the architecture, eruptive history, and evolution of the table rock complex… · 2015. 10. 4. ·...

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

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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

⁎ Corresponding author.E-mail addresses: [email protected] (B.D. Brand), ama

(A.B. Clarke).1 Tel.: +1 480 965 6590; fax: +1 480 965 8102.

0377-0273/$ – see front matter © 2008 Elsevier B.V. Aldoi:10.1016/j.jvolgeores.2008.10.011

a b s t r a c t
a r t i c l e i n f o
Available online 28 October 2008

Keywords:

The Table Rock Complex (TR1971. Tuff rings; examples fr76, 5615-5626.], is a large an

phreatomagmaticSurtseyanmaarbase surgeTable Rock Complex

d well-exposed mafic phreatomagmatic complex in the Fort Rock–Christmas LakeValley Basin, south-central Oregon. It spans an area of approximately 40 km2, and consists of a large tuff cone inthe south (TRC1), and a large tuff ring in the northeast (TRC2). At least seven additional, smaller explosion craterswere formed along the flanks of the complex in the time between the two main eruptions. The first period ofactivity, TRC1, initiated with a Surtseyan-style eruption through a 60–70 m deep lake. The TRC1 deposits are

C; Pliocene–Pleistocene), first documented and described by Heiken [Heiken, G.H.,om the Fort Rock-Christmas Lake valley basin, south-central Oregon. J. Geophy. Res.

dominated by multiple, 1-2 m thick, fining upward sequences of massive to diffusely-stratified lapilli tuff withintermittent zones of reverse grading, followed by a finely-laminated cap of fine-grained sediment. The massivedeposits are interpreted as the result of eruption-fed, subaqueous turbidity current deposits; whereas, the finelylaminated cap likely resulted from fallout of suspended fine-grained material through a water column. Othercommon features are erosive channel scour-and-fill deposits, massive tuff breccias, and abundant soft sedimentdeformation due to rapid sediment loading. Subaerial TRC1 deposits are exposed only proximal to the edifice, andconsist of cross-stratified base-surge deposits. The eruption built a large tuff cone above the lake surface endingwith an effusive stage, which produced a lava lake in the crater (365 m above the lake floor). A significant reposeperiod occurred between the TRC1 and TRC2 eruptions, evidenced by up to 50 cmof diatomitic lake sediments atthe contact between the two tuff sequences. The TRC2 eruptionwas the last and most energetic in the complex.General edifice morphology and a high percentage of accidental material suggest eruption through saturatedTRC1 deposits and/or playa lake sediments. TRC2 deposits are dominated by three-dimensional dune featureswithwavelengths200–500mperpendicular to theflow, and20–200mparallel to thedirectionofflowdependingon distance from source. Large U-shaped channels (10–32m deep), run-up features over obstacles tens ofmetershigh, and a large (13 m) chute-and-pool feature are also identified. The TRC2 deposits are interpreted as theproducts of multiple, erosive, highly-inflated pyroclastic surges resulting from collapse of an unusually higheruption column relative to previously documented mafic phreatomagmatic eruptions.

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1. Introduction

Hydromagmatic eruptions occur when rising magma violentlyfragments after intersecting andmixing with shallow surface water orgroundwater (Sheridan and Wohletz, 1983). Fragmentation in thisstyle of volcanism is driven principally by the energetic interactionbetween magma and external water (Houghton and Wilson, 1989),although expansion of magmatic volatiles can occur depending on thevolatile content of the magma, and may provide a secondarymechanism of fragmentation (i.e., Houghton and Wilson, 1989;Houghton et al., 1999; Cole et al., 2001; Brand et al., in press). Thedegree of fragmentation associated with magma–water interaction

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has been experimentally and theoretically determined to be a functionof melt composition, magma flux, water–melt mass ratio, confiningpressure, magma viscosity, and the degree of turbulent mixing ofmagma 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 inunderstandinghow the hydromagmatic deposits of tuff cones, tuff rings,andmaars relate to the eruptivedynamics anddepositionalmechanismsthat produced them (e.g., Fisher and Waters, 1970; Crowe and Fisher,1973; Lorenz, 1974; Sheridan and Wohletz, 1983; Kokelaar, 1983; Fisherand Schmincke, 1984; Houghton and Hackett, 1984; Kokelaar, 1986;Sohn and Chough, 1989; Dellino et al., 1990; White, 1996; Houghtonet al., 1999; White, 2001; Nemeth et al., 2001; Cole et al., 2001; Mastinet al., 2004; Brand andWhite, 2007; Brand et al., in press). To assess therelative influence of external water (water–magma ratio) on overalleruption dynamics, researchers look at deposits for evidence for liquid

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204 B.D. Brand, A.B. Clarke / Journal of Volcanology and Geothermal Research 180 (2009) 203–224

water at the timeof deposition.Wet conditions are typically indicated bya combination of features such as accretionary lapilli, fine-grainedvesiculated tuff, pervasive soft-sedimentdeformation, surge cross-stratawith stoss-side accretion and evidence for plastering, mud cracks,debris-flow filled erosion channels, and lahar deposits (Waters andFisher, 1971; Lorenz, 1974; Wohletz and Sheridan, 1983; Sohn andChough,1989;Dellinoet al.,1990;Nemethet al., 2001;White, 2001; Coleet al., 2001; Brand and White, 2007; Brand et al., in press). Wet-phreatomagmatic conditions are interpreted to occur when externalwater at the zone of magma–water interaction is not efficientlyconverted to steam (Sheridan and Wohletz, 1983; Wohletz andMcQueen, 1984) and abundant liquid water is retained in the eruptioncolumn and density currents.

Dry conditions are distinguished by their lack of features indicativeof liquid water at the time of pyroclastic deposition. These depositscontain no evidence of accretionary lapilli, little to no soft-sedimentdeformation, low-angle cross strata with dominantly lee-side accre-tion (although antidunes are also common), and no evidence formuddy deposits plastering against obstacles or on the stoss side ofdunes (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-phreatomag-matic conditions represent efficient conversion of external water tosteam at the site of magma–water interaction, or subsequently in theeruption plume. As a consequence, little or no liquid water is retainedin the eruption column or proximal density currents, although steammay condense in density currents and rising plumes with distancefrom the source (e.g., Sheridan and Wohletz, 1983; Wohletz andMcQueen, 1984). These conditions are thought to represent highlyefficient conversion of thermal energy to kinetic energy, yieldingthe 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, tuffrings, 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 ofhydrovolcanism in the Fort Rock-Christmas Lake Valley Basin, was firstmapped and described by Heiken (1971). We have revisited TRC withthe goal of studying the stratigraphy in detail in order to reconstructthe volcanic evolution of the complex. Our results provide evidencefor two major eruptions, TRC1 and TRC2, for which we constrain thetemporal evolution, dominant depositional mechanisms, the influ-ence of liquid water on deposit characteristics, and relative eruptionenergy. The objectives of our work are to build on the existingframework for hydromagmatic pyroclastic deposits by continuing toidentify relationships between deposit characteristics and eruptivedynamics of mafic hydromagmatic eruptions.

The main topics of this paper include (1) subaqueously emplaceddensity currents of a large Surtseyan-style eruption early in thecomplex's history (TRC1); and (2) The formation of multiple, large-scale base surge deposits produced during a later, highly energeticmaar-forming eruption (TRC2). Finally, a more general, broad-scalereason for studying tuff cones, tuff rings, and maars is that they offeran opportunity to observe, describe, and study the deposits of manypyroclastic processes on easily accessible vertical and lateral scales.

1.1. Geologic setting

The Fort Rock–Christmas Valley basin is 64 km long by 40 kmwide,and was the location of an extensive, ancient Pliocene to Pleistocenelake (lake boundary dashed in Fig. 1; Heiken, 1971). The basin isoccupied by a variety of basaltic eruptive features which trendnorthwest to southeast. The ages of these features are poorlyconstrained, 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 dominantlytuff cones and tuff rings; whereas, towards the boundaries thefeatures are dominated by maars and cinder cones (Heiken, 1971).

