The initiation and development of a caldera-forming Plinian eruption (172 ka Lower Pumice 2 eruption, Santorini, Greece)

Accepted Manuscript The initiation and development of a caldera-forming Plinian eruption (172ka Lower Pumice 2 eruption, Santorini, Greece)

J.M. Simmons, R.A.F. Cas, T.H. Druitt, R.J. Carey PII: DOI: Reference:

S0377-0273(16)30249-9 doi: 10.1016/j.jvolgeores.2017.05.034 VOLGEO 6119

To appear in:

Journal of Volcanology and Geothermal Research

Received date: Revised date: Accepted date:

12 August 2016 20 May 2017 30 May 2017

Please cite this article as: J.M. Simmons, R.A.F. Cas, T.H. Druitt, R.J. Carey , The initiation and development of a caldera-forming Plinian eruption (172ka Lower Pumice 2 eruption, Santorini, Greece), Journal of Volcanology and Geothermal Research (2017), doi: 10.1016/j.jvolgeores.2017.05.034

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ACCEPTED MANUSCRIPT The initiation and development of a caldera-forming Plinian eruption (172 ka Lower Pumice 2 eruption, Santorini, Greece) J.M. Simmonsa, R.A.F Casa, b, T.H. Druittc, R.J. Careyb School of Earth, Atmosphere and Environment (EAE), Monash University, Clayton, Victoria, Australia 3800 b School of Physical Sciences/CODES, University of Tasmania, Hobart, Tasmania, Australia c Laboratoire Magmas et Volcans, Université Clermont Auvergne, CNRS, IRD, OPGC, 6 Avenue Blaise Pascal, 63178 Clermont-Ferrand, France

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Keywords: Lower Pumice 2, Santorini, Precursor, Plinian, Caldera, Magma decompression

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ABSTRACT

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The rhyodacitic 172 ka Lower Pumice 2 (LP2) eruption terminated the first magmatic cycle at Santorini (Greece), producing a proximal <50 m thick succession of pyroclastic fall

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deposits, diffusely-stratified to massive ignimbrites and multiple lithic breccias. The eruption

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commenced with the development of a short-lived precursory eruption column, depositing a <15 cm blanket of 1-2 cm sized pumice fragments at near vent localities (LP2-A1). The precursor deposits are conformably overlain by a <30 m thick sequence of reversely-

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graded/ungraded pumice fall deposits that reflect opening and widening of a point-source

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vent, increasing mass discharge rates up to 108 kg s-1, and the development of a 36 km high eruption column (LP2-A2, A3).

The progressive increase in maximum vesicle number

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density (NVF) in rhyodacitic pumice, from 3.2 x 109 cm-3 in the basal fall unit of LP2-A2-1 to

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9.2 x 109 cm-3 in LP2-A3, translates to an increase in magma decompression rate from 18 to 29 MPa s-1 over the course of LP2-A. This is interpreted to be a consequence of progressive vent widening and a deepening of the fragmentation surface.

Such interpretations are

supported by the increase in lithic clast abundance vertically through LP2-A, and the occurrence of basement-derived (deep) lithic components in LP2-A3. The increasing lithic clast content and the inability to effectively entrain air into the eruption column, due to vent widening, resulted in column collapse and the development of pyroclastic density currents (PDCs; LP2-B). A major vent excavation event or the opening of new vents, possibly

ACCEPTED MANUSCRIPT associated with incipient caldera collapse, facilitated the ingress of water into the magmatic system, the development of widespread PDCs and the deposition of a <20 m thick massive phreatomagmatic tuff (LP2-C). The eruption cumulated in catastrophic caldera collapse, the enlargement of a pre-existing flooded caldera and the discharge of lithic-rich PDCs,

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depositing proximal <9 m thick lithic lag breccias (LP2-D).

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1. INTRODUCTION The low recurrence times of caldera-forming explosive eruptions necessitate the need for modern appraisals of pre-historic Plinian eruption sequences and their source volcanoes. The stratigraphy of such volcanoes often document a prolonged history of explosive

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volcanism, associated with a range of explosive phenomena, including precursor, plume,

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ignimbrite and caldera-forming events (e.g., Froggatt, 1981; Walker, 1981a; Walker, 1981b;

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Houghton et al., 2003; Houghton et al., 2010; Houghton et al., 2014; Wilson et al., 1995; Chesner and Rose, 1991; Chesner et al., 1991; Chesner, 2012; De Silva et al., 2015; Simmons Santorini Volcano, in the Aegean Sea, Greece, is one such example,

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et al., 2016).

representing the product of twelve major explosive eruptions (with recurrence periods of 10-

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30 kyr) and five major caldera collapse events (Druitt et al., 1989; Druitt and Francaviglia, 1992; Druitt et al., 1999; Simmons et al., 2016; Simmons et al., 2017). Santorini is one of the

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most intensively studied volcanoes on Earth, infamous for the VEI 6-7 Minoan eruption and

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its part in the displacement of the Minoan society, from 1627 BCE (Bond and Sparks, 1976; Pichler and Friedrich, 1980; Heiken and McCoy Jr, 1984; Druitt et al., 1989; Druitt et al.,

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1999; Heiken and McCoy, 1990; Sparks and Wilson, 1990; Cioni et al., 2000; Pfeiffer, 2001; Sigurdsson et al., 2006; Druitt, 2014; date from Friedrich et al., 2006). The Minoan eruption

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was the largest and most recent Plinian explosive eruption of Santorini, terminating the second of two magmatic cycles. The earlier rhyodacitic 172 ka Lower Pumice 2 (LP2) eruption, thought to be of equivalent magnitude to the Minoan event, concluded the first magmatic cycle, forming one of the early calderas of Santorini. Despite these affinities, detailed research on the LP2 sequence is still in its infancy (e.g., Druitt et al. 1989; Druitt et al.,1999; Gertisser et al., 2009; Cadoux et al., 2014; Keller et al., 2014). Here, we use the stratigraphy of the LP2 deposit sequence, its components and their textures to reconstruct the

ACCEPTED MANUSCRIPT eruption conditions, style of fragmentation (i.e., magmatic versus phreatomagmatic), dispersal and depositional processes, and the evolution of the vent system during one of the earliest recorded cataclysmic events of the Aegean region.

2. SANTORINI VOLCANO

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Santorini is one of five volcanic provinces that collectively define the Hellenic

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Volcanic Arc, resulting from the northward subduction of the African-Arabian plate beneath

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the Aegean microplate plate (McKenzie, 1970; Papazachos and Comninakis, 1971; Mercier et al., 1989; Jackson, 1994). The Santorini volcanic province consists of three volcanic fields,

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including Santorini caldera (e.g., Fouque, 1879; Nicholls, 1971; Bond and Sparks, 1976; Druitt et al., 1989; Perissoratis, 1995; Druitt et al., 1999), the Kolumbo Volcano to the NE

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(Druitt et al., 1999; Nomikou et al., 2012; Kilias et al., 2013; Nomikou et al., 2013; Cantner et al., 2014; Hübscher et al., 2015), and the Christiana Islands to the SW (Druitt et al., 1999;

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Keller et al., 2010). The modern caldera of Santorini is bound by the islands of Thera,

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Therasia and Aspronisi. The intra-caldera islands of Palea Kameni and Nea Kameni record the latest phase of volcanism at Santorini, which involved the growth of a lava flow and

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dome complex (Fig. 1A; Nicholls, 1971; Druitt et al., 1999; Pyle and Elliott, 2006). Santorini Volcano overlies continental crust of the Cycladic massif, which consists of

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non-exposed Proterozoic gneiss (Hercynian basement) and Mesozoic igneous and sedimentary assemblages.

These assemblages underwent high and medium pressure

metamorphism during the Eocene and Miocene, respectively (Alpine Orogeny; Nicholls, 1971; Makris, 1978; Andriessen et al., 1979; van der Maar and Jansen, 1983; Andriessen et al., 1987; Papanikolaou, 1987; Schliestedt et al., 1987; Liew and Hofmann, 1988; Urai et al., 1990; Skarpelis et al., 1992; Lips et al., 1999). Altered Triassic limestone and Eocene metapelites form exposures at Mt Profitis Ilias and Athinios, on Thera, the main island of

ACCEPTED MANUSCRIPT Santorini (Nicholls, 1971; Druitt et al., 1999). Granitoids have also been identified in drill cores near the town of Megalochori, on Thera (Skarpelis et al., 1992). Accessory limestone fragments, marbles, skarns, granitoids and other diverse metasedimentary lithic components preserved in the caldera wall deposits, provide additional insights into the diverse crustal assemblages of the Aegean.

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Early volcanism at Santorini was characterised by effusive to localised explosive

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activity on the Akrotiri Peninsula, producing basaltic to rhyodacitic submarine and subaerial

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lavas, hyaloclastite breccias and other pyroclastic deposits (650-550 ka Early Volcanic Centres of the Akrotiri Peninsula; Nicholls, 1971; Druitt et al., 1999). From 530 ka, a

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stratovolcano, named Peristeria, developed on the northern margin of the present-day caldera, producing a succession of basaltic-andesitic lavas.

This occurred concurrently with

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stombolian style eruptions on the Akrotiri Peninsula (e.g., Cape Balos, Cape Kokkinoperra, Cape Mavrorachida/Red Beach - 450-330 ka; Druitt et al., 1999). Twelve major explosive

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eruptions and at least five major caldera collapse events have occur since 360 ka (Fig. 1A, B;

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Heiken and McCoy Jr, 1984; Druitt et al., 1989; Druitt and Francaviglia, 1992; Druitt et al., 1999). The deposits from these eruptions define two mafic to silicic magmatic cycles that

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ended with the 172 ka Lower Pumice 2 and 3.6 ka Minoan eruptions, respectively (Fig. 1B;

2014).

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Druitt et al., 1989; Druitt et al., 1999; Gertisser et al., 2009; Cadoux et al., 2014; Keller et al.,

In historic times, volcanism has been restricted to Nea Kameni, characterised by effusive lava flows and vulcanian explosive activity. The last documented eruption occurred in 1950 (Druitt et al., 1999; Pyle and Elliott, 2006). From 2011 to 2012, the intrusion of > 1.4x107 m3 of magma beneath Santorini, resulted in inflation of the caldera floor and increased seismicity (Chouliaras et al., 2012; Newman et al., 2012; Kaviris et al., 2015;

ACCEPTED MANUSCRIPT Papadimitriou et al., 2015; Parks et al., 2015). At the time of publication, seismicity is at baseline levels.

