Textural, geochemical, and volatile evidence for a Strombolian-like eruption sequence at Lō`ihi Seamount, Hawai`i

Journal of Volcanology and Geothermal Research 207 (2011) 16–32

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

Textural, geochemical, and volatile evidence for a Strombolian-like eruption sequence at Lō`ihi Seamount, Hawai`i C. Ian Schipper a, b, c,⁎, James D.L. White c, Bruce F. Houghton d a

Institut des Sciences de la Terre d'Orléans, CNRS – Université d'Orléans, 1A rue de la Férollerie, 45071 Orléans, Cedex 2, France Soil and Earth Sciences, INR, Massey University, PB 11-222, Palmerston North, 4474, New Zealand Geology Department, University of Otago, PO Box 56, Leith St., Dunedin, 9016, New Zealand d Geology and Geophysics, University of Hawai`i at Mānoa, 1680 East-West Road, Honolulu, Hawai`i 98622, USA b c

a r t i c l e

i n f o

Article history: Received 5 April 2011 Accepted 1 August 2011 Available online 10 August 2011

a b s t r a c t Despite recent advances and observations, many questions remain about the range of eruption styles that are possible at submarine volcanoes. We report deposit characteristics, clast microtextures, and geochemical and volatile data for a ~ 22 m thick pyroclastic succession on the summit plateau of Lō`ihi Seamount, Hawaii, with the goal of determining eruption style directly from features preserved in the submarine deposits. Deposit stratigraphy indicates emplacement of lapilli-dominated ejecta from a nearby vent or series of vents. Lapilli contain heterogeneous vesicle (Nv = 1.6 × 10 4 to 9.5 × 10 5 cm − 3) and gradational microlite populations that matured below the vent, before fragmentation. The homogeneous, tholeiitic matrix glasses are degassed of CO2, but retain variable concentrations of dissolved H2O (0.41 to 0.82 wt.%). Olivine crystals are mostly highly magnesian (N Fo88) crystals, bearing inclusions entrapped at ~ 1300 ± 37 °C that suffered post-entrapment Fe-loss in contact with cooler (~1136 ± 50 °C) bulk melt over calculated time periods up to ~ 3 months, and a small subset of lower-Mg (b Fo85.5) phenocrysts bearing glass inclusions entrapped at ~ 1217 ± 12 °C. Inclusions are also mostly degassed of CO2, and retain variable H2O concentrations (0.36 to 0.74 wt.%). No single data type definitively proves what style of eruption formed this deposit, and each is treated in-turn with realistic discussion of their respective limitations. Textural features indicate Strombolian-like (e.g., episodic) ejection of magma, in pulses that allowed recharging magma to pond and partially cool in the shallow conduit between bursts. Evidence against H2O exsolution between source and vent suggest that the decoupled volatile slugs driving the eruption(s) must have been dominated by CO2 derived from unerupted magma deeper in Lō`ihi' s plumbing system. This study advances our catalog of analyses for submarine pyroclastic rocks, but also highlights that we are still far from a full understanding of submarine explosive volcanic processes. © 2011 Elsevier B.V. All rights reserved.

1. Introduction 1.1. Submarine explosive volcanism Until the late 1980's, submarine explosive eruptions of basalt were thought to be impossible at depths exceeding 500 m below sea level (mbsl). Hydrostatic pressure was assumed to suppress the volatile exsolution required to drive vigorous magma discharge (McBirney, 1963) at abyssal vents. This concept of a depth limit to basalt explosivity persisted for decades (e.g., Bonatti, 1970; Tazieff, 1972; Fisher, 1984; Fisher and Schmincke, 1984; Bonatti et al., 1988), and was strengthened by the observation that the ocean floor is largely covered with effusive volcanic products. ⁎ Corresponding author at: Institut des Sciences de la Terre d'Orléans, CNRS – Université d'Orléans, 1A rue de la Férollerie, 45071 Orléans, Cedex 2, France. Tel. + 33 6 66 30 38 73; fax: + 33 2 38 63 64 88. E-mail address: [email protected] (C.I. Schipper). 0377-0273/$ – see front matter © 2011 Elsevier B.V. All rights reserved. doi:10.1016/j.jvolgeores.2011.08.001

The last 30 years have seen a complete re-evaluation of this concept. With ever-widening use of submersible technology, the seafloor is now being observed and sampled in far greater detail than ever before. Volcaniclastic and/or pyroclastic deposits have now been found to depths exceeding 4000 mbsl (Batiza et al., 1984; Fouquet et al., 1998; Hekinian et al., 2000; Clague et al., 2003; Eissen et al., 2003; Davis and Clague, 2006; Clague et al., 2008; Sohn et al., 2008), and two low-flux submarine explosive eruptions have now been observed by remotely operated vehicles (Embley et al., 2006; Chadwick et al., 2008; Clague et al., 2009; Resing and Embley, 2009; Deardorff et al., 2011). These discoveries have brought to the fore uncertainties regarding volatile systematics and magma-water interactions during submarine explosivity. The submarine explosive eruptions at N.W. Rota 1 (Embley et al., 2006; Chadwick et al., 2008; Deardorff et al., 2011) and West Mata (Clague et al., 2009; Resing and Embley, 2009) have advanced our understanding of significant aspects of submarine eruptions, but they were both of very low eruptive flux, and were not in ridge or ocean