The oldest formation identified in the area is the Picture Rockbasalt, a 230 m thick sequence of 10 m thick basalt flows interbeddedwith 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 immediatelysouth of TRC, is the beginning of a 25 km wide structural arch thatforms the southwest boundary of the lake basin (Heiken, 1971).Driller's logs show that ∼220 m of flat-lying lacustrine sediments andinterbedded tuffs overlie the Picture Rock basalt formation at thecenter of the basin, and that these sediments thin to 0 m towards thebasin boundary (Hampton, 1964). At the center of the basin (beneathTRC), these sediments consist primarily of diatomites, whereas closerto the basin margins they consist of coarse clastic sediments, lavaflows, and volcanic breccias (Heiken, 1971).

The underlying stratigraphy exposed on the west side of TRCconsists of a 5–10 m thick basaltic lava flow, followed by ∼8 m ofinterbedded volcanic litharenites, lithic arkoses, and diatomaceoussiltstone and mudstones (Heiken, 1971). In contrast, the non-volcanicstratigraphy beneath the eastern portion of TRC is dominated by well-bedded diatomites (Heiken, 1971). The interbedded, non-volcanicunits exposed on the west side of the edifice are interpreted asoutwash apron deposits from the Connley Hills to the northwest, a6.4 km wide by 19-km-long volcanic feature consisting of a basalticshield and intermediate to silicic domes, which likely formed an islandin 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 ofpresent day Silver Lake, which is likely the small remnant of the muchlarger lake that occupied the Fort Rock–Christmas Valley Basin inPleistocene time (Heiken, 1971). TRC has an elongated-oval shape,trends NNW (along strike with the other phreatomagmatic complexesin the Fort Rock–Christmas Valley region), and covers an areaapproximately 8 km by 5 km (Heiken, 1971). Two large phreatomag-matic edifices; a large southern tuff conewith a capping solidified lavalake at 395m above the basin floor (TRC1), and a low, broad tuff ring inthe northeast (TRC2; Fig. 2), make up the complex. Additionally, sevensmaller tuff rings and vents were identified along the flanks of thecomplex, yielding a complicated network of tuff ring–tuff conedeposits. For the purposes of this paper, only the two largest andmost significant eruptions of TRC1 and TRC2 will be discussed indetail. 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 sectionswere completed to determine the eruptive history of TRC (Fig. 2),resulting in the identification of 42 lithofacies based on variations ingrain size, composition, and sedimentary structures (following Fisher,1961; Schmid, 1981; Table 1). Detailed descriptions of each lithofaciesare available as online supplementary data. These lithofacies weregrouped into six Facies Associations (FA) according to commonbedding styles, juvenile fragment morphology, and the percentageand type of accidental components (Table 2). Five of these FAscorrespond to the TRC1 eruption, and one corresponds to the TRC2eruption. Inferred eruptive conditions and depositional mechanismare discussed below in the text.

Grain percentage was estimated in the field and supported bysubsequent petrographic thin section analysis of each FA frommultiple locations around the complex. In the field, juvenile clastswere distinguished from accidental basalt by their fresh appearance,

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 imageshows the location of the Table Rock Complex, as well as many of the other hydromagmatic edifices and cinder cones within the basin.

205B.D. Brand, A.B. Clarke / Journal of Volcanology and Geothermal Research 180 (2009) 203–224

irregular shapes, or twisted-fluidal shapes, greater vesicularity, andquenched glassy rinds with vesicular cores. Thirty thin sections,sampled from multiple lithofacies of the TRC1, TRC2, and flank venteruptions, were analyzed and point counted to better constrain thecomposition and percentage of accidental versus juvenile clasts,especially in the matrix. For each thin section, 300 points werecounted 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 thelarge tuff cone with a capping solidified lava lake in the south-centralpart of the complex (Fig. 2). This was the first eruption in the sequenceof tuffs, as it directly overlies the pre-existing lake and lacustrinesediments described above. The deposits are well exposed along thesouth, east, and north sides of the edifice, and intermittently exposedon the west side as they were eroded away or covered by latereruptions. The sequence is consistent around the vent with minorvariations in depositional characteristics such as grain size andbedding 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 ofthe deposits also contain abundant fine-to-medium ash matrix.Petrographic analysis of the matrix reveals that ∼98% of the matrixand grains within the thin sections are of juvenile origin (a detailedthin section analysis is presented in Section 2.3 below).

2.1.1. PH1: massive to stratified tuff and lapilli tuffsPH1 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 amassive, matrix supported (95–100% fine-medium grained ash) depositwith occasional imbricated pebble stringers of mudstone and rareaccidental 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 coarselapilli to block-sized juvenile grains consist of scoriaceous, twisted,fluidal, and often flattened juvenile grains with quenched rinds (foundboth in LT2, and at random, concentrated horizons within LT1). Thesejuvenile spatter-bomb clasts are not found above the PH1 horizon. Softsediment deformation is prevalent throughout PH1.

2.1.2. PH1 interpretationT1 and LT1 are likely a combination of tephra fallout and

subaqueous sediment gravity flows. The angular grains and moderatesorting within LT1 may be indicative of fallout, possibly througha density current (Valentine and Giannetti, 1995), whereas thelenticular interbeds, erosive contacts, and reverse grading oftenfound within LT1 are consistent with a deposition by lateral move-ment of grains (White, 1996, 2000). The fine-grained, massive, poorlysorted, and non-graded deposits of T1 suggest deposition from aconcentrated suspension density current with little tractional trans-port, which inhibits the development of grading (Sparks, 1976; Sparkset al., 1978; Chough and Sohn, 1990; Freundt and Bursik, 1998). Thereverse-graded, coarse ash to fine block deposits of LT2 are interpreted

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


Table 1Lithofacies 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, anddominance of juvenile fragments [table modeled after Nemeth et al. (2001) and Nemeth and White (2003)]. Detailed descriptions of individual lithofacies are available as onlinesupplementary data

Lithofacies Tuff Breccia (TB) Lapilli Tuff (LT) Tuff (T) Magmatic (M) Non-volcanic (NV)

Pre-TRC1 eruption NV1

TRC1 eruptionClast supported

Massive LT4, LT6Massive-to-diffuse stratification LT3, LT5Stratified LT1Stratified, reverse grading LT2Scour-fill massive TB1Scour-fill bedded LT8, LT9Cross-stratified LT18

Matrix supportedMassive-to-diffuse stratification T1Diffusely stratified T18Stratified, laminated T2, T3Scour-fill bedded LT7Planar-to-wavy beds T17

Strombolian deposits, lava lake, and radial dikes M1

Post-TRC1 eruption NV2

TRC2 eruptionClast supported

Massive, non-graded TB2Massive, reverse grading LT17Massive-to-diffusely stratified bed in a well-stratified deposit sequence LT10 T13, T15Stratified, reverse grading LT14, LT15, LT16 T21Scour-fill massive LT12Planar-to-wavy stratification, occasional lenticular deposits LT11Wavy-to-cross-stratified T20

Matrix supportedMassive T10Massive-to-diffusely stratified bed in a well-stratified deposit sequence T9, T16Stratified — pinch and swell T4, T5Stratified T12Stratified, laminated T11, T14,Scour-fill bedded T6, T7, T8Wavy-to-cross stratified T19


as a combination of fallout and density current deposits. The coarsedeposits of LT2 are interpreted as having an initial fallout origin that,after landing, transitioned into thin, laterally moving, traction carpetdominated, subaqueous sediment gravity flows. The finely-laminated,well-sorted, normal-graded ash beds that overlie the coarser depositsare interpreted to have been deposited by ash-fallout settling througha column of water. These beds are typically highly deformed under theweight of the overlying material, further suggesting that they wereunconsolidated and water-saturated during emplacement of subse-quent units (Nichols et al., 1994; White, 2000). The field and thinsection analyses show that N98% of the grains are juvenile andtypically angular to subrounded, indicating minor to no abrasion (seeSection 2.2 below; also see online supplementary data).

The angular, blocky glass shards indicatemagma–water interactionas the dominant fragmentation mechanism. However, the fluidal,scoriaceous juvenile coarse lapilli-to-fine blocks suggest that, in theearly stages of the eruption, both hydromagmatic and magmaticprocesses were occurring simultaneously (Houghton and Wilson,1989; Cole, 1991; Houghton et al., 1999).