Figure 1. A. Distribution of the Thera Pyroclastics on Santorini (including the LP2 sequence; after Druitt et al., 1999). S1, S2, S3 etc. denotes the field locations (e.g., Site 1, Site 2, Site 3

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etc.) studied as the basis for this research (coordinates for each field location are provided in

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the supplementary material). B. Stratigraphy of the Thera Pyroclastics (after Druitt et al.,

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1999). Black bars denote periods of erosion between eruptions. Eruption dates from Druitt et al. (1999), Keller et al. (2000), Friedrich et al. (2006), Lee et al. (2013) and Fabbro et al.

METHODS

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3.

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(2013).

3.1. Facies and stratigraphic analysis

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Stratigraphic analysis of the LP2 sequence was conducted at 13 localities inside the

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caldera wall and at other selected extra-caldera exposures on the southern coast of the Akrotiri Peninsula (Fig. 1). This involved using the facies and terminology approach of Cas

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et al. (2009) to distinguish units based on lithological characteristics (e.g., componentry,

textures, geochemistry), depositional structures (e.g., diffuse bedding, ripples, dunes, cross

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bedding, laminations etc.) and fabrics (e.g., imbrication). Graphic logs recording grainsize, componentry, sorting, grading, clast angularity, bedforms, and contact relationships between successive lithofacies were constructed at each locality. The proportions of juvenile and lithic clast assemblages were determined by counting the number of clasts (>1 cm), of a specified component type, for a defined area of interest (typically 0.5-2 m2 depending on unit thickness).

Fallout dispersal patterns were assessed by recording the thicknesses of

individual fallout units at each field locality, then constructing isopach diagrams. In addition,

ACCEPTED MANUSCRIPT the average maximum diameters of the five largest juvenile and lithic clasts were measured at each locality, and isopleth diagrams were constructed for each fallout deposit.

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thicknesses of ignimbrites and other associated deposits were also mapped, where accessible, to identify variations in unit geometry. Juvenile pyroclasts and volcanic ash were sampled at each major stratigraphic level at Cape Balos West (Site 4) and Fira (Site 9) to identify

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variations, or absence thereof, in texture, morphology, density and chemistry through the

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eruption sequence. Ash morphologies were imaged using a JEOL 7001F Scanning Electron

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Microscope (SEM) at the Monash Centre for Electron Microscopy (MCEM) to characterise the style of magma fragmentation (magmatic versus phreatomagmatic). The composition of

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juvenile pyroclasts (Table 1) was determined using the method detailed in Simmons et al. (2016). Lithic clasts were also sampled and characterised using petrographic analysis, to

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over the course of the LP2 eruption.

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assess the temporal variations in fragmentation depth and the evolution of the vent system

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3.2. Pyroclast density and vesicularity

The densities, total, connected and isolated vesicularities of juvenile pumice

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pyroclasts were measured at the Laboratoire Magmas et Volcans, in Clermont-Ferrand, France. Fifty to one hundred rhyodacitic white pumice pyroclasts (16–24 mm in diameter),

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selected from specified stratigraphic units of the LP2 sequence, were cleaned using distilled water and dried at 100°C for a period of 24 hours. Pumices which exceeded 24 mm in diameter were excluded from study, to minimise the effects of post-fragmentation vesiculation and thermal expansion (e.g., Thomas and Sparks, 1992). The fragments were first weighed using Metler Toledo DB602 electric scales (MP - sample mass). The volumes of pumice (VP) were then determined using glass beads (600-800 µm), as per the method of Silva (2008), which is detailed in Simmons et al. (2017). These parameters define pyroclast

ACCEPTED MANUSCRIPT density (kg m-3). To determine the mean solids density (DSOL - in kg m-3), selected pyroclasts were crushed using a Rocklabs tungsten carbide mill and weighed. The solids volume was then determined using an AccuPyc II 1340 He-Pycnometer. Total vesicularity was then

Equation 1.

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Total Vesicularity (VeT) = 1 – (MP / (DSOL x VP))

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determined using the following equation:

Connected (VeC) and isolated (VeI) vesicularities were subsequently

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determined as per the following equations:

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pycnometery.

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The combined glass and isolated vesicle volumes (VC) were also determined using helium

Equation 2.

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Connected Vesicularity (VeC) = 1 – (VC / VP)

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Isolated Vesicularity (VeI) = VeT – VeC

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Equation 3.

3.3. Analysis of vesicles textures

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Vesicle textures in pumice were imaged using a JEOL 7001F Scanning Electron Microscope (SEM) at Monash University’s Centre for Electron Microscopy (MCEM). Samples from each stratigraphic level and selected density classes (low, modal and high) were impregnated in epoxy and polished at Monash University’s petrographic facility. Vesicle textures were then imaged at four different magnifications (section scan, x25, x100, x500 on our JEOL 7001F SEM) to capture each vesicle size class, as per the method of Shea et al. (2010). These images were processed using the Matlab program, FOAMS (Shea et al.,

ACCEPTED MANUSCRIPT 2010), which uses the stereological conversion technique of Sahagian and Proussevitch (1998) to quantify vesicle volume distributions and vesicle number densities (NV).

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minimum vesicle size diameter of 1 μm was used in the calculation of NV, which was then corrected for vesicularity (NmV). Vesicle number densities were also calculated for vesicle populations with equivalent diameters between 1 and 10 μm (NVF), to facilitate the

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determination of magma decompression rate (Toramaru, 2006; Shea et al., 2011; Simmons et

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al., 2017).

4. RESULTS

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Druitt et al. (1989) and Pyle (1990a) previously defined four key stratigraphic units that commonly occur throughout the LP2 deposit sequence. These include:

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 a very-thick basal pumice fall deposit (LP2-A), overlain by  a diffusely-stratified to massive ignimbrite (LP2-B), which is overlain by

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 a widespread phreatomagmatic tuff (LP2-C), and finally

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 a lithic-rich lag breccia (LP2-D).

Subsequent analysis by Gertisser et al. (2009) delineated a minor precursory deposit

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(LP2-A1) and additional subunits within LP2-A (A2-1, A2-2, A3) and LP2-B (B1, B2). These units disconformably overlie deposits of the earlier 184 ka Lower Pumice 1 (LP1)

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eruption (Simmons et al., 2016; Simmons et al., 2017), cropping out in the caldera wall from Cape Loumaravi to Fira, and irregularly at Oia (Site 10, 11), Cape Perivola (Site 12) and Mt Megalo Vouno (Site 13) in the north. Vertical sections through the stratigraphy also occur at Mesa Pigadia Beach (Site 1) and Kambia Beach (Site 2), on south coast of the Akrotiri Peninsula (Fig 1A). Plinian fall deposits (LP2-A) have also been identified on the island of Anafi, ~31 km east of Santorini (Keller et al., 2014), while centimetre thick tephra deposits (V-1 tephra layer), considered to be the distal deposits of LP2 eruption, occur in deep-sea

ACCEPTED MANUSCRIPT cores throughout the Aegean (Keller et al., 1978; Keller, 1981; Keller et al., 2000). Revisions of the LP2 sequence, conducted as the basis for this study, has defined an additional twelve previously undescribed subunits. These subunits are discussed herein with reference to the established stratigraphic framework of the LP2 sequence.

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4.1. Componentry

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4.1.1. Juvenile Pyroclasts

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Three juvenile clast types are distinguished in the stratigraphy. These include white pumice pyroclasts, grey and banded pumice pyroclasts, and dark cauliform pyroclasts (Fig.

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2A-C). White pumice pyroclasts are dacitic to rhyodacitic in composition (65-71 wt. % SiO2; Fig. 2A, 3 - blue symbols), highly-vesicular (51-90 vol. %) and crystal-poor (< 5 vol. %; 6-8 wt. %; Cadoux et al., 2014), characterised by dispersed

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Gertisser et al., 2009;

phenocrysts of plagioclase, orthopyroxene and minor clinopyroxene (Gertisser et al., 2009;

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Cadoux et al., 2014). Amphibole, ilmenite and magnetite also occur (Gertisser et al., 2009).

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Analysed melt inclusions in minerals have maximum water contents ranging from 4.7 wt. % (Gertisser et al., 2009) to 5.3 wt. % (Cadoux et al., 2014). Grey and banded pumice

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pyroclasts are dacitic to rhyodacitic in composition (63-68 wt. % SiO2; Fig. 2B, 3 - orange symbols), have high vesicularities (42-85 vol. %) and low to moderate crystal contents (5-15

inclusions.

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vol. %), which include plagioclase, orthopyroxene, clinopyroxene and felspathic mafic Cauliform pyroclasts are tabular to oval in morphology and have variably

crenulated surfaces (Fig. 2C). Analysed clasts are basaltic to basaltic andesite in composition (47-56 wt. % SiO2; Fig. 3 - green symbols), dense (~2000-2500 kg m-3) and variably crystalrich (20-40 vol. %), characterised by phenocrysts of plagioclase, olivine, two pyroxenes and iron oxides (Ti-magnetite and ilmenite; Gertisser et al., 2009). Maximum water contents of 4.0 wt. % were measured in olivine-hosted melt inclusions (Gertisser et al., 2009).

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Table 1. Whole-rock major and trace element data. bd - below detection limit. W - White pumice. G - Grey and banded pumice. C - Cauliform pyroclasts.

4.1.2. Lithic Pyroclasts

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A variety of lithic clast assemblages have been identified within the LP2 sequence,

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which have been characterised broadly as either volcanic or basement-derived. Volcanic

sequence.

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lithic components, of all shapes (round to angular) and sizes (<2 m), occur throughout the Noted examples include poorly to moderately-vesicular basalts, crystal-rich

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basalts, aphanitic basalts and variably flow-banded andesite/dacites (Fig. 2D-G). Basementderived lithics are typically smaller (15-20 cm max) and less abundant (5 vol. %). Identified

schists (Fig. 2H-K).