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island settings, so it is scientifically insupportable to assume that they represent the full range of explosive eruption styles possible in the deep ocean. Furthermore, observations at subaerial volcanoes indicate that magmas of similar chemistry can erupt in styles of significantly different character, in discrete events closely spaced in time and/or space or even simultaneously (e.g., Andronico et al., 2005; 2009). What remains to be developed is a catalog of the full range of eruption styles that occur at depth, given the practical constraint that most eruptions go unobserved, so that traces of their dynamics are preserved only in the particles and bedding characteristics of the submarine rock record. In this paper, we examine a ~ 22 m thick pyroclastic deposit exposed in a fault scarp ~ 1150 m below sea level (mbsl) on Lō`ihi Seamount, Hawaii. Using deposit characteristics, vesicle and crystal textures in lapilli, and the geochemistry of matrix glasses and olivinehosted glass inclusions, we explore the conditions of magma ascent, storage, degassing, and fragmentation that determined the eruptive conditions of this particular batch of magma. 1.2. Lō`ihi Seamount, Hawaii Lō`ihi Seamount lies ~35 km southeast of the island of Hawaii (Fig. 1A). Two subparallel, shallow-dipping rifts impart an elongate north–south geometry to this young volcano (Malahoff et al., 1982; Fornari et al., 1988), which abuts the submarine flanks of neighboring Mauna Loa and Kilauea volcanoes, yet is largely underlain by Pacific Ocean crust (Garcia et al., 1998; 2006). The western and eastern flanks dip more steeply than the rifts, both having been extensively oversteepened by large-scale mass wasting (Malahoff et al., 1982; Moore et al., 1989). Lō`ihi rises ~5 km to a ~12 km 2 summit plateau, inferred to host a caldera-like collapse structure, at ~ 1200 mbsl (Fig. 1B). The summit plateau hosts three collapse pits, the youngest of which formed in the most recent (1996) eruption (Lō`ihi Science Team, 1997), as well as numerous cones, the largest of which rises to just under 1 km depth. The summit plateau and rift zones generally lack sediment cover, and the former in particular is inferred to be the locus of the most recent eruptive activity at Lō`ihi. Due to its proximity to the intensively monitored island of Hawaii, Lō`ihi is remarkably well studied by submarine-volcano standards. Extensive studies of Lō`ihi's petrology, seismic structure, hydrother-

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mal activity, and volcaniclastic deposits have been carried out since large seismic swarms in the 1970's alerted the scientific community to the fact that Lō`ihi was very much an active volcano. For more information, readers are directed to the recent review by Garcia et al. (2006) and references therein. 2. Methods The deposit (Fig. 1C) was examined during dive series P4-160 to P4-164 with the Hawaiian Undersea Research Laboratory's Pisces IV submersible, in October 2006. Extensive traverses, both upsection and along the near-vertical face of the deposit, yielded direct observations and video footage to determine the stratigraphy and facies geometry. Volcaniclastic material was scooped directly from beds of the deposit with the submersible manipulator arm wielding a sediment scoop. Density/vesicularity of lapilli populations from seven samples were measured using techniques of Houghton and Wilson (1989). Representative lapilli from the low, modal, and high-density classes were selected for textural analysis, and yielded standard polished thin sections after impregnation with epoxy to preserve internal textures. Vesicle, phenocryst, and microlite textures were examined using petrographic and scanning election (SEM) microscopes. Vesicle populations were quantified following Shea et al. (2010). Nested images included 1200 dpi scans and SEM images at 25× (252 pixels/mm) and 75× (770 pixels/mm). After several tests, magnifications higher than 75× were deemed unnecessary to capture the full size range of vesicles present. Measured vesicle areas were converted to volumes using the stereological conversion procedure of Sahagian and Proussevitch (1998). Phenocryst populations were measured on 1200 dpi scans of lapilli with ImageJ software. Since phenocryst shapes are far from spherical, no stereological correction was used, and their abundances are reported in terms of area. Microlites were examined qualitatively with petrographic microscopes and SEM. Whole rock major element geochemistry was determined by X-ray fluorescence spectroscopy (XRF), on a Phillips PW2400 sequential XRF at the University of Otago. The major element chemistry of matrix glasses, glass inclusions, and olivine crystals was determined by electron microprobe (EMP)

Fig. 1. Location and bathymetry. A: The Hawaiian island chain showing position of Lō`ihi with respect to the big island of Hawaii. B: Oblique view showing summit plateau and deposit location (created by J.R. Smith) by merged deep-tow Reson Seabat 8101 and shipboard SIMRAD EM-12 multibeam data. C: The fault scarp-exposed deposit examined in this study (marked by white star), showing location relative to the previously described Northern Cone (NC) deposit (Schipper et al., 2010a,b), and a previously studied, non-primary volcaniclastic sequence (Clague et al., 2003; Clague, 2009; Schipper and White, 2009).