2.1.3. PH2: normal graded, massive to diffusely stratified sequencesThe PH2 facies association consist of repeating packages of lithofacies

(i.e., LT6–LT5–T3 and LT4–LT3–T2; Tables 1 and 2). Each package beginswith a coarse-grained (angular to subangular grains of fine-to-coarselapilli), 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 samebed (Fig. 4), however thepackage as awhole graduallyfines upwards (LT3,

LT5), and is capped by a planar-to-cross laminated,medium-to-coarse ash(grains subangular to subrounded) bed that commonly displays softsediment deformation (20–40 cm thick, T2, T3; Fig. 4). Individuallithofacies pinch and swell laterally, but the sequences overall are laterallycontinuous for 10's up to 100's of meters away from source.

Field and petrographic analysis show that the grains in all of thesedeposits are dominantly juvenile. Accidental grains of mudstone anddense, angular basalt compose b5% of the deposits. However, many ofthe block-sized ballistic clasts with underlying sags are composed ofaccidental basalt.

2.1.4. PH2 interpretationThe repeated, fining upward packages of PH2 are interpreted to

correspond to the traction carpet and suspension sedimentation stagesof high-density, turbulent, subaqueous sediment gravity flows (Nemecet al.,1980; Lowe,1982; Postma,1986). Lithofacies LT6–LT5 and LT4–LT3are thick, diffusely stratified deposits with alternating coarsening- andfining-upward sequences. They have no distinct bedding surfaces, andhave an overall fining upward trend. Changes in grain size throughout agiven deposit reflect the grain size variation of supplied sediment withtime, and/or intermittent waxing and waning flow conditions (Sohn,1997). The overall fining upward sequence is interpreted to be a con-sequence of waning flow conditions and loss of sediment supply overthe duration of the current. The fine-grained, stratified lithofacies (T2,T3), which overlie the coarser, thicker deposits, represent the last stageof sedimentation from the collapsing turbulent, suspended-load region.

As with the deposits of PH1, grains from PH2 are subangular tosubrounded, and consist dominantly of juvenile volcanic clasts. This

Table 2The lithofacies have been grouped into Facies Associations (FA) according to common occurrences, depositional features, and compositional similarities. These are discussed in detailin 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 theTRC1 deposits on the north side of the complex.

PH2: Normal graded, massive-to-diffusely stratified sequences with occasionalzones 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 TRC2crater on east side.

PH4: Cross-bedded depositsT14, T15, LT14, LT15, LT16

T17, T18, LT18PH5d — These are interbedded and alternating beds found in the southern part ofthe complex.M1: Spatter, lava lake deposit, and radially intruding dikes within tuff

cone walls of TRC1.T19, T20, T21, T22

M1PH5e — Found filling in large scours— likely closely related to first 20 m of thePH5a deposits.

LT12, T6, T7, T8

PH5f—Massive tuffs, lapilli tuffs, and tuff breccias found within the large dune formsthroughout the east and southeast quadrants

TB2, T16, LT17


observation, combined with the lack of highly abraded and inter-bedded non-volcanic clasts and sediments, suggests that the densitycurrents were eruption-fed, and that grain size andmorphology of theclasts reflect eruptive conditions at the vent (White, 2000). Additionalcoarse tephra dispersed through the water column, either by falloutfrom tephra jets or from dilute density currents traveling across thesurface of the water, may have caused ash and lapilli to becontinuously rained into the subaqueous flows (White, 2000). Thisprocess may have suppressed the formation of distinct thin layering(Lowe, 1988; Arnott and Hand, 1989; White, 2000), and could furtherexplain 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 lakesediments overlie the deposits of PH2 (representing the end of theTRC1 eruptive sequence distal from source). This observation,combined with the lack of subaerial features such as base surgedeposits and well-sorted strata from fallout, further supports thesubaqueous depositional environment, and suggests the lake existedwell after the eruption ceased.

2.1.5. PH3: scour and fill depositsPH3 consists of channel-shaped, scour-and-fill deposits within

the FAs of PH1 and PH2. The lower contact is invariably erosiveinto underlying substrate, and is filled with a variety of lithofaciesincluding massive, poorly sorted tuff breccias (TB1), and variouslapilli tuffs that fill in large scour features (LT7, LT8, and LT9; seeTable 1 for description of these lithofacies; also see online sup-plementary data). Pervasive soft sediment deformation such as sagsand flame structures are common, and fine-grained beds on thechannel walls are often deformed into small convolute folds towardsthe axis of the channel.

2.1.6. PH3 interpretationThe scour and fill features of PH3 are interpreted to be the result

of erosive subaqueous sediment gravity flows, and are primarilyfound N1 km from the vent on the flanks of the growing tuffcone platform. The highly deformed and occasionally folded beds

that fill the large scours suggest that the sediment gravity flowswere saturated, and therefore likely emplaced subaqueously anddeformed by the weight of the subsequently deposited overlyingmaterial. The sediment gravity flows could have originated fromone of two sources: (1) remobilization of tephra avalanching downthe steepening slopes of the outer tuff cone as large quantities ofnew sediment were added to the system from the erupting vent;or (2) high-concentration, eruption-fed density currents derivedfrom column collapse. The fine strata that fill the scour features arelikely the deposits of the waning sediment gravity flows, or fill fromsubsequent currents and fallout.

2.1.7. PH4: cross-bedded depositsPH4 was found only proximal to the TRC1 tuff cone, above 1451 m.

PH4 deposits more distal from the tuff cone are not exposed, andeither were eroded, or never deposited. PH4 deposits are dominatedby wavy- to cross-stratified tuffs and lapilli-tuff beds (T17, T18, andLT18; Tables 1 and 2; also see online supplementary data). Dunes aresymmetrical, have wavelengths from 10–20m parallel to the directionof flow, and amplitudes of 0.5–1 m. Individual beds within a singledune consist of one of the aforementioned lithofacies (e.g., a dunewith a 1 m amplitude may be composed of multiple, 10–20 cm layersof 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 contactbetween PH2 and PH4 is not exposed, but is constrained to within 10vertical meters.

2.1.8. PH4 interpretationBased on the cross-stratified nature of these deposits, PH4 is

interpreted to be the result of dilute density currents flowing across asubaerial, gently-sloping platform. These were determined to origi-nate from the TRC1 eruption based on location and dip inwardstowards the TRC1 vent, and consistency in dip with the underlyingTRC1 deposits. The presence of preserved base surge deposits suggeststhat in this part of the stratigraphy, tephra was being deposited abovethe level of the lake. The stratigraphy therefore indicates that the level

ig. 3. a) Generalized stratigraphic column representing PH1 (composed of lithofacies T1, LT1, and LT2). The dark blobs represent the spatter-bomb clasts which range frommediumpilli 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,encil for scale); e) LT2 (photograph taken in section SW1, 0.6 km from source).


Flap

of the lake at the time of the TRC1 eruptionwas approximately 1450mabove sea level.

2.1.9. M1: spatter deposits and solidified lava lakeThe 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) isfound 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 interpretationOver the duration of the TRC1 eruption a large, symmetrical

tuff cone was built above the surface of the lake water. This coneas seen today is ∼1530 m in diameter at the base, and ∼360 m indiameter at the top. M1 is interpreted to represent a final,magmatic stage of the TRC1 eruption, in which the vent of thegrowing tuff cone was gradually isolated from interaction withexternal water, resulting in a fire-fountaining, Strombolian stage(cinders and spatter deposits), and a final effusive stage whichformed a crater filling lava lake. The dikes appear to radiate out

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, person1.5 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 conewalls, which have now been eroded away.

2.2. TRC1 thin section

Petrographic analyses of samples from the TRC1 eruption (includ-ing all FAs) show that the grains are composed dominantly ofpalagonitized glass shards of varying sizes. Some glass has elongated,swirly, and fluidal textures, but most are angular, fractured, and blockywith little evidence of rounding and abrasion. Many of the largerglass shards (N300 µm) have tiny, round, thick-walled vesicles, whichare on average 20 to 80 µm in diameter, but can be found up to240 µm in diameter. These vesicles compose 1% to 26% of the glassgrain; however, coalesced bubbles are rare. Subhedral to euhedralphenocrysts of plagioclase, orthopyroxene, and olivine are commonwithin the glass fragments, and are also found broken and occa-

sionally abraded within the fine ash matrix. The altered, glassy, fine-ash matrix contains features similar to mud cracks (on the micronscale).