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basement assemblages include granitoids, altered limestone, marbles, sandstones, slate and Approximately 20-30% of clasts, particularly basalts, exhibit

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hydrothermal alteration.

Figure 2. Pyroclast assemblages in the LP2 sequence. Juvenile pyroclasts. A: White

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pumice. B: Grey and banded pumice. Note the alternating bands of light and dark coloured glass. C: Tabular, crenulated, cauliform basaltic pyroclasts, with adhered white pumice.

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Volcanic lithic assemblages. D: Mildly-vesiculated basalt. E: Dense, crystal-rich basalt. F: Aphanitic basalt.

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Flow-banded dacite.

Basement-derived lithic assemblages.

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Granodiorite. I: Altered limestone. J: Sandstone. K: Slate.

Figure 3. Total alkali silica (TAS) diagram illustrating the compositional spectrum of juvenile pyroclasts (white pumice - blue symbols, grey and banded pumice (G/B) - orange symbols, and cauliform pyroclasts - green symbols) from the LP2 sequence. TAS diagram

ACCEPTED MANUSCRIPT from Le Maitre et al. (1989). Alkaline (dark grey) - subalkaline (light grey) boundary from Irvine and Baragar (1971). Symbol shapes (circle and square) represent included data from Druitt et al. (1999) and Gertisser et al. (2009), respectively.

4.2. Stratigraphic and lithofacies architecture

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The LP2 sequence consists of 18 lithological subunits that define five major

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stratigraphic divisions. These include, from the bottom up, a thin, fine-grained, pumice

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lapillistone unit (LP2-A1), a sequence of massive pumice lapillistones (LP2-A2, A3), a series of diffusely-stratified to massive, pumice lapilli-tuffs (LP2-B1-1, B1-2, B1-3, B2), a massive,

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breccias (LP2-D1, D2, D3, D4; Fig. 4, 5).

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poorly-sorted, pumice-lithic tuff (LP2-C1-1, C1-2 , C2-1, C2-2, C3-1, C3-2), and lithic

Figure 4. Stratigraphic unit correlations for the LP2 sequence, showing the locations of

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measured sections on Santorini.

Figure 5. Stratigraphic subdivisions and architecture of the LP2 sequence, showing

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variations in juvenile (white, grey and banded pumice, and cauliform pyroclasts) and lithic

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clast (volcanic and basement lithics) assemblages.

4.2.1. LP2-A - Pumice lapillistones A sequence of moderately to well-sorted, pumice lapillistones, termed LP2-A1, A2-1, A2-2 and A3, are laterally widespread units that occur ubiquitously at the base of the deposit sequence (Fig. 6A, B; Druitt et al., 1989; Druitt et al., 1999; Gertisser et al., 2009; Keller et al., 2014). LP2-A1 is characterised by a clast-supported framework of rhyodacitic white pumice (⌀ - diameter ~2 cm; 14 cm thick near Athinios; Fig. 6C). On the Akrotiri Peninsula,

ACCEPTED MANUSCRIPT LP2-A1 is thin (<3 cm thick), defined by very diffuse stratification of indurated ash (Fig. 6D, 7A). LP2-A2-1 (<2 m thick), which conformably overlies LP2-A1, has a lithic-rich base (<5 cm thick; ⌀ <1 cm; 40 vol. % lithics; Fig. 6C). LP2-A2-1 is also massive and reverselygraded (⌀ ~3 mm to 14 cm), characterised by a clast-supported framework of rhyodacitic white (⌀ 14 cm; 96 vol. %) and grey (⌀ 8 cm; 2 vol. %;) pumice, and dispersed volcanic

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lithics (⌀ <4 cm; 2 vol. %). LP2-A2-2 is massive and nongraded, distinguished from LP2-

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A2-1 by a larger fraction of grey and banded pumice lapilli (⌀ 7 cm; 3 vol. %), and volcanic

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lithic components (⌀ <3 cm; 3 vol. %; Gertisser et al., 2009). LP2-A3 is similarly massive and nongraded (Fig. 6E, F), defined by a framework of rhyodacitic white pumice (⌀ <30 cm;

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85-93 vol. %). However, LP2-A3 also has andesitic-dacitic grey and banded lapilli/blocks (⌀ <35 cm), and cauliform basalt clasts (⌀ 10 cm), which decrease (grey and banded pumice 11

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to 3 vol. %) and increase (cauliform basalts 7 to 15 vol. %) in abundance vertically upwards in the stratigraphy. Near Fira, grey and banded pumice pyroclasts are characterised by

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elongate shapes, with long axis parallel to the depositional surface (Fig. 6E). Elsewhere,

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pyroclasts of this type show no preferential plane of alignment. Volcanic (⌀ < 60 cm; 7 to 13 vol. %) and basement-derived (⌀ < 15 cm; 0 to 2 vol. %) lithic assemblages also increase in

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abundance vertically. Isopach and isopleth patterns for LP2-A2 and A3 show significant

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decreases in unit thickness, juvenile and lithic clast diameters towards the south, with dispersal axis towards the east (Fig. 7B-E). Such characteristics suggest the deposits formed via fallout from a buoyant eruption column (Druitt et al., 1989; Druitt et al., 1999; Gertisser et al., 2009; Keller et al., 2014). Lithic horizons and reverse grading of juvenile pyroclasts at the base of LP2-A2-1 are typical of vent opening/widening and wind shifts or waxing column heights, respectively (Wilson et al., 1980; Walker et al., 1984; Wilson and Walker, 1987; Edgar et al., 2002; Folch, 2012; Jordan et al., 2016).

ACCEPTED MANUSCRIPT Figure 6. LP2-A - Massive pumice fall deposits. A: Pumice fall deposits within the caldera wall near Fira (top left of picture). Note the decrease in unit thickness towards the south (to the right of the picture). B: >20 m thick pumice fall deposits (LP2-A2, A3) overlying lithic lag breccias of the earlier LP1 eruption (Fira - Site 9). C, D: LP2-A1 precursory fall deposits, overlain by massive, reversely-graded, pumice lapillistone deposits (LP2-A2, A3) at (C)

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proximal vent localities (note the lithic-rich horizon at the base of LP2-A2-1; Fira - Site 9)

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and (D) distal vent localities (Cape Balos West - Site 4). E: Spatter-like grey and banded

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pumice in a clast-supported framework of white pumice pyroclasts. Note that the direction of clast elongation is parallel to the depositional surface (LP2-A3; Fira - Site 9). F: Massive

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pumice lapillistone with dispersed cauliform basalts/andesites (LP2-A3; Cape Balos West -

Figure 7.

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Site 4).

LP2 dispersal characteristics. A: Isopach maps detailing variations in unit

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thickness for A: LP2-A1, B: LP2-A2, C: LP2-A3. D, E: Isopleth maps for LP2-A3 detailing

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variations in (D) juvenile and (E) lithic clasts diameters. Thickness variations for units F:

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LP2-B, G: LP2-C, and H: LP2-D.

4.2.2. LP2-B - Diffusely-stratified to massive, poorly-sorted ignimbrite

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Diffusely-stratified to massive, poorly-sorted, pumice lapilli-tuff beds, conformably overly the fall deposits of LP2-A (Fig. 8A, B; Druitt et al., 1989; Druitt et al., 1999; Gertisser et al., 2009). Gertisser et al. (2009) previously defined two key subunits within LP2-B, termed LP2-B1 and LP2-B2. In this study, we have identified four key subunits, termed LP2B1-1, B1-2, B1-3 and B2, in reference to the nomenclature of Gertisser et al. (2009). LP2B1-1 (<50 cm thick) is diffusely-stratified, characterised by a matrix-supported framework of imbricated white (⌀ <3 cm; 15 vol. %) and grey (⌀ <2 cm; 3 vol. %) pumice, and volcanic

ACCEPTED MANUSCRIPT lithic assemblages (⌀ <2 cm; 15 vol. %). LP2-B1-2 (<30-40 cm thick) is massive and poorlysorted, with pinch and swell (dune-like) geometries (scales of several metres), consisting of white pumice (⌀ <10 cm; 40 vol. %), cauliform pyroclasts (⌀ <5 cm; 2 vol. %) and volcanic lithics (⌀ <6 cm; 5 vol. %). LP2-B1-3 (<80 cm thick) transitions vertically upwards from an initially diffusely-stratified base (with possible cross-bedding) to more massive depositional

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structures. The subunit consists of white pumice (⌀ <8 cm; 30 vol. %) and volcanic lithic

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components (⌀ <6 cm; 15 vol. %). LP2-B2 (<6 m thick) is entirely massive, charactered by

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inverse grading of white pumice (⌀ 20 cm; 35-40 vol. %) and subordinate volcanic lithic assemblages (⌀ 10 cm; 5 vol. %). A 30 cm interval at the base of LP2-B2 is pumice-poor and

lithic assemblages.

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lithic-rich (⌀ 20 cm; <40 vol. % lithics; Fig. 5, 8B). LP2-B1-3 and B2 lack basement-derived The ash matrix of each subunit consists of minor plagioclase and

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orthopyroxene crystal fragments, volcanic lithic grains and cuspate glass shards (Fig. 9A). On the Akrotiri Peninsula, these subunits range in thickness up to 6.6 m, although their

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thickness varies significantly between localities (up to several metres difference; Fig. 7F).

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Between Athinios and Fira, the deposits are generally thin or absent, highlighting the influence of pre-existing topography on the dispersal of LP2-B. These deposit characteristics

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(diffusely-stratified to massive, poorly-sorted, matrix-supported, cuspate glass shards) suggest the deposition of pumice, lithics and ash from a moderate to high-particle

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concentration, valley-confined, pyroclastic density current (PDC; Sparks et al., 1973; Cas and Wright, 1987; Branney and Kokelaar, 2002), associated with a magmatic style of eruption (Heiken and Wohletz, 1991; Cioni et al., 1992; van Otterloo et al., 2013).

