatomagmatic eruptions in Karymskoye Lake, Kamtchaka, Russia. Spec. Publ. - Int. Assoc. Sediment. 30, 35–60. Bouma, A.H., 1962. Sedimentology of some Flysch deposits; a graphic approach to facies interpretation. Elsevier Pub. Co., Amsterdam. 168 pp. Brand, B.D., White, C.M., 2007. Origin and stratigraphy of phreatomagmatic deposits at the Pleistocene Sinker Butte Volcano, Western Snake River Plain, Idaho. J. Volcanol. Geotherm. Res. 160, 319–339. Brand, B.D., Clarke, A.B., Semken, S. Eruptive Conditions and Depositional Processes of Narbona Pass Maar Volcano, Navajo Volcanic Field, Navajo Nation, New Mexico (USA). Bull. Volcanol. in press. doi:10.1007/s00445-008-0209-y.

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