Accidental material consists of (1) subrounded to rounded, crystal-line basalt lava flow clasts with a plagioclase-rich groundmass inwhichthe elongated microlites and phenocrysts are aligned; (2) rounded,highlyweatheredmudstone clasts,which in some locations contain silt-sized particles of altered glass in the matrix; and (3) silicic pumice withelongated, stretched, thin vesicles walls. The silicic pumice are identicalin composition and micro-scale texture to samples of the NV1 depositsjust belowPH1of TRC1 (Table 1). Onaverage, samples fromTRC1 containless than 5% accidental clasts (detailed petrographic analysis available asonline supplementary data). The subaerial deposits (PH4) contain 1–4%armored and accretionary lapilli, but overall the components do not varysignificantly from the beginning to the end of the pyroclastic sequence.Details of the spatter and lava flow deposit can be found as online

Table 3This 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. Themaximum 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 80–300 m 19° 80–200 m 10° 300–500 m 19° − –

Parallel 80–150 m 30° 80–120 m 33° 20–200 m 27° 20–100 m 28°


supplementary data, and compositional data can be found in Heikenet al. (1981).

2.3. TRC2 Facies Associations: descriptions and interpretations

2.3.1. PH5: long-wavelength dune bedformsThe TRC2 deposits originated from the eruption that formed the

large depression and tuff ring in the north–northeast part of thecomplex (Fig. 2). The TRC2 eruption has only one FA, PH5, which

Fig. 5. (a) TRC2 dune form above section W4 in the western TRC. Flow direction is oblique tphoto was taken from an adjacent ridge to the south, and the flow direction was from right teast (right). Person 1.5 m circled for scale in (a) and (b); (c) TRC2 dune in the eastern sector ofthe crest, the strata dip away from the crest. This dune is parallel to flow direction (flow di

contains 26 lithofacies (Tables 1 and 2; also see online supplementarydata). There are two scales on which the deposits of TRC2 must bedescribed. The first focuses on the large features (tens to hundreds ofmeters), and the second focuses on individual strata (centimeters todecimeters).

On the large scale, the TRC2 deposits consist of long-wavelength,three-dimensional, symmetrical dune structures with wavelengthsthat vary from 20 to 200 m in the direction parallel to flow and from80 to 500 m perpendicular to flow. Table 3 presents the range of dune

o the plane of the photograph; (b) Large lobate feature in the TRC2 surge deposits. Theo left. Note that the strata dip to the west (left), shallow in a trough, and then dip to theTRC. The crest of this dune is on the right side of the photograph. To the right and left ofrection from left to right), and is close to 200 m in length; (d) trace of dune in (c).

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 regionsaround the TRC2 tuff ring (west–northwest, north–northeast, east,and southeast). Wavelengths in the east tend to exceed those in otherregions. Dips parallel to flow tend to exceed dips perpendicular to flow(30° vs. 10°–20°), with similar values in all regions. Dune wavelengthsin the direction perpendicular to flow exceed wavelengths in thedirection parallel to flow, consistent with the dip data. The amplitudeof the dunes are consistently approximately 1/10th of the wavelengthparallel to the direction of flow (Fig. 5).

The dips shallow towards the middle and tops of the waveforms tob1°, often forming topographic saddles and flat crests. Similar featureswere noted at El Chichon and described as transverse, sinuous ridgesup to 100 m in length with horizontal form index (breadth/wavelength,or in our case perpendicular wavelength/parallel wavelength) rangingfrom 1 to 10 (Sigurdsson et al., 1987). The horizontal form index for theTRC2 dunes ranges from 1 to 5, with the longest perpendicular flowdirectionwavelength equal to 500 m. The difference between the TRC2

and El Chichon dunes is that the TRC2 dunes are symmetricwith dips onthe lee and stoss sides of 27–33°, whereas the El Chichon dunes areasymmetrical with steep stoss and gently sloping (20° dip) lee sides.

The deposits of PH5 also include features such as (1) largeU-shaped channels (10–32 m deep; Heiken, 1971; Heiken et al.,1981) where the steeply dipping deposits along the channel walls areplastically deformed, slumped and folded towards the channel axis(PH5e; Tables 1 and 2, also see online supplementary data); (2) Large-scale chute-and-pool features (up to 13 m tall); and (3) depositsplastered up and around pre-existing obstacles. Also, each set of dunestruncate pre-existing deposits and/or are truncated by later deposits,which results in large-scale hummocky cross-stratification. Thesefeatures will be discussed in more detail in Section 3.3 below.

On the small scale, each dune consists of a range of lithofacies, themost dominant being centimeter to decimeter thick, laterallycontinuous, wavy-planar strata that are internally massive, poorly tomoderately well-sorted tuff and lapilli tuff beds. The large dunes also

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- tomedium-grainedash, and contain abundant accretionary and armored lapilli, vesicu-lated tuff, and pervasive soft sediment deformation. The grains consistdominantly of juvenile scoria, but also contain a higher percentage ofaccidental clasts than those of TRC1 (35–55%). Accidental lithics aredominated by platy, aphanitic basalt, and to a lesser degree by angular,vesicular basalt and mudstone clasts (same as found in deposits ofTRC1). Grain types includes 10–40% accidental basalt, 5–15% mud-stone, and 45–85% juvenile clasts and matrix.

Where exposed, the base of PH5 has a sharp and erosive contactwith underlying TRC1 deposits. However, in a few areas up to 50 cm ofdiatomitic lake sediments exist between TRC1 and TRC2 deposits, atthe contact between FAs PH2 and PH5. Facies Association PH5 consistsof several subgroups (PH5a–PH5f; Tables 1 and 2, online supplemen-tary data).

2.3.2. PH5 interpretationPH5 is interpreted to be the result of multiple, high velocity, column

collapse-induced, dilute pyroclastic density currents (e.g., Fisher andWaters, 1970; Schmincke et al., 1973). These currents, also known aspyroclastic base surges, are unsteady, density-stratified gas-particulatecurrents where turbulence is the dominant particle transport mechan-ism (Fisher andWaters, 1970; Crowe and Fisher, 1973; Sigurdsson et al.,

1987; Valentine, 1987; Druitt, 1992; Wohletz, 1998). The variation indepositional characteristics from one bed to the next within the largedune features (i.e., fines rich, fines poor, massive, stratified, inversegraded, non-graded), and lack of distinct or discernible vertical beddingpatterns between various internal strata, reflect spatial and temporalvariations in bed loadflowdynamics as the deposits aggraded vertically.

Deposits consisting of laterally continuous beds with grain align-ment likely reflect laminarflowconditions in the depositional region ofthe current (Druitt, 1992), whereas deposits consisting of massivedeposits indicate inertial grain flow (Wohletz and Sheridan, 1979).Deposits consisting of thin inversely graded strata suggest tractioncarpet transport in the depositing layer of the bed load (Lowe, 1982;Sohn, 1997), whereas cross-stratified dunes and antidunes indicateturbulent flow during sedimentation (Valentine, 1987; Druitt, 1992), orenergetic, turbulent sweeps through abasal granularfluid, as suggestedby Brown et al. (2007). Therefore, variations in small-scale bedformfeatures are interpreted as a consequence of temporal unsteadiness inthe flow, similar to the interpretation for isolated dune features withinthe distal Mt St Helens blast deposits (Druitt, 1992).

Where present, the small-scale dune features within the largerTRC2 bedforms are dominated by antidunes, which suggest that theflow velocity exceeded internal wave speed in the bed load at the timeof deposition, and that standing internal waves developed within thestratified flow (Crowe and Fisher, 1973; Hand, 1974; Allen, 1982;Valentine, 1987). The same argument could be made for the larger,

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 therectangular 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 alsoimplies standing waves within the broader context of the overall bedload, and therefore possibly supercritical flow conditions. In thisregard, we interpret the larger dune forms (20–200mwavelengths) inthe TRC2 deposits to reflect internal waves in the thick bed load regionof the greater flow, and the smaller-scale features as the result of layer-by-layer deposition and variations in internal waves and flowdynamics within the lowermost depositional bed load region.