Figure 8. LP2-B - Diffusely-stratified to massive PDC deposits. A: Vertical transition from diffusely-stratified deposits at the base, with pinch and swell geometries, to massive, unstructured, tabular ignimbrites (Cape Balos - Site 5). B: Planar-stratified deposits of LP2-

ACCEPTED MANUSCRIPT B1-1, conformably overlying pumice fall deposits of LP2-A3, overlain by a massive, ‘dunelike’ PDC deposit of LP2-B1-2, a diffusely-stratified (predominantly planar with cryptic cross-bedding) to massive PDC deposit of LP2-B1-3, and a massive ignimbrite of LP2-B2. Note the occurrence of a lithic horizon at the base of LP2-B2 (Cape Balos - Site 5).

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Figure 9. Morphological features of ash shards from the LP2 sequence. A: Cuspate glass

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shards associated with a magmatic style of eruption (LP2-B2). B: Angular glass shards

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associated with a predominantly phreatomagmatic style of eruption (LP2-C2-2).

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4.2.3. LP2-C - Massive, fines-poor, lithic breccias and pumice-lithic lapilli-tuffs A package of massive lithic breccias (LP2-C1-1, C2-1, C3-1), and thick (<18.5 m),

of LP2-B (Fig. 10A, B).

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ash-rich, pumice-lithic lapilli-tuffs (LP2-C1-2, C2-2, C3-2) directly overly the PDC deposits LP2-C1-1 is generally tabular and unconstrained by

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paleotopography, although pinch and swell geometries on scales of up to five metres are

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common where the breccias have infilled erosional gullies (Fig. 10C; <8 m thick; Gertisser et al., 2009). The subunit is characterised by a clast-supported framework of white pumice

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lapilli (⌀ <6 cm; 20-25 vol. %), and a variety of volcanic (⌀ <1-2 m; <70 vol. %) and basement-derived (⌀ <5 cm; <2 vol. %) lithic assemblages (Fig. 10D). LP2-C1-1 grades

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sharply upwards into a pumice-lithic lapilli-tuff unit (LP2-C1-2), characterised by a matrixsupported framework of white pumice (⌀ <30 cm; 30 vol. %;) and volcanic lithic assemblages (⌀ 60 cm; 20 vol. %). LP2-C2-1 and C3-1 are similar to LP2-C1-1, except they are laterally discontinuous, less than one metre thick, matrix-supported and lack basementderived lithic assemblages (Fig. 10B). LP2-C2-2 and C3-2 have a similar componentry to LP2-C1-2. The ash matrix of LP2-C1-2, C2-2 and C3-2 consists of crystal fragments, volcanic lithic grains and blocky to cuspate glass shards (Fig. 9B). These units repeatedly

ACCEPTED MANUSCRIPT increase and decrease in thickness over several tens to hundreds of metres within the caldera wall sequence, parallel to the Akrotiri Peninsula (Fig. 7G). The massive, pumice-lithic lapilli-tuff deposits are interpreted as ‘Layer 2b’ ignimbrites that were deposited rapidly from a PDC (Sparks et al., 1973; Cas and Wright, 1987; Branney and Kokelaar, 2002). The abundance of lithic components and fine ash, with blocky to cuspate textures, in addition to

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the absence of accretionary lapilli, suggest these flow types were the product of a relatively

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‘dry’ phreatomagmatic style of eruption (Heiken and Wohletz, 1991; Cioni et al., 1992; van

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Otterloo et al., 2013; Jordan et al., 2014). The lithological properties of the LP2-C1-1 massive lithic breccia, and its association with the overlying pumice-lithic lapilli-tuff, are

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characteristic of Layer 1 ground breccias, which form through intense fluidisation at the head of a PDC (Walker et al., 1981; Druitt and Sparks, 1982; Wilson and Walker, 1982; Walker,

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1983; Walker, 1985; Allen and Cas, 1998) or at turbulent flow boundary layers (Fisher, 1990; Buesch, 1992). Subsequent lithic breccias (LP2-C2-1, C3-1) within the ignimbrite, which are

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matrix-supported, are interpreted to reflect pulses within the PDC and discreet, high intensity

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explosive events.

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Figure 10. LP2-C - Massive, fines-poor, lithic (ground) breccia overlain by a massive, poorly-sorted, pumice-lithic tuff. A: Massive phreatomagmatic PDC deposits of LP2-C

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overlying the LP2-A and B (Cape Balos - Site 5). B: Massive ignimbrite with discontinuous lithic horizons (lh; LP2-C; Oia 2 - Site 11). C: Ground breccia, with pinch and swell morphologies (LP2-C1-1), grading sharply upwards into a massive phreatomagmatic ignimbrite (LP2-C1-2; Cape Balos - Site 5). D: Matrix to clast-supported ground breccia (LP2-C1-1; Cape Balos West - Site 6).

4.2.4. LP2-D - Massive to diffusely-stratified, fines-poor, poorly-sorted, lithic breccia

ACCEPTED MANUSCRIPT The LP2 sequence is capped by a thick package of rhyodacitic, lithic-rich breccias (Druitt et al., 1989; Druitt et al., 1999; Gertisser et al., 2009) that transition variably upwards from massive and matrix-supported (LP2-D1, D3) to diffusely-stratified and clast-supported (LP2-D2, D4; Fig. 11A, B). Subunits of this facies consist of varying proportions of white pumice (⌀ <5 cm; 3-10 vol. %), volcanic (⌀ <60 cm; 70 vol. %;) and basement-derived lithic

Although relatively widespread and tabular in geometry, LP2-D varies in

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(Fig 11A).

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assemblages (⌀ <20 cm; 2-5 %). Scour type structures are common at the base of LP2-D1

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thickness by up to several metres between closely spaced field sites (4.8 m at Site 5 to 8.4 m at Site 6; Fig. 7H). The abundance of lithics in the preserved deposit sequence, and the lack (i.e., lithic breccias don’t transition upwards into

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of identifiable facies associations

ignimbrite units), suggests that these breccias were deposited in deflation zones close to the

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vent whereby PDCs were unable to support the relatively dense lithic blocks that were incorporated into the collapsing eruption column (Wright and Walker, 1977; Druitt and

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Sparks, 1982; Druitt, 1985; Walker, 1985; Allen and Cas, 1998; Branney and Kokelaar, 2002;

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Pittari et al., 2008). Scour type structures at the base of LP2-D suggest the lithic-rich PDCs were highly erosive. These deposit characteristics are typical of lithic lag breccias and are

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expected to grade laterally (away from the vent) into lithic-poor ignimbrites (e.g., Minoan

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sequence; Druitt et al., 1999; Druitt, 2014).

Figure 11. LP2-D - Massive to diffusely-stratified, poorly-sorted, lithic lag breccias. A: Alternating horizons of massive and diffusively-stratified lithic breccias, with scour type structures at the base (LP2-D; Cape Balos West - Site 6). B: Massive, matrix-supported lithic breccia. (LP2-D1; Cape Balos West - Site 6).

4.3. Pyroclast properties

ACCEPTED MANUSCRIPT 4.3.1. Density and Vesicularity The median density of rhyodacitic pumice pyroclasts decrease (LP2-A to LP2-B) and then increase (LP2-B to LP2-D) within the stratigraphy of the LP2 sequence (Fig. 12A). Pumice from LP2-A2-1 have density and vesicularity ranges of 420-1230 kg m-3 and 52-83 vol. %, respectively, with median densities of 620 kg m-3. Pumice clasts from LP2-A2-2 and

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LP2-A3 have much lower densities of 350-1030 kg m-3 (VeT – total vesicularity 60-86 vol. %)

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and 360-950 kg m-3 (VeT 63-86 vol. %), respectively, with medians of 550 kg m-3 and 530 kg m-3. The median density of pumice in LP2-B2 is 460 kg m-3 (VeT 82 vol. %), ranging from

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280-640 kg m-3 (VeT 75-89 vol. %). Pumice clasts in LP2-C2-2 and LP2-D2 have similar

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minimum densities (270 kg m-3 and 260 kg m-3) to LP2-B2, but have significantly higher median (480 kg m-3; VeT 81 vol. %; LP2-C2-2, D2) and maximum densities of 710 kg m-3

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(VeT 72 vol. %; LP2-C2-2) and 1000 kg m-3 (VeT 60 vol. %; LP2-D2; Fig. 12A). Isolated vesicularity (VeI) ranges consistently between 1-13 vol. %, irrespective of stratigraphic

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position. However, a very-weak relationship between VeI and VeT suggests that increasing

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vesicle connectivity results as a function of increasing total vesicularity (Figure 12B). Based on these analyses, low, modal and high density pumice pyroclasts were sectioned for 2D

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image analysis and the characterisation of vesicle textures.

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Figure 12. Density and vesicularity characteristics for rhyodacitic pumice pyroclasts from the fall deposits (LP2-A2-1, A2-2, A3), magmatic ignimbrite (LP2-B2), phreatomagmatic tuff (LP2-C2-2) and lithic lag breccia (LP2-D). A: Density histograms. Bin size 100 kg m-3. Stars denote density classes from which samples were sectioned for 2D image analysis. B: Bivariate plots of connected versus total vesicularity. Oblique dashed lines represent isolated vesicularity (5 vol. % intervals).

ACCEPTED MANUSCRIPT 4.3.2. Vesicle textures and size distributions Vesicles in analysed pumice from the LP2 sequence define symmetrical to coarseskewed unimodal to bimodal size distributions, with modes of 30 to 75 μm (Fig. 13A-C). Pumices from LP2-A (A2-1, A2-2, A3) have largely homogenous microvesicular textures, characterised by small, round, closely-spaced vesicles (diameters <60 μm) with thin glass

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walls (1-10 μm), irrespective of pyroclast density (Fig. 14A, B). Vesicle walls 1-5 μm thick

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are common between vesicles of similar size and shape, reflecting the simultaneous growth of

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closely-spaced bubbles (Klug et al., 2002). Vesicles >60 μm often have more elongate or irregular shapes, resulting from stretching or rupturing of the bubble wall during coalescence

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(Fig. 14A; e.g., Klug et al., 2002). Between 52-74 % of vesicles in each sample are <60 μm, while 13 to 32 % are >100 μm. In LP2-A1 and LP2-A3 the fraction of vesicles >100 μm

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decreases as a function of increasing pumice density (32 to 13% in LP2-A2-1; 24 to 16 % in LP2-A3); no obvious relationship exists for pumices from LP2-A2-2. This suggests that

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pyroclast density is controlled, at least in part, by the abundance of large vesicles (Fig. 14B).