The average clast size most proximal to the vent (b1 km) is less than0.5 cm. However, the particle sizes increase to an average grain size of1–5 cm with distance from source, suggesting maintained high flowvelocities and current competencewithdistance fromsourceup to 1.5 km,and thereafter declining current competence up to 3 km distance wherethe average grain size again decreases to b1 cm. Abraded, block-sizedclasts are found intermittently at bedding horizons, and are interpreted asballistic clasts that were entrained into the flow after impact, as it isunlikely that block-sized clasts remained in suspension over the distanceof the current. Their abraded nature also suggests that the block-sizedclastswere tumbled or bounced along via saltation at the base of the flow.

Other than the large-scale dune forms, additional features whichattest to the high velocity of the base surges include deep U-shapedchannels in the northeastern part of the complex, the 13 m tall chute-and-pool feature observed in TRC2, and evidence for currentssurmounting and depositing across 21 to 45 m tall obstacles N2 kmfrom source (discussed further in Section 3.3). The presence ofaccretionary lapilli, vesiculated tuff, steam-escape structures, plas-tered beds, and pervasive soft sediment deformation suggests that

these currents contained a significant amount of liquid water duringtransport and deposition (Wohletz and Sheridan, 1979; Allen, 1982;Wohletz, 1998).

2.4. TRC2 thin section analysis

Twelve samples were taken from various strata throughout thePH5 sequences (detailed descriptions of these samples are availableas online supplementary data). Petrographic analysis shows thatgrain morphology within the matrix is consistent in each of thesamples, and varies only in grain size depending on the coarseness ofthe deposits they were collected from. Grains consist of subangularto rounded, clear-brown to dark and sub-opaque (altered) brownglass with 2–25% rounded vesicles. The glass shards are blocky andhave fine fractures running across their surfaces. Vesicles within theglass shards are rounded, have thick bubble walls, and range from20–100 µm in size. Plagioclase, orthopyroxene, and olivine are alsopresent as individual grains in the matrix. They are also angular tosubrounded, except when found as phenocrysts within the glassgrains where they are subhedral to euhedral. The only new grainwithin these deposits (i.e., not found in the TRC1 samples) is a dark,dense, most commonly irregularly shaped, but occasionally sub-rounded juvenile grain. It differs from the glass grains in that itcontains 10–18% plagioclase needles that are dispersed throughoutthe sample, rather than found as radiating clusters. The dense,altered matrix of the new juvenile grain, which contains 2–5%, andrarely up to 25% rounded vesicles, composes the rest of the juvenile


clast. Accidental material within PH5 consists of the same sub-rounded to rounded crystalline basalt, rounded, highly-weatheredmudstone, and silicic pumice that were found in the TRC1 deposits.Most grains, especially those larger than 300 µm, have a thin fine ashcoating.

The grains within these samples have very similar textures andmorphology to the TRC1 eruptive products; therefore, it is not possibleto tell if the glass shards in these samples were derived from juvenilematerial from the TRC2 eruption, or entrained from the deposits ofTRC1.

Fig. 9. Measured stratigraphic sections of N4 and NE1 (Fig. 1 and 10) and associated FAs. N4

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 withspeckled fill represent the approximate locations of the various vents.These are surrounded by dotted lines which represent the approximatelocation of the inferred crater rims. Numerous stratigraphic sectionswere measured in four sectors around the vent (West, North, East, andSouthern arm) in order to reconstruct the temporal evolution of theeruption aswell as visualize facies variationswith distance from source.

, the western-most measured section is on the right, and the eastern-most on the left.


3.1. Western TRC exposures

Four vertical sections were measured in the western sector of thevolcano; SW1 at 0.6 km from TRC1 source, W4 at 2.4 km from TRC1source, W8 at 2.6 km from TRC1 source, andW10 at 3.4 km from TRC1source (Fig. 7d, c, b, and a, respectively). The sections reveal aprogression from less than 10 m of PH1 (first set of eruption-fedturbidity current deposits) to a few tens of meters of PH2 (second setof eruption-fed turbidity current deposits). PH3 (TRC1 scour and filldeposits) appears in only one section, near the base of sectionW8, butwas traced to more distal locations in the field. PH4 (TRC1 surgedeposits) is found only proximal to source, andwas likely eroded awayin other locations, replaced by the later deposits of TRC2-PH5, or neverdeposited in the first place. The stratigraphic sequence (other than inareas proximal to the TRC1 tuff cone) is overlain by up to 45 m ofPH5 (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 olderlake sediments, the paleo-lake at the site of the TRC1 vent was roughly60–70 m deep.

The contact between the lake sediments and pyroclastic deposits isexposed in sectionsW8 andW10 (Fig. 7a, b). The sediments dip inwardsbeneath thepyroclastic section at 30–40° (Fig. 8a), and formclastic dikesand flame structures into the overlying pyroclastic sediments of TRC1(Fig. 8c). Thin (b3 cm), interbedded, lenses of NV1 lake sediments orlarge (up to 40 cm long), intact pieces of the NV1 substrate (Fig. 8b) arecommonly 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 kmfrom the TRC1 source (Figs. 2 and 9), were measured on the northernside of the complex. Both sections begin much lower in elevation thanthe western sections, suggesting a deepening of the lake to the north.

The two most interesting features are the increased number ofballistics in the N4 section, and the finer grain size (average medium-to-coarse ash) in the NE1 section. The ballistic clasts in the N4 sectionare on average 0.3–0.5m in diameter, but occasionally up to ∼1m, andall have deep bomb sags (e.g., at base of the stratigraphic column, a97×66 cm accidental clast of basalt deforms the underlying strataN1.5 m, Fig. 9b). Ballistic clasts comprise 5–7% of clasts, are foundrandomly scattered throughout the deposits and are present in muchhigher proportions than in more proximal sections. Dewateringfeatures such as squeeze-up and flame structures (0.1–1 m) anddiscontinuous offsets and faults (usually N1 m in length) are alsocommon in this section. A 40 cm thickness of NV2, the non-volcaniclithofacies between the TRC1 and TRC2 deposits, was found about80 m southwest of the N4 outcrop.

Fig.10 is a schematic fence diagram for the north face of the complexover a cross section shown in Fig. 2. Stars indicate where the crosssection bends (Fig. 10). The fence diagram extends from east (left) towest (right; Fig. 10a). On the east side of the complex, the contactbetween NV1 and PH1 is exposed at lower elevations (approximately1326 m a.s.l.), suggesting that the lake deepened to the east.

The TRC1 deposits are grouped in Fig. 10 to illustrate their lateralextent and relationship with the overlying TRC2 deposits. The contactbetween the TRC2 and TRC1 deposits is irregular and varies in elevation,which we attribute mostly to scouring by TRC2 surges. However, NV2lake sediments were found in several locations in thewest (40 cm thicknear section N4), and in the east (east of section NE1), indicating thatthe TRC1–TRC2 contact is not erosive in all locations.

Three smaller flank vents, FV 1, FV 2, and FV 3 (Fig. 10b and c; firstrecognized by Heiken, 1971), cut through the TRC1 deposits but areoverlain by the deposits of TRC2. This indicates that the flankeruptions 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

∼250m in diameter; and the crater of FV 2 is ∼190m in diameter witha surrounding tuff ring ∼440 m in diameter. The lateral tuff ringdeposits have been eroded away, and the exposed outer tuff ringdeposits are highly weathered and form inaccessible cliffs makingdetailed stratigraphic sections impossible. FV 3, previously namedvent 8 by Heiken (1971), is also located in the northern section. It isa small flank vent, only 100 m at the base and 200 m in diameter atthe top, and is well exposed in the cliffs of the northern face.

Three large U-shaped channels were identified in the northeasternside of the complex (location shown by the three arrows in Fig. 2; alsofirst recognized by Heiken, 1971). The channels, which scour the TRC1deposits, are located close to themeasured section of NE1. The upstreamside 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 isroughly 30m further from source, is 13m tall,10mwide at the base, and35 mwide at the top (Fig. 11d). Similar increases in channel dimensionswith distance from source have been noted at Koko crater, HI (Fisher,1977) and Barcena Volcano, Mexico (Richards, 1959). Near the westchannel, parallel to the flow direction the strata are observed to plasterup and over the pre-existing TRC1 deposits (Fig. 11c), and drape theobstruction perpendicular to direction of flow (Fig. 11a).