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Pumice from LP2-B2, C2-2 and D2 have a smaller fraction of vesicles <60 μm, between 48 and 69 %. The fraction of vesicles >100 μm also decreases with increasing pumice density

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(38 to 17 % in LP2-B2; 33 to 21 % in LP2-C2-2; 40 to 27 % in LP2-D2), similar to clasts from LP2-A (Fig. 14A). However, unlike LP2-A, clasts from LP2-B2, C2-2 and D2 exhibit

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distinct microvesicular textural domains (Fig. 14A - LP2-B2, C2-2). These domains are typically bound by larger, more expanded vesicles that range in shape from round to stretched and highly irregular. Stretched vesicles in the samples reflect localised and preferential shearing (Klug et al., 2002; Houghton et al., 2010; Shea et al., 2012), while the irregular vesicle population reflects bubble coalescence (Fig. 14A; Klug et al., 2002). Vesicles >100 μm in diameter show evidence for either coalescence, resulting in irregular vesicle shapes, or significant bubble expansion. With respect to the latter, the surrounding vesicle population is

ACCEPTED MANUSCRIPT often variably deformed, appearing to wrap around or trace the boundary of the expanded vesicle (Fig. 14A - LP2-B2). Many clasts from the PDCs also have coarsely vesicular interiors and small vesicle rinds. Such textures likely reflect post-fragmentation bubble growth in the eruption column and encompassing PDCs (e.g., Houghton et al., 2010). Such

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textures are absent in clasts from LP2-A.

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Figure 13. Vesicle size distributions and number densities for low, modal and high density

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clasts from the fall deposits (LP2-A2-1, A2-2, A3), magmatic ignimbrite (LP2-B2), phreatomagmatic tuff (LP2-C2-2) and lithic lag breccia (LP2-D). A: Low density. B: Modal

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density. C: High density.

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Figure 14. Vesicle characteristics for pyroclasts from the fall deposits (LP2-A2-1, A2-2, A3), magmatic ignimbrite (LP2-B2), phreatomagmatic tuff (LP2-C2-2) and lithic lag breccia

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(LP2-D). A: Characteristic vesicle textures for modal density clasts from each unit captured

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at low (x25), medium (x100) and high (x500) magnifications. B: Vesicle characteristics for low, modal and high density pumice clasts from each major stratigraphic division of the LP2

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sequence (x 100 magnification). Note the general increase in vesicle wall thickness and the decrease in the abundance of large vesicles with increasing pumice density or decreasing

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vesicularity. R - Ruptured vesicle wall. P - Planar glass wall. D - Donut-like structures. mv - Microvesicular textures.

Post frag expan. – Vesicle (bubbles) that experienced post

fragmentation expansion. Pl. – Plagioclase.

4.3.3. Vesicle number density Vesicle number density (corrected for vesicularity; NmV) in pumice ranges over an order of magnitude from 1.1 x 109 to 5.1 x 1010 cm-3. The median and maximum NmV

ACCEPTED MANUSCRIPT increases vertically upwards within the basal fallout sequence with no overlap, from 4.0 x 109 - 5.1 x 109 cm-3 in LP2-A2-1, to 5.6 x 109 - 6.2 x 109 cm-3 in LP2-A2-2, and 6.2 x 109 - 1.0 x 1010 cm-3 in LP2-A3. However, the minimum NmV for clasts from each subunit is smaller, at 1.7 x 109 cm-3 for LP2-A2-1, 3.0 x 109 cm-3 for LP2-A2-2, and 1.2 x 109 cm-3 for LP2-A3 (Fig. 15). NmV in pumice from the PDC phases are significantly more variable. Pumice from

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LP2-B2 have similar minimum and median NmV of 2.3 x 109 and 2.7 x 109 cm-3, respectively,

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with maximum NmV over an order of magnitude larger, at 5.1 x 1010 cm-3. The NmV in

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pumice from LP2-C2-2 and LP2-D2 are generally similar (1 x 109 to 3 x 109 cm-3), with the exception of the modal density clast in D2 which has clusters of small round vesicles and

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associated number densities of 1.3 x 1010 cm-3. Excluding two samples from LP2-B2 and D2, which have anomalously high NmV, pumice from the PDC phases typically exhibit lower NmV

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relative to LP2-A (Table 2; Fig. 15). This is attributed to the occurrence of larger vesicles in

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clasts from LP2-B, C and D, related to post-fragmentation expansion.

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Table 2. Vesicularity data for low, modal and high density pyroclasts from selective units of the LP2 sequence. VeT - Total vesicularity. NmV - vesicularity corrected vesicle number

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10 μm.

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density for vesicles >1 μm. NVF - corrected vesicle number density for vesicles between 1 and

Figure 15. Variations in NmV between units of the LP2 sequence, compared with vesicle number density data for other well-known silicic eruptions (Adams et al., 2006; Carey et al., 2010; Houghton et al., 2010; Klug et al., 2002).

NmV for the LP2 clasts have been

recalculated with a minimum vesicle cut-off of 4 μm to enable comparison with other studies. Note that the NmV in pumice from the Crater Lake eruption was previously determined based

ACCEPTED MANUSCRIPT on a minimum vesicle cut-off size of 1 μm (Klug et al., 2002). Tie line denotes median NmV in clasts from the LP2 sequence.

5. DISCUSSION 5.1. Eruption initiation and magma ascent

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The high vesicularity of pumice, in addition to the broad spectrum of vesicle sizes and

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the observed textural characteristics of those clasts, attest to a prolonged history of magma

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degassing, bubble nucleation and growth. Symmetrical to coarse-skewed, unimodal vesicle size distributions in pumice from the fall deposits indicate a near continuous history of

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magma degassing (e.g., Klug et al., 2002; Alfano et al., 2012), with the largest vesicle population representing the earliest recorded stage of bubble nucleation and growth.

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Experimental results indicate that at confining pressures of 100 MPa and magma temperatures of 856°C (as estimated for the rhyodacitic LP2 magma), magma saturation was

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reached at 3.4 wt. % H2O (Cadoux et al., 2014). At higher confining pressures of 200 MPa,

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5.6-8 wt. % water is necessary to achieve saturation (Cadoux et al., 2014). Measured preeruptive water contents of 5-6 wt. % and calculated storage pressures of 200 MPa suggest

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that the magma was close to or at saturation at reservoir depths (Cadoux et al., 2014), and possibly experienced a period of equilibrium degassing prior to dyke propagation and magma

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ascent. Because H2O has a low specific volume of <3 cm3/g in 800-900°C magma at 200 MPa (McBirney, 1963), dyke propagation and magma ascent, driven purely by volatile exsolution, is unlikely. Instead, the increase in magmatic pressure, driven by the intrusion of compositionally diverse (higher temperature) magmas into the main rhyodacitic reservoir (evidenced by the occurrence of grey and banded pumice, and cauliform basaltic clasts in the stratigraphy of LP2-A), is considered to be the primary trigger for dyke propagation and magma ascent.

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5.2. Magma decompression rates At the critical point (22 MPa; crustal depths <1 km depending on density), exsolved water transitions from a supercritical fluid to a gas (bubbles). As a result, the rate of specific volume growth of H2O accelerates, leading to increased rates of decompression and magma

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ascent (20 cm3/g at 22 MPa to 100 cm3/g at 5 MPa; McBirney, 1963).

According to

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nucleation theory, the rate of decompression can be projected if the properties of magmas are

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known (Toramaru, 2006; Shea et al., 2011). Magma decompression rates were determined by combining the methods of Toramaru (2006) and Shea et al. (2011). Vesicles between 1

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and 10 μm (NVF) were considered to reflect the final stage of magma degassing and bubble nucleation during late-stage magma ascent (Shea et al., 2011). Initial water contents (5.3 wt.

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%; Cadoux et al., 2014) were corrected for magma degassing based on the porosity of vesicles >10 μm as per the equilibrium models of Gardner et al. (1999) and Shea et al.

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(2011). Magma temperatures of 856oC (Cadoux et al., 2014) and an effective surface tension

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of 0.075 N m-1, characteristic of homogeneous nucleation in dacitic to rhyolitic melts (Gardner et al., 2013), were considered in all subsequent calculations (see supplementary

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material for further details). The maximum NVF in pumice from the fall units were used in the calculation of magma decompression rate. Decompression rates were not determined for

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the PDC phases due to evidence of post fragmentation bubble modification in the analysed clasts. Based on maximum NVF of 4.6 x 109 cm-3, 5.1 x 109 cm-3 and 9.2 x 109 cm-3 for LP2A2-1, A2-2 and A3, respectively, magma decompression rate is calculated to have increased from a maximum of 18 MPa s-1 to 29 MPa s-1, over the course of the initial Plinian phase.

5.3. Eruption characteristics 5.3.1. Mass discharge rates and plume height

ACCEPTED MANUSCRIPT Mass discharge rates (MDR) are conventionally determined by measuring the volume of magma ejected over a specified period of time. In the absence of well-defined time constraints, MDR can be inferred based on eruption column height. Assessing the heights of volcanic plumes typically require well-exposed fall deposits and constraints on their dispersal (e.g., Carey and Sparks, 1986). Because most of the outcrop is within the caldera wall,

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relatively few isopleths could be extrapolated with certainty. Nonetheless, we have attempted

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a calculation, and estimate minimum plume heights of ~36 km for LP2-A3, based on crosswind distances of 10.7 km for a 6.4 cm isopleth (Fig. 7E; Carey and Sparks, 1986).

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These plume heights are consistent with interpreted MDR of 108 kgs-1 (Wilson et al., 1980;

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Sparks, 1986; Mastin et al., 2009; Woodhouse et al., 2013), and are the same as proposed

1990; Bonadonna and Costa, 2013).

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5.3.2. Tephra volume

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plume heights for the Minoan eruption (36 km; Sigurdsson et al., 1990; Sparks and Wilson,

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The minimum volume for the LP2-A2 and A3 fall deposits was determined as per the method of Legros (2000), which calculates tephra volume based on the thickness and area of

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a selected isopach (V = 3.69 x thickness x area). Using this method, minimum dense rock volumes of 0.08 km3 and 0.27 km3 were determined for LP2-A2 and A3, respectively (Fig.

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16). In contrast, Keller et al. (2014) estimates a combined minimum dense rock tephra volume of 0.60 km3 (LP2-A2 and A3), based on the method of Fierstein and Nathenson (1992). The volumes for the PDC phases could not be determined due to burial by the deposits of successive eruptions, deposition in a marine environment and multiple caldera collapse events.

ACCEPTED MANUSCRIPT Figure 16. Isopach maps for the Plinian fall deposits (LP2-A2, A3), incorporating thickness data from Keller et al. (2014) on the island of Anafi. The minimum volume of each fall deposit was determined based on the deposit thickness and inferred area for each defined isopach, where V = 3.69 x thickness x area (Legros 2000).

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5.4. Transitions between eruption phases

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5.4.1. Precursor (LP2-A1) to Plinian (LP2-A2, A3)

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The characteristics of the precursor unit (LP2-A1) at the base of the deposit sequence attest to an initial short-lived plume from an opening phase of the eruption. Since the contact between

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it and the overlying main fallout unit (LP2-A2) is conformable at every observable location, and no paleosoil exists, it’s possible that only a few hours to weeks separated the two events.

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A repose period of more than a few weeks is unlikely given the excellent preservation of LP2-A1 (c.f., a 2-day hiatus during the 1600 eruption of Huaynaputina resulted in localised

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erosion of the earlier Phase 1 deposits, prior to the initiation of Phase 2; Adams et al. 2001).

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The dispersal and grainsize characteristics of LP2-A1 are comparable to Unit P0 of the Minoan sequence, interpreted as fallout from a 7-10 km plume (MDRs of ~106 kg s-1) during

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an initial phreatomagmatic event (Heiken and McCoy, 1990; Cioni et al., 2000; Druitt, 2014),

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suggesting similar plume heights and eruption conditions.

5.4.2. Plinian eruption column (LP2-A2, A3) to pyroclastic flows (LP2-B) The dispersal characteristics of the Plinian fall units (LP2-A2, A3) indicate a subaerial point source vent located NE of present-day Nea Kameni (Druitt et al., 1989; 1999; Keller et al., 2014). The concentration of lithics at the base of LP2-A2-1, in addition to evidence of reverse grading of pumice lapilli, attest to initial vent opening and clearing, followed by the progressive increase in eruption column height to ~36 km, comparable to phase P1b of the

ACCEPTED MANUSCRIPT Minoan eruption (Bonadonna and Costa, 2013).

The progressive increase in lithic

components vertically within the stratigraphy attest to continuous erosion of the conduit and vent over the course of LP2-A. The resulting change in conduit and vent geometry led to an increase in magma decompression rate (18 to 29 MPa s-1), a deepening of the fragmentation surface and the incorporation of basement-derived lithic components into the eruption

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column. The former also occurred during phases EU1 and EU2 of the 79 AD Vesuvius

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eruption (Shea et al., 2011), while a deepening of the fragmentation surface, as evidenced by the occurrence of plutonic lithics towards the top of Unit P1b, occurred during the initial

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Plinian phase of the Minoan eruption (Druitt, 2014). Inferred MDRs of 108 kg s-1 suggests

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that LP2-A was the product of at least four hours of near continuous fallout from a buoyant eruption column (based on a minimum DRE volume of 0.6 km3; Keller et al., 2014).

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Minimum volume estimates for P1 of the Minoan eruption of 0.24 km3 DRE (Pyle, 1990b) are four to eight times smaller than maximum dense rock estimates of 1-1.2 km3 (Pyle, 1990b)

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and 2 km3 (Sigurdsson et al., 1990, 2006). This discrepancy suggests significantly larger

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maximum dense rock volumes for LP2-A of 2.4-5 km3, which translates to ~16-32 hrs of pumice fallout prior to eruption column collapse. Because the mean density of juvenile

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pyroclasts decreased between LP2-A and B, and the influence of water during the early eruptive phases was negligible, we propose that continued vent widening and the inability to

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entrain air into the eruption column resulted in eruption column collapse (LP2-B; e.g., Andrews and Gardner, 2009).

5.4.3. Pyroclastic density currents (LP2-B) and the incorporation of water into the vent system (LP2-C) Eruption column collapse and the development of PDCs resulted in the emplacement of diffusely-stratified to massive ignimbrites (LP2-B). The absence of basement-derived

ACCEPTED MANUSCRIPT lithics within the ignimbrite indicates a shallowing of the fragmentation surface at the onset of this eruption phase, while the vertical change from diffusely-stratified to massive units reflect an increase in MDR and deposit aggradation rate or the progressive transition from an unsteady to steady density current. The former also occurred during the transition to the initial PDC phase of the Minoan eruption (P1 to P2; Druitt, 2014). Enlargement of the

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existing vent, or the development of new vents during incipient caldera collapse, facilitated a

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deepening of the fragmentation surface (as inferred by the occurrence of basement-derived lithic assemblages in LP2-C1-1) and the ingress of water into the magmatic system. This

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resulted in the development of lithic-rich PDCs, associated with a “dry” phreatomagmatic

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style of eruption. Multiple subunits of LP2-C attest to several discrete vent excavation events. The absence of basement-derived lithic assemblages, within the main body of these

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depositional subunits (excluding the basal unit LP2-C1-1), again suggests shallowing of the

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fragmentation surface over the course of this phase.

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5.4.3.1. The source of water driving phreatomagmatism We suggest that the main vent and any subsequent vents that developed over the

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course of the eruption, formed within a pre-existing flooded depression of Santorini (most likely a caldera associated with the earlier 196 ka Cape Therma 3 and 184 ka Lower Pumice

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1 eruptions; Simmons et al., 2016; Simmons et al., 2017). The ingress of water from the Aegean Sea (as is the case for present-day Santorini), or rainfall, may have flooded the early caldera. Although no direct evidence exists for either process, sapropel layers (layer S6 175-165 ka) in deep sea cores from the Mediterranean Sea (cores DED87-08 and KET80-04) and Tyrrhenian Sea (core MD84-641) show high organic carbon content and relatively low δ18O of G. bulloides and G. ruber (δ18O 0-0.2; foraminifer), suggesting high rainfall with low surface temperatures around the time of the LP2 eruption (12°C; Kallel et al., 2000). While

ACCEPTED MANUSCRIPT not discounting the possibility that the caldera was ‘open’ to the Aegean Sea, this data lend credence to the hypothesis that rain water, and rain water runoff, may have equally infilled a pre-existing caldera.

5.4.4. Transition to caldera collapse (LP2-D)

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The occurrence of lithic-rich lag breccias at the top of the deposit sequence, and a

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northward dipping unconformity (which truncates and terminates at the top of the lag

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breccias) beneath Fira (Fig. 17), suggests that caldera collapse accompanied and followed LP2-D. In particular, magma evacuation over the course of the eruption and subsequent

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under pressurisation of the magma reservoir led to the break-up of the reservoir roof block, the widening of fracture as vents, the discharge of lithic-rich PDCs and finally roof-block

Simmons et al., 2017).

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subsidence (Druitt and Sparks, 1984; Branney, 1995; Lipman, 1997; Cole et al., 2005; Paleomagnetic measurements indicate the lithic breccias were

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deposited at relatively low temperatures of approximately 200 °C, although several clasts

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from the same depositional unit indicate emplacement temperatures ranging from 100-600 °C (Bardot, 2000; Bardot et al., 1996). This variability has been attributed to the fragmentation

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and incorporation of crustal fragments from a range of crustal levels (shallow and deep), which have temperatures that fluctuate in accordance with the geothermal gradient, and the

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location and depth of the magma reservoir (Bardot, 2000). The occurrence of both volcanic and basement-derived lithic assemblages in LP2-D, support this conclusion, indicating deep fragmentation depths.

Such interpretations are consistent with field-based studies (e.g.,

Edgar et al., 2002; Simmons et al., 2017) and numerical models (e.g., Martı et al., 2000), which propose a deepening of the fragmentation surface as a consequence of decompression and caldera collapse.

ACCEPTED MANUSCRIPT Figure 17. Northward dipping unconformity truncating the LP2 sequence, reflecting latestage caldera collapse or post eruption slumping. Bedding of the Middle tuff series (ap4) is parallel to the unconformity. Deposits of the Middle tuff series also pinch towards the south, over the unconformity.

aa - Andesites of Cape Alai, ap1 - Cape Therma 1 Tuff, ra -

Rhyodacites of Cape Alonaki, rp3 - Lower Pumice Tuffs, ap4 - Middle tuffs, as2 - Andesites

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and basalts of Cape Skaros, rt - Rhyodacites of Therasia, rpb7 - Minoan tuff (nomenclature

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from Druitt et al., 1999).

5.5. Implications

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Changes in the dynamics and dispersal processes of Plinian eruptions can result as a consequence of a variety of processes, including changes in magma properties (e.g.,

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composition, density, volatile content, viscosity; Wilson et al., 1980; Carey et al., 1990; Civetta et al., 1997; Balcone-Boissard et al., 2011; Shea et al., 2011), variations in magma

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discharge rates (Wilson et al., 1980; Pensa et al., 2015), the presence or absence of external

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water (Morrissey et al., 2000), and the geometry of the conduit and vent (Wilson et al., 1980; Carey and Sigurdsson, 1989; Woods and Bower, 1995; Koyaguchi et al., 2010). Interestingly,

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many sequences transition repeatedly from plume to flow events. For example, the 4.6 ka Fogo A sequence preserves evidence for simultaneous plume and ignimbrite-forming events

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resulting from instabilities in the eruption column and rapid variations in MDR, possibly associated with changing magma chemistries (Pensa et al., 2015). Alternating and repeated transitions from plume to flow during the 79 AD Vesuvius eruption have been interpreted as resulting from an increase in juvenile clast density (Shea et al., 2011), while interbedded phreatomagmatic surge deposits in the Minoan fallout deposit reflect the episodic influx of water into the magmatic system (Heiken and McCoy Jr, 1984). Such rapid and repeated transitions in dispersal processes are not typically associated with changes in vent

ACCEPTED MANUSCRIPT morphology. These examples are in contrast to the LP2 eruption, whereby vent widening facilitated eruption column collapse. Thus, in the presence of an unchanging single-vent system, fluctuations in dispersal processes may occur over relatively short temporal scales. However, given that all vents evolve over the course of an explosive eruption, we suggest that once a critical vent diameter/area is achieved, eruption column collapse and the

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development of PDCs is inevitable.