3.3. Eastern TRC complex-PH5

The eastern flank of the complex is dominated by the surgedeposits of TRC2. The TRC1 deposits are poorly exposed in one smallarea in the east, and another in the southeast (Fig. 2). TRC1 depositswere either eroded by the base surges of the TRC2 eruption, or wereeroded by non-volcanic processes prior to the TRC2 eruption.

The hummocky topography created by the PH5 deposits is mostobvious in the eastern side of the complex. A well exposed outcrop inthe east, beginning with the diamond labeled E4 on Fig. 2, andextending 500 m to the south, exposes ∼40 vertical meters of sectionin the direction perpendicular to flow (Fig. 12a). The corresponding500 m long dune truncates pre-existing TRC2 deposits on the north(right) side (Fig. 12c); and is truncated by another large dune on thesouth (left) side (Fig. 12b). Fig. 12d represents the direction parallel toflow, and shows two 40–50 m long dunes that are part of the largerfeature. These dunes dip ∼13° to the southeast, consistent with the dipshown in Fig. 12c (crest of dune on right side of photograph). Thesephotographs illustrate the three-dimensionality of the dune forms.

What is interpreted as a large chute-and-pool structure, firstmentioned in Section 2.3.1 above, is found at location E5 (Fig. 2). Thisexposure is 1.6 km from the source, extends 40 m parallel to thedirection of flow, and is 13 m tall (Fig. 13). The deposits parallel to theflow direction contain wavy-planar and horizontal strata for thefirst ∼35 m of the outcrop. These strata range from 10–30 cm thick,and are composed of alternating beds of matrix supported (up to100% fine ash beds) and poorly sorted, grain supported beds. Thestrata begin to bend upwards at angles of 33 to 44°, where the bedsabruptly thin and fine to an average 2–7 cm thick (Fig. 13). The stratacontinue to steepen and thin towards the downstream flow directionto the full outcrop height of 13 m before they begin bending backtowards horizontal (Fig.13). The rest of the downstream side of thisfeature has been eroded away. The exposed flat-lying strata are 4 mthick on the left side of the feature (but are probably closer to 6 m inthickness given the exposed strata a few meters further to the east),and are overlain by slightly thicker and somewhat more diffuselystratified deposits which also cover the steeply dipping strata to theright. Although poorly exposed, the upward bending strata can betraced N100 m north of this feature, at the same elevation anddistance from vent.

Chute-and-pool features are commonly observed in base-surgedeposits, and represent a hydraulic jump where supercritical flow(chute, Fr N1,Fr = V=

ffiffiffiffiffiffigh

p, where V=velocity, g=gravitational accelera-

tion, and h=depth of flow) abruptly changes to subcritical flow

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 stratawith thinwhite lines. The top of FV 3 is 200m across, andthe 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 aand 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 typicallymuch smaller in scale (meters rather than 10s of meters). Similarlarge chute-and-pool features were identified in the 1991 Pinatubodeposits and at the base of Mt St Helens in the 1980 pyroclasticflow deposits, but have since been eroded away (Steve Self, personalcommunication). However, structures as large as this have neverbefore been identified in basaltic hydromagmatic eruptions.

An alternative interpretation for this feature is a surge current thatsurmounted a pre-existing obstacle that is no longer exposed, although

there is no evidence of any obstruction in this region (i.e., no remnanttuff rings or older deposits). Furthermore, given that the morphology ofthe feature is the same as previously identified chute-and-pool deposits(JoplingandRichardson,1966; Fisher andWaters,1970; Schmincke et al.,1973) and that the ratio of deposit thickness (downstream to upstreamside of the jump, 2.2) is similar to other well documented chute-and-pool features (Schmincke et al.,1973;Weirich,1988), the chute-and-poolinterpretation is preferred. This indicates supercritical flow conditionsup to 1.6 km from source.

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 toflow, 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 atthe 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 thecomplex is composed primarily of bedded tephras from one ormore ofthe surrounding flank vents, and is overlain by the deposits of PH5. Atleast four obvious flank vents exist in this area (FV 4–7, Fig. 2). Flankvents 4 and 5 are located in the west–southwest part of the complex(Fig. 2) and represent two highly eroded inner craters, which arepartially nested within each other.

Strike and dip data suggest that another sizeable crater may haveexisted in the southern part of the complex (FV 6; Fig. 2), but thedeposits have since been eroded away beyond confident recognition.Based on distinct differences in accidental components and deposi-tional characteristics, the first 15 m of strata along the southern armwere determined to have originated from vents other than TRC1 orTRC2, and are likely the products of either FV 6 or 7. The upper ∼60 mof bedded tephra deposits in the southern arm consist of strata fromFA PH5 of the TRC2 eruption, and contains dunes 20 m to 80 m in

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 ofthe 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.

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

FaciesAssociation

Lithofacies 0.6 km(SWI) maxthickness(cm)

2.4 km(W4) maxthickness(cm)

2.7 km(WI0) maxthickness(cm)

3.5 km(NEI, N4) maxthickness(cm)

PHI LT2 32 10 10 6PH2 LT4 25 70 40 50PH2 LT3 25 50 30 50PH2 T2 10 10 10 20PH2 LT6 0 100 80 0PH2 LT5 0 80 40 0


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

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 freshwaterlake sediments and the subaqueous–subaerial pyroclastic contact.Where exposed, the NV1 lake sediments dip toward the vent beneaththe pyroclastic section at 10–30°, soft sediment deformation struc-tures are pervasive, and clastic dikes varying from decimeters up to6 m in length and 1–100 cm in width are common. These featuressuggest that the NV1 sediments were unconsolidated and saturated at

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

Fig. 14. Wavelength versus distance from source for several well-documented surgedeposits (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 withdistance from source are shown on this plot, thus the total runout distance is notrepresented for each example.


The beginning of the pyroclastic section is dominated by eruption-fed aqueous density currents, some of which probably transformedafter initial fallout. These deposits were interpreted to have beeneruption-fed rather than reworked due to the dominance of juvenilematerial (N95%), general consistency in the bedding style, and lack ofsignificant grain abrasion and rounding. Furthermore, the repeatingpackages of lithofacies (i.e., LT1–T1, LT2, LT4–LT3–T2) indicate that thiswas a pulsating eruption, and each repeated lithofacies or set oflithofacies (i.e., LT6–LT5–T3) represents one explosive pulse. Indivi-dual, fine-ash glass shards within the matrix are fractured, blocky, andangular, indicating that the dominant mechanism of fragmentationwas magma–water interaction (Heiken and Wohletz, 1985).

White (2000) argues that although subaqueous eruption-fedflows commonly involve water-supported transport, the transportand depositional processes are controlled by the nature of theeruption and its interactionwith the surrounding water. The contrastin the bedding style and thickness, the juvenile grain morphologybetween individual lithofacies, and the considerable differencesbetween FAs PH1 and PH2 demonstrate this well; both FAs weredetermined to have been emplaced subaqueously, however, weattribute the differences in deposit characteristics to differences ineruptive style.

PH1 begins with thick, alternating sequences of tuff and lapilli tuffthat often contain either lenses of the underlying lake sediment, or insome cases large, elongated, and intact slabs of the substrate (sectionW10; Fig. 8b), suggesting that the first deposits were dominated bylateral erosive transport. The sharp transition to LT2 deposits, whichare repeating 10–30 cm thick beds of reverse-graded coarse ash tolapilli, followed by finely laminated fine ash, are interpreted as acombination of fallout-initiated, thin density currents, and water-settled, fine ash tuff. While the initial 5–8 m of PH1 deposits representthick, concentrated flows, the more thinly bedded deposits of LT2suggest that the eruption eventually attained a higher frequency ofjetting or explosivity.