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6. CONCLUSIONS

The LP2 eruption was triggered by the intrusion of primitive magmas into the more

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evolved rhyodacitic magmatic system, which facilitated dyke propagation and the initial discharge of magma from a point-source vent. The eruption commenced with the development of a short-lived precursor plume that

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

deposited small pumice lapilli at near vent localities (LP1-A1). Vent widening over the course of the initial Plinian phase (LP2-A2, A3) resulted in a

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

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deepening of the fragmentation surface and increased rates of magma decompression (18 MPa s-1 to 29 MPa s-1), which was followed by eruption column collapse and the

A major vent excavation event, possibly associated with incipient caldera collapse,

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

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development of pyroclastic density currents (LP2-B).

resulted in a rapid increase in the fragmentation depth, followed by shallowing and the ingress of water into the magmatic system (LP2-C). 

The development of fracture vents, associated with under-pressurisation of the magma reservoir, resulted in a deepening of the fragmentation surface, the discharge of lithicrich pyroclastic density currents and caldera collapse (LP2-D).



Although magma properties (composition, volatile content, temperature, viscosity etc.) are first order factors influencing the intensity of eruptions, the geometry and structure

ACCEPTED MANUSCRIPT of the vent system plays a very significant role in the course and temporal evolution of highly explosive, caldera-forming Plinian eruptions.

ACKNOWLEDGMENTS This research was conducted as part of a PhD research dissertation supported by

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discretionary research funds of Ray Cas and a Faculty of Science Dean’s Postgraduate

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Research Scholarship. The paper has benefited from constructive reviews by S. Allen and

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one anonymous reviewer whom we thank for their suggestions. Thank you also to Jörg Keller for his correspondence regarding the V-1 (LP2) tephra layer in deep sea cores

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throughout the Aegean. The authors acknowledge use of facilities within the Monash Centre for Electron Microscopy (MCEM), which are supported under the Australian Research

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Council's Centres of Excellence funding scheme (COE for Design in Light Metals). This is

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Laboratory of Excellence ClerVolc contribution number xx.

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ACCEPTED MANUSCRIPT Newman, A.V., Stiros, S., Feng, L., Psimoulis, P., Moschas, F., Saltogianni, V., Jiang, Y., Papazachos, C., Panagiotopoulos, D., Karagianni, E. and Vamvakaris, D., 2012. Recent geodetic unrest at Santorini Caldera, Greece. Geophysical Research Letters, 39(6). Nicholls, I.A., 1971. Petrology of santorini volcano, cyclades, Greece. Journal of Petrology,

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Nomikou, P., Carey, S., Papanikolaou, D., Croff Bell, K., Sakellariou, D., Alexandri, M. and

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Bejelou, K., 2012. Submarine volcanoes of the Kolumbo volcanic zone NE of Santorini Caldera, Greece. Global and Planetary Change(0).

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Nomikou, P., Papanikolaou, D., Alexandri, M., Sakellariou, D. and Rousakis, G., 2013. Submarine volcanoes along the aegean volcanic arc. Tectonophysics, 597-598: 123-

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Papadimitriou, P., Kapetanidis, V., Karakonstantis, A., Kaviris, G., Voulgaris, N. and

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Makropoulos, K., 2015. The Santorini Volcanic Complex: A detailed multi-parameter

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seismological approach with emphasis on the 2011-2012 unrest period. Journal of Geodynamics, 85: 32-57.

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Papanikolaou, D., 1987. Tectonic Evolution of the Cycladic Blueschist Belt (Aegean Sea, Greece). In: H. Helgeson (Editor), Chemical Transport in Metasomatic Processes.

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ACCEPTED MANUSCRIPT Pensa, A., Cas, R., Giordano, G., Porreca, M. and Wallenstein, N., 2015. Transition from steady to unsteady Plinian eruption column: the VEI 5, 4.6 ka Fogo A Plinian eruption, São Miguel, Azores. Journal of Volcanology and Geothermal Research, 305: 1-18. Perissoratis, C., 1995. The Santorini volcanic complex and its relation to the stratigraphy and structure of the Aegean arc, Greece. Marine Geology, 128(1–2): 37-58.

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Pfeiffer, T., 2001. Vent development during the Minoan eruption (1640 BC) of Santorini,

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Doumas (Editor), Thera and the Aegean World II. Thera and the Agean World, London, pp. 15-30.

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Pittari, A., Cas, R.A.F., Wolff, J.A., Nichols, H.J., Larson, P.B. and Martí, J., 2008. Chapter 3 The Use of Lithic Clast Distributions in Pyroclastic Deposits to Understand Pre- and

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Polacci, M., 2005. Constraining the dynamics of volcanic eruptions by characterization of pumice textures. Annals of Geophysics, 48(4-5): 731-738.

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Pyle, D., 1990a. Geological and radiochemical studies on young volcanoes: Santorini, Greece and Oldoinyo Lengai, Tanzania, University of Cambridge. Pyle, D., 1990b. New estimates for the volume of the Minoan eruption. In: Hardy, D.A., Keller, J., Galanopoulos, V.P., Flemming, N.C., Druitt, T.H. (Eds.), Thera and the Aegean World III, vol. 2. The Thera Foundation, London, pp. 113–121. Pyle, D.M. and Elliott, J.R., 2006. Quantitative morphology, recent evolution, and future activity of the Kameni Islands volcano, Santorini, Greece. Geosphere, 2(5): 253-268.

ACCEPTED MANUSCRIPT Sahagian, D.L. and Proussevitch, A.A., 1998. 3D particle size distributions from 2D observations: stereology for natural applications. Journal of Volcanology and Geothermal Research, 84(3–4): 173-196. Schliestedt, M., Altherr, R. and Mathews, A., 1987. Evolution of the Cyclandic crystalline complex: petrology, isotope geochemistry and geochronology. In: R. Helgeson

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Shea, T., Gurioli, L. and Houghton, B.F., 2012. Transitions between fall phases and pyroclastic density currents during the AD 79 eruption at Vesuvius: Building a

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transient conduit model from the textural and volatile record. Bulletin of Volcanology, 74(10): 2363-2381.

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Shea, T., Gurioli, L., Houghton, B.F., Cioni, R. and Cashman, K.V., 2011. Column collapse and generation of pyroclastic density currents during the A.D. 79 eruption of

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Simmons, J. M., Cas, R.A.F., Druitt, T.H. and Folkes, C.B., 2016. Complex variations during

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co-ignimbrite breccias and climactic lag breccias: The 184 ka Lower Pumice 1 eruption sequence, Santorini, Greece. Journal of Volcanology and Geothermal

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Research, 324: 200-219.

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Volcanology, 48(1): 3-15.

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Sparks, R.S.J. and Wilson, C.J.N., 1990. The Minoan Deposits: A review of their Characteristics and Interpretation. In: D.A. Hardy, J. Keller, V.P. Galanopoulos, N.C.

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ACCEPTED MANUSCRIPT Urai, J.L., Schuiling, R.D. and Jansen, J.B.H., 1990. Alpine deformation on Naxos (Greece), Geological Society Special Publication, pp. 509-522. van der Maar, P. and Jansen, J.B., 1983. The geology of the polymetamorphic complex of Ios, Cyclades, Greece and its significance for the Cycladic Massif. Geologische Rundschau, 72(1): 283-299.

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van Otterloo, J., Cas, R.A.F. and Sheard, M.J., 2013. Eruption processes and deposit

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characteristics at the monogenetic Mt. Gambier Volcanic Complex, SE Australia:

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Implications for alternating magmatic and phreatomagmatic activity. Bulletin of Volcanology, 75(8): 1-21.

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Walker, G.P.L., 1981a. Plinian eruptions and their products. Bulletin Volcanologique, 44(3): 223-240.

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Walker, G.P.L., 1981b. The Waimihia and Hatepe plinian deposits from the rhyolitic Taupo Volcanic Centre. New Zealand Journal of Geology and Geophysics, 24(3): 305-324.

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Walker, G.P.L., 1983. Ignimbrite types and ignimbrite problems. Journal of Volcanology and

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Geothermal Research, 17(1-4): 65-88. Walker, G.P.L., 1985. Origin of coarse lithic breccias near ignimbrite source vents. Journal of

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Volcanology and Geothermal Research, 25(1-2): 157-171. Walker, G.P.L., Self, S. and Froggatt, P.C., 1981. The ground layer of the taupo ignimbrite:

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ACCEPTED MANUSCRIPT Wallace, P.J., Plank, T., Edmonds, M. and Hauri, E.H., 2015. Chapter 7 - Volatiles in Magmas. In: H. Sigurdsson (Editor), The Encyclopedia of Volcanoes (Second Edition). Academic Press, Amsterdam, pp. 163-183. Wilson, C.J.N., Houghton, B.F., McWilliams, M.O., Lanphere, M.A., Weaver, S.D. and Briggs, R.M., 1995. Volcanic and structural evolution of Taupo Volcanic Zone, New

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Zealand: a review. Journal of Volcanology and Geothermal Research, 68(1-3): 1-28.

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Wilson, C.J.N. and Walker, G.P.L., 1982. Ignimbrite depositional facies: the anatomy of a

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Wilson, L. and Walker, G.P.L., 1987. Explosive volcanic eruptions - VI. Ejecta dispersal in plinian eruptions: the control of eruption conditions and atmospheric properties.

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Geophysical Journal International, 89(2): 657-679.

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Woodhouse, M., Hogg, A., Phillips, J., Sparks, R., 2013. Interaction between volcanic plumes and wind during the 2010 Eyjafjallajökull eruption, Iceland. Journal of

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Geophysical Research B: Solid Earth, 118(1), 92-109. Woods, A.W. and Bower, S.M., 1995. The decompression of volcanic jets in a crater during

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explosive volcanic eruptions. Earth and Planetary Science Letters, 131(3–4): 189-205. Wright, J.V. and Walker, G.P.L., 1977. The ignimbrite source problem: Significance of a coignimbrite lag-fall deposit. Geology, 5(12): 729-732.

Figure 1

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Figure 10

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Figure 11

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Figure 12

Figure 13

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Figure 14

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Figure 16

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Figure 17

ACCEPTED MANUSCRIPT TABLE 1.

Major Elements (Wt. %) SiO2

TiO2

Al2O3

Fe2O3

MnO

MgO

CaO

Na2O

K2O

P2O5

SO3

LOI

Tot.

W

64.72

0.45

13.92

3.32

0.12

0.91

1.74

5.59

2.94

0.08

bd

6.78

100.58

LP2-A2

W

63.33

0.48

14.16

3.45

0.12

0.82

1.97

5.58

2.85

0.1

0.02

7.97

100.86

LP2-A3

W

66.4

0.53

14.82

3.87

0.12

0.84

2.22

4.77

2.8

0.1

bd

4.26

100.73

LP2-A3

G

62.78

0.81

15.74

5.44

0.15

1.72

3.98

5.14

2.06

0.22

0.01

2.71

100.76

LP2-A3

C

63.01

0.84

16.71

4.86

0.12

1.46

4.44

4.79

2.12

0.18

0.08

2.22

100.83

LP2-B1-1

W

60.2

0.41

13.03

3.07

0.1

0.67

1.84

5.19

2.73

0.08

bd

13.51

100.82

LP2-B2

W

68.28

0.47

14.54

3.37

0.12

0.54

1.77

5.01

3.15

0.09

bd

3.63

100.97

LP2-C1

W

67.18

0.47

14.38

3.32

0.11

0.53

1.94

LP2-C2

W

67.1

0.45

14.39

3.26

0.11

0.53

1.9

LP2-C3

W

67.5

0.47

14.52

3.39

0.11

0.54

LP2-D3

W

68.1

0.49

14.6

3.33

0.11

0.62

4.84

3.12

0.09

bd

4.73

100.71

4.8

3.21

0.08

0.01

5.02

100.87

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LP2-A1

4.75

3.25

0.09

bd

4.16

100.65

1.77

4.51

3.16

0.1

bd

3.91

100.71

NU

1.87

Zr

Nb

Mo

Cd

MA

Trace Elements (ppm) Sn

Sb

Cs

Ba

La

Ce

Pr

Nd

Sm

Eu

W

274

11.7

2.40

0.14

4.29

0.35

3.17

410

26.3

54.5

6.26

23.8

5.48

1.07

LP2-A2

W

263

11.6

2.48

0.13

2.90

0.26

2.67

424

26.4

53.0

6.44

24.8

5.74

1.18

LP2-A3

W

294

12.9

2.66

0.14

3.08

0.35

3.12

466

28.6

59.3

7.02

26.9

6.19

1.28

LP2-A3

G

225

11.2

2.23

0.12

2.48

0.25

2.23

379

24.6

51.4

6.35

25.5

6.06

1.51

LP2-A3

C

223

10.7

LP2-B1

W

253

11.3

LP2-B4

W

305

12.8

LP2-C1

W

309

13.0

LP2-C2

W

316

13.4

LP2-C3

W

312

LP2-D3

W

312

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LP2-A1

2.89

0.11

2.42

0.26

2.54

437

24.6

50.3

5.98

23.3

5.44

1.37

1.92

0.10

2.46

0.27

2.55

377

26.5

50.8

5.96

22.8

5.55

1.05

2.35

0.11

2.73

0.37

3.08

468

31.3

60.2

6.98

26.8

6.48

1.23

2.78

0.16

3.78

0.30

3.36

478

29.4

60.5

7.16

27.2

6.29

1.25

0.15

3.29

0.31

3.34

511

29.8

61.1

7.15

27.2

6.26

1.25

2.77

0.15

3.17

0.31

3.25

470

29.6

60.7

7.15

27.4

6.30

1.27

13.6

2.70

0.15

3.18

0.39

3.11

456

28.4

59.1

6.93

26.6

6.15

1.15

AC

CE

2.90

13.3

ACCEPTED MANUSCRIPT Trace Elements (ppm) Li

Be

Sc

V

Cr

Co

Ni

Cu

Zn

Ga

As

Rb

Sr

Y

W

17.4

2.04

8.52

16.2

17.6

3.17

6.49

12.9

105

15.2

9.51

101

75.1

45.6

LP2-A2

W

14.1

1.88

8.74

10.7

16.4

2.43

1.46

5.60

81.2

15.4

5.66

87.9

98.0

46.8

LP2-A3

W

14.5

2.13

10.6

22.7

27.3

3.82

2.53

5.91

75.1

17.3

5.46

103

102

50.6

LP2-A3

G

11.7

1.71

15.1

62.4

24.0

7.63

4.75

9.36

98.8

17.6

4.26

75.0

162

46.7

LP2-A3

C

15.3

1.78

15.2

90.4

19.0

7.42

1.64

16.4

78.1

18.2

11.4

84.3

211

41.9

LP2-B1

W

11.8

1.81

9.25

10.6

15.5

2.79

2.68

9.31

61.9

13.5

2.63

87.8

73.9

44.3

LP2-B4

W

11.3

2.11

9.40

11.6

22.1

2.25

1.58

6.26

57.0

16.2

3.74

104

84.2

52.4

LP2-C1

W

10.7

2.16

8.87

8.45

27.8

2.08

0.61

4.52

67.0

16.6

4.71

109

83.8

52.0

LP2-C2

W

10.2

2.30

9.17

8.21

22.6

2.12

0.94

3.72

82.3

17.1

6.74

111

85.3

52.5

LP2-C3

W

11.1

2.22

9.30

9.28

29.8

2.17

0.63

3.51

68.0

17.1

6.80

106

86.9

52.4

LP2-D3

W

11.7

2.26

9.77

17.9

62.9

2.90

2.06

5.00

82.3

16.6

7.71

108

85.4

51.6

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ACCEPTED MANUSCRIPT Trace Elements (ppm) Gd

Tb

Dy

Ho

Er

Tm

Yb

Lu

Hf

Ta

Tl

Pb

Th

U

W

6.00

1.01

6.52

1.46

4.52

0.70

4.70

0.73

6.44

0.84

0.71

22.5

16.0

4.69

LP2-A2

W

6.31

1.06

6.80

1.52

4.64

0.72

4.80

0.75

6.26

0.83

0.43

17.9

15.4

4.42

LP2-A3

W

6.85

1.15

7.36

1.64

5.04

0.78

5.24

0.81

7.04

0.92

0.50

21.8

16.9

5.07

LP2-A3

G

6.71

1.12

7.09

1.56

4.69

0.71

4.66

0.72

5.53

0.79

0.39

16.5

12.7

3.83

LP2-A3

C

5.98

1.00

6.31

1.39

4.20

0.63

4.23

0.65

5.40

0.79

0.55

18.5

14.7

4.79

LP2-B1

W

5.70

1.04

6.75

1.59

4.63

0.69

4.70

0.74

6.44

0.82

0.46

17.3

14.9

4.50

LP2-B2

W

6.65

1.22

7.98

1.88

5.48

0.81

5.57

0.88

7.78

0.95

0.59

41.6

17.9

5.52

LP2-C1

W

6.95

1.18

7.53

1.69

5.23

0.81

5.46

0.85

7.47

0.95

0.63

18.8

18.1

5.46

LP2-C2

W

6.90

1.18

7.53

1.69

5.20

0.80

5.40

0.85

7.48

0.95

0.58

18.8

18.1

5.43

LP2-C3

W

6.96

1.18

7.57

1.69

5.21

0.81

5.42

0.85

7.41

0.95

0.54

19.0

17.9

5.40

LP2-D3

W

6.74

1.16

7.42

1.66

5.17

0.80

5.41

0.85

7.46

1.01

0.54

19.0

18.4

5.53

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LP2-A1

ACCEPTED MANUSCRIPT TABLE 2.

LP2A3

LP2A2-2

0.0496

High

60.7

0.0496

Low

86.2

0.0297

Modal

81.3

High

75.2

0.0496

Low

84.9

0.0496

Modal

82.5

High

77.4

Low

85.3

0.0297

0.0745

Modal

79.0

0.0297

0.0745

High

73.7

Low

84.3

Modal

77.8

High

73.7

LP2A2-1

82.5

AC

Low

0.0297

0.0297

0.0745

0.0745

0.0297

Modal

77.8

0.0297 0.0496 0.0297 0.0297

High

75.6

0.0496

0.0297

2.71 x 109 5.04 x 1010 2.28 x 109 6.20 x 109 1.03 x 1010 1.16 x 109

1.95 x 109 4.88 x 1010 1.28 x 109 4.38 x 109 9.15 x 109 6.08 x 108

5.60 x 109 6.19 x 109 3.04 x 109

3.85 x 109 5.13 x 109 2.66 x 109

1.68 x 109 3.95 x 109 5.12 x 109

1.07 x 109 3.17 x 109 4.63 x 109

0.0745

0.0297

N/A

N/A

PT

81.3

1.60 x 109 1.20 x 1010 9.63 x 108 2.15 x 109 1.59 x 109 6.20 x 108

RI

Modal

2.50 x 109 1.30 x 1010 1.35 x 109 3.18 x 109 2.81 x 109 1.07 x 109

Decom. Rate (MPa s-1)

SC

0.0496

NVF (cm-3)

NU

88.2

NmV (cm-3)

MA

LP2B2

Low

D

LP2C2-2

2nd Mode (mm)

VeT %

PT E

LP2D2

1st Mode (mm)

Density Class

CE

Unit

N/A

18.9 28.9 4.7 17.0 19.7 12.5 7.4 14.4 18.2

ACCEPTED MANUSCRIPT Highlights

Plinian eruption column to the development of magmatic and phreatomagmatic pyroclastic density currents, and lithic-rich density currents during late-stage caldera collapse.

PT

decompression rates, a

RI

deepening of the fragmentation surface and the development of pyroclastics density currents.

AC

CE

PT E

D

MA

NU

SC

evolution of caldera-forming, Plinian eruptions.