PH1 contains coarse lapilli to block-sized juvenile scoria withquenched rinds. These coarse juvenile grains are found eitherrandomly dispersed throughout the T1 and LT1 lithofacies andsupported in a fine-grained matrix, or in concentrated horizons inthe reverse-graded strata of lithofacies LT2. While it has beendetermined based on ash morphology and texture that magma–water interaction was the dominant fragmentation mechanism, thefluidally-shaped, scoriaceous juvenile coarse lapilli-to-fine blockssuggest a concurrent fire-fountaining stage at the beginning of theeruption (Houghton and Schmincke, 1986; Mueller and White, 1992).As all evidence points to magma interacting with an abundant sourceof lake water, the early phase of fire-fountaining may suggest an

initially high magma mass flow rate at the onset of the eruption thatcould have isolated some of the erupting magma from the externalwater, 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 fluidaljuvenile clasts suggests that the flows which produced these depositscontainedwater as the continuous intergranular phase rather than hotgas, which is consistent for deposits of Surtseyan-style explosiveeruptions (White, 1996; Sohn, 1997; White, 2000). In contrast, hot gasis speculated to be the interstitial fluid in high-temperaturepyroclastic flows into subaqueous environments when the flows arederived from sustained explosive eruptions with well-developed gasthrust regions (Sparks et al., 1980; Cas andWright, 1991; Kokelaar andBusby, 1992; Schneider et al., 1992; White and McPhie, 1997; White,2000). Therefore, PH1 deposits likely originated from eruption-fedaqueous density currents including combinations of low- and high-concentration turbidity currents (Lowe, 1982; Postma et al., 1988;Kneller and Branney, 1995), cohesionless sediment gravity flowsor grain flows (Lowe, 1976; Postma, 1986; Nemec, 1990), and a con-current fallout component. As only the coarsest juvenile clasts show afluidal nature, we suggest that the larger clasts remained insulated byself-generated steam jackets within a flow in which water was theinterstitial fluid phase (White, 2000).

The PH2 sequence is interpreted to be the result of multiple, high-density, eruption-fed turbidity currents, where the overall finingupward nature and progressive development of different beddingstructures (i.e., normal and reverse grading, diffuse stratification, fine-grained interbeds) in the deposits represents waxing and waning flowconditions over the duration of deposition (Bouma,1962; Nemec et al.,1980; Lowe, 1982; Postma, 1986). This repeating set of lithofacieselucidates the pulsating nature of the eruption, and the variation ingrain size and depositional characteristics from one package toanother represents variations in initial particle concentration, grainsize, and velocity of each flow.

The subaqueous deposits thicken away from vent, consistent witha growing platform with shallow slopes (1–7°) into deeper water.Table 4 shows the variation of individual lithofacies thickness withdistance from vent. The coarsest deposits, LT6 and LT5, are notrecognized in either proximal or distal locations from source. It isinterpreted that the initial high-density turbidity currents carryingthe coarser-grained loads of LT6–LT5 had a higher initial capacity andcompetence to carry the coarse clasts N0.6 km away from the ventbefore sedimentation occurred. The lack of these lithofacies at themost distal exposures suggests that the coarsest grains settled out atmedial locations. This also suggests that, while we have distinguishedLT6–LT5–T3 and LT4–LT3–T2 stratigraphically, the deposits of LT4–LT3in the distal regions may represent a lateral facies change from thecoarser-grained LT6–LT5 currents.

It is also interesting to note that the deposits of LT4–LT3 thickenwith distance from source, with the thickest deposits occurring at adistance of ∼2.4 km. These deposits subsequently thin to a distance of2.7 km. This likely represents a low initial rate of deposition, followedby an increasing rate of sedimentation with distance from source,which finally tapers off with decreasing current load at distallocations. Note also that the most distal deposits of LT4–LT3 thickenagain, and then follow the thinning trend of LT6–LT5 (Table 4), whichalso supports a lateral facies change from LT6–LT5 to LT4–LT3.

Significant soft sediment deformation (SSD) was found throughoutPH1 and PH2 deposits. Some SSD occurs in the steep walls of scour andfill channels, where it is interpreted that thinly bedded, fine-grained filldeposits slipped and formed convolute folds towards the axis of thechannel. Other SSD in the form of flame features and localized foldedstrata appear to be due to the weight of overlying and likely quicklydeposited pyroclastic deposits. SSD is also found beneath ballistic blocks.All of these SSD features attest to the subaqueous and saturatedconditions at the time of deposition. In several distal locations around


the TRC1 remnants, lake sediments of NV2were identified at the contactof the TRC1 and TRC2 deposits, suggesting that the lake level was wellabove the height of the distal TRC1 deposits long after the end of theTRC1 eruption, and that much of the distal deposits may have beeneroded below the wave base. Additionally, this suggests a significantrepose period between the TRC1 and TRC2 eruptions.

Given the distance from the subaqueous–subaerial contact to thesolidified lava lake, the subaerial portion of the TRC1 tuff cone grew atleast 285 m above the lake level. The most proximal, final pyroclasticdeposits recognized before the transition to spatter at the top of TRC1contain similar accidental clasts as in the lower tuff cone, and alsohave similar juvenile glass morphologies including bubble numberdensity and size. This suggests that as the eruption evolved, thelocation of magma–water interaction did not excavate into thesubsurface to entrain deeper, or a higher percentage of, country rockfragments. Additionally, these deposits are highly palagonitized andcontain a high proportion of accretionary and armored lapilli, whichsupports the interpretation that abundant liquid water was present atthe 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 isevidenced by a rapid increase in progressively coarser-grained scoria(1–3 cm) over two vertical meters just below the Strombolian-spatterdeposits. This observation suggests that the vent was quickly sealedoff from the influence of external water to produce the Strombolian-spatter 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 theTRC. The crater rim is ∼2.7 km in diameter, and the deposits, whichform a broad tuff ring surrounding the crater, are dominated by large-scale pyroclastic base surge deposits. The TRC2 eruption occurred asignificant time after the TRC1 eruption, as evidenced by the 25–50 cmof lake sediments found at the TRC2–TRC1 contact. Additionally, thereappears to have been a significant drop in the regional lake level priorto the TRC2 eruption. Primary depositional characteristics of surgecurrents are found as low as 1325 m, which suggests the lake, ifpresent at the time of the TRC2 eruption, was no deeper than 15 m.Wave cut terraces within the TRC2 deposits are present at 1356 melevation, which indicates that at some point after the eruption thelake water rose to at least 46 m deep.

The TRC2 surge deposits drape features to the east–southeast andnorthwest of the TRC1 tuff cone, which are likely well below themaximum height of the original tuff cone walls (Fig. 2). This suggestseither that much of the TRC1 tuff cone was eroded away prior todeposition of the TRC2 deposits, or it was eroded by the TRC2 eruptionitself.

A difficult question to answer is whether or not the TRC2 eruptionoccurred due to magma–water interaction at depth to create a maarand surrounding tuff ring, or if it erupted in a near surface or playa-lake setting to create a tuff ring with the crater floor above the level ofpaleotopography. While the accidental clast component is muchhigher in the TRC2 deposits (between 30 and 55%) than in the TRC1deposits (b5%), it is not quite as high as proportions recognized forother maar deposits where the accidental component composes up to80% of the total tuff deposits (Cas andWright,1987). However, the trueaccidental component cannot be quantified with certainty as it isimpossible to distinguish between the accidental clasts derived fromthe TRC1 deposits and the juvenile clasts of the TRC2 deposits. Thus,30–55% is a minimum estimate for the total amount of accidentalmaterial 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 occurredthrough the flank of the remnant TRC1 tuff cone, and therefore mayhave 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 themagma–water interaction occurred dominantly within the TRC1deposits, and that the source of the water was related to a shallowplaya lake, saturated, unconsolidated or very poorly lithified TRC1deposits, and/or near-surface, water-saturated playa-lake sediments.Therefore the TRC2 feature could be considered to be a shallow maarin the sense that magma–water interaction occurred below the pre-existing surface, and accidental clasts derived from the underlyingcountry rock are common, although subordinate to the juvenile clasts.

The rest of the TRC2 discussion will focus on the large scale surgebedforms that comprise the tuff ring, and the insight they provide intoeruption dynamics. The pyroclastic surge deposits radiate axi-symmetrically from the center of the crater. Vertical cross-sectionsthrough the TRC2 dune features reveal large-scale hummocky cross-stratification. The fact that each set of dunes truncate the previousones, and form the same three-dimensional, hummocky featuressuggests 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 theedifice with the passing of each surge current. These features demon-strate the pulsating nature of the explosive eruption.

The dune forms at TRC2 are more than an order of magnitude largerthan dune features recognized in other basaltic hydrovolcanic deposits(e.g., Fisher andWaters,1970;Waters and Fisher,1971; Croweand Fisher,1973; Schminckeet al.,1973; SohnandChough,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 awayfrom the vent (e.g., 20-mwavelengths recorded4.7 kmfromsource). Theflowswere highly erosive and scoured into the TRC1deposits on all sidesof the tuff ring, truncate earlier surge deposits from the same eruption,and form large U-shaped channels in the north. The channels wereinterpreted by Heiken (1971) to have been carved out by an extremelyerosive surge, similar to theU-shaped channels described at Koko Crater,Hawaii (Fisher,1977).Most U-shaped channels have been recognized onthe 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 thoseat TRC2 indicate much larger and faster surges than those required toform similar features in steeper tuff cones (Fisher, 1977).

Whenwe plot the TRC2 data (wavelength vs. distance from source)along with data from the literature, we see that wavelengths of theTRC2 bedforms are much longer than those recorded at other basaltichydrovolcanoes, and appear to follow the trends of larger scaleeruptions such as El Chichon and Mt St Helens (Fig. 14). At El Chichon,similar three-dimensional features with shorter wavelengths wereobserved at distances greater than 4 km, and the authors suggest thatif the dunewavelengths at greater distanceswere extrapolated towardthe source, they should havewavelengths of 100 to 200 m (Sigurdssonet al., 1987), consistent with those at TRC2. The TRC2 surges areunusual in the sense that wavelengths of this scale are typicallyfound 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 flowtuffs (e.g., Ohakuri-ignimbrite forming eruption in the central TaupoVolcanic Zone; Gravley et al., 2007).

The unusually long wavelengths of the dune features proximal to thevent are indicative of highly inflated currents at the onset of the flows.The most likely way to obtain highly inflated flows proximal to source isto entrain air during rise and collapse of an eruption column (Sigurdssonet al., 1987). Therefore, the base surge currents are interpreted to begenerated by collapse from an unusually high column. This is furthersupported by the runout distance, which, based on the 20 mwavelengths found at the distal exposures, was likely much greaterthan the exposed 4.7 km. By comparison, eruption column collapses atAmbae Island, Vanuatu, on the order of a couple hundred meters,produced surges to distances of 300 m (Nemeth et al., 2006), and the


phreatomagmatic eruption in Karymskoye Lake, Kamtchaka, collapsedfrom 1 km and produced surges to distances of b2 km (Belousov andBelousova, 2001). A general relationship can be made from these dataand weapons test-induced surge currents (e.g., Sedan weapons test;Rohrer,1965); surge runout distance is approximately 1.3 to 1.5 times thecolumn collapse height. The TRC2 surge deposits extend at least 4.7 kmfrom source implying column collapse heights of 3–4 km, which isunusual for basaltic hydrovolcanic eruptions.

On the fine-scale, the deposits within TRC2 are dominated byblocky, angular shards of fine ash, which suggests highly efficienthydromagmatic fragmentation. The juvenile glass shards have thick-walled bubbles, and low bubble fractions (0.01 to 0.25), implying thatmagmatic volatiles did not play an important role in magmafragmentation. The abundant fine-ash fraction combined with thelarge-scale, long runout surge features, are evidence of efficientconversion of thermal to kinetic energy and a highly energeticeruption, respectively. However, the deposits contain ample evidencefor abundant liquid water in the surges during transport andemplacement, which is contradictory, as these features are oftenattributed to water–magma ratios above that which flashes all waterto steam, and thus are associatedwith lower explosivity (Sheridan andWohletz, 1983; Wohletz and McQueen, 1984). Based on all observedfeatures, it is interpreted that the eruption was highly energetic andlikely had an efficient conversion of thermal to mechanical energy,vaporizing most if not all external water at the site of magma waterinteraction. It is then hypothesized that the water vapor subsequentlycondensed due to entrainment-induced cooling during ascent andcollapse from the eruption column to produce the “wet” deposits.

The small-scale, centimeter to decimeter-thick strata within thelarger dune features are interpreted to represent temporal variation inthe 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 dunewavelengths, from the largest that create the hummocky topographyaround the tuff ring, to the medium dunes (up to 50 m) that appearnested within the larger wavelength dunes, to the smallest antidunesrecorded in the individual 1–2m thick packages of sediment (3–6m inwavelength), support the Valentine (1987) model of depositionwithina density stratified current. The various scales of dunes likelyrepresent internal gravity waves on different length scales withinthe larger flow, as discussed in Section 2.3.2 above.

Extensive base surge runout distances have been recognized forother phreatomagmatic eruptions. The Taal Volcano produced basesurges 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 andHiscott, 1997; Brown et al., 2007). The explosive, basaltic eruptionat Taal volcano matches TRC2 closest in terms of runout distance,but the bedform wavelengths at Taal were more than an orderof magnitude smaller (Fig. 14; Moore, 1967). Thus, the TRC2 dunewavelengths expand the range possible for basaltic phreatomag-matic eruptions.

5. Conclusions

TRC is a basaltic, polygenetic volcano that formed during severalstyles of hydrovolcanism, varying from Surtseyan to maar-styleeruption. The first eruption to occur in the complex is representedby a large tuff cone (TRC1), which is the tallest feature in the complex.It began erupting through a 60–70 m deep lake, initially depositingmultiple eruption-fed turbidity currents until a gently-dipping plat-form was built up above the surface of the lake. The hydromagmatic

activity subsequently constructed a steep tuff cone with a craterfloor ∼365 m above the base of the lake.

The beginning of the eruption is interpreted to have had both ahydromagmatic and a simultaneous fire-fountaining stage, however, theeruption changed to an entirely hydromagmatic phase after the initialactivity. The rest of the subaqueous deposits contain a mixture of fineash and lapilli, suggesting moderately efficient hydromagmatic mixing.The subaerial deposits of TRC1 are not well preserved due to post-eruptive erosion, but where present consist of base surge and falldeposits. The focus of explosivity was shallow and did not core into thepre-volcanic substrate, and the final eruptive stages are marked by bothefficient hydromagmatic fragmentation and evidence for liquid water atthe time of deposition, suggesting relatively high water-to-magmaratios. The eruption ended with a transition from hydromagmatic toStrombolian and effusive activity and produced a crater-filling lava lakethat partially intruded the unconsolidated tuff cone walls.

Once the eruption ceased, much of the medial to distal tuff cone wasbelowwave base allowing portions of it to be eroded away, and allowingthe TRC1 deposits to remain saturated long after the eruption. A longrepose period occurred before later volcanic activity resumed.

The seven flank vents around the complex erupted some timeduring or after the TRC1 activity, but before the TRC2 eruption. TRC2 isrepresented by a 2.7 km diameter crater in the northeast of thecomplex. It was the most energetic eruption in the complex, and likelyoccurred when magma interacted with a shallow lake, playa-lake,and/or saturated TRC1 tuff sediments. The TRC2 eruption producedmultiple, highly inflated, erosive pyroclastic surges that radiated outfrom the base of a collapsing column. Dunes in the TRC2 surgedeposits have longer wavelengths than other basaltic hydromagmaticdeposits, and are instead more consistent with dune bedformsproduced by larger eruptions such as the 1980 blast at Mt St Helens,and the column collapse surges at El Chichon in 1982. Such large dunewavelengths are rare for low-aspect ratio tuff rings, and suggest highlyinflated flows produced by an unusually high eruption column.

Acknowledgements

The authors would like to thank Grant Heiken for suggesting thisfield area and spending several days in the field with B.B. We areespecially grateful for the assistance of Mike, Jean, and Josh Bandfield,and Adam Frus in the field during summer 2006. This work greatlybenefited from the thorough and insightful reviews of John Smellieand Corina Risso, and guest editor Károly Németh. Funding for thisresearch was provided by an Arizona NASA Space Grant ConsortiumStudent Fellowship, and the National Science Foundation, USA (EAR0538125).

Appendix A. Supplementary data

Supplementary data associated with this article can be found, inthe online version, at doi:10.1016/j.jvolgeores.2008.10.011.

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