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analysis of on a JEOL JXA-8600 Superprobe, at the University of Otago. Quantitative energy-dispersive (EDS) analyses used 15 keV accelerating voltage, 1 nA beam current, and 200 s counting times, and were calibrated using Smithsonian basaltic glass standards VG-2, VGA-99, Indian Ocean glass, and San Carlos olivine standard (Jarosewich et al., 1980). H2O and CO2 concentrations in matrix glasses and glass inclusions were determined by Fourier transform infrared spectroscopy (FTIR), on areas of matrix glass and inclusions that were completely microlite free. Measurements were made with a Nicolet 6700 Fourier transform infrared spectrometer (FTIR), with a Nicolet Continuum FTIR microscope, at Massey University, NZ. Transmission IR spectra in the range 4000–1200 cm − 1 were measured using a KBr beamsplitter, and a liquid nitrogen cooled HgCdTe2 (MCT) detector. Peak heights for total H2O at 3530 cm − 1, molecular H2O at 1630 cm − 1, and carbon dioxide as CO3− at 1515 and 1430 cm − 1, were measured graphically on straight line backgrounds using the OMNIC software. Peak heights were converted to concentrations using Beer's Law, with molar absorptivity values of 63 ± 3 l mol − 1 cm − 1, 28 ± 2 l mol − 1 cm − 1, and 353 ± 7 l mol − 1 cm − 1 for the 3535 cm − 1, 1630 cm − 1, and 1515– 1430 cm − 1 doublet, respectively (Dixon et al., 1995; Dixon and Pan, 1995; Dixon and Stolper, 1995; Dixon and Clague, 2001). The Beer's Law conversion from absorptivity to concentration is sensitive to the FTIR beampath, or sample thickness. FTIR for matrix glasses was on doubly-polished glass chips. Thicknesses at each analytical spot were determined using the spacing of interference fringes on FTIR reflectance spectra (Nishikida et al., 1996). Transmission and reflectance spectra were collected in-turn, without moving the sample, ensuring that the thickness at each analytical spot was determined precisely, as opposed to using a micrometer-measured average sample thickness (Wysoczanski and Tani, 2006). FTIR of inclusions was on singly-exposed inclusions in doubly-polished crystals, following the olivine spectra subtraction technique of Nichols and Wysoczanski (2007). Transmission and reflection spectra were collected in-turn through each inclusion and then through pure host olivine directly adjacent to each inclusion, without changing the crystallographic orientation of each olivine between inclusion and host spectra collection. Conversion from interference fringe wavelength to thickness requires refractive index of the given material, for which we used a fixed value of 1.546 for basalt (Kumagai and Kaneoka, 2003), and the linear relationship between olivine forsterite content and refractive index in Deer et al. (1992). Readers are directed to Nichols and Wysoczanski (2007) for formulations and procedures of the olivine subtraction technique. Glass inclusions ranging in radius from 21 to 144 μm were examined under petrographic and scanning electron (SEM) microscopes. Image J software was used to measure the area of inclusions and their included vapor bubbles (where present) on transmitted light photomicrographs of FTIR sections. Areas were converted to equivalent circular (and thus, spherical) radius. Saturation pressures (Psat) and equilibrium vapor compositions (CO2v) were calculated using the VolatileCalc software (Newman and Lowenstern, 2002), which is appropriate for tholeiitic basalts at pressures b100 MPa (Papale et al., 2006; Moore, 2008). Errors for EMP analyses were calculated for each analytical point by the Moran Scientific data reduction software, incorporating counting statistics, matrix effects and peak overlaps. These errors were combined in quadrature to yield standard errors for each element analyzed, shown as general error bars on figures as appropriate. Error bars on individual raw data points represent the standard deviations determined from replicate analyses of individual glasses and inclusions. Standard errors for FTIR measurements become larger, and more difficult to quantify, as peak resolution depends in part on the thickness of the sample. This is especially true for olivine-hosted inclusions, where the inclusion glass commonly makes up only a small proportion of the total beampath (Nichols and Wysoczanski, 2007).

In this case, in addition to uncertainties of sample thickness and absorptivity coefficient, there is error generated by the fact that thin inclusions yield small absorptions. For example, for an inclusion with pathlength of 20 μm, a change of 0.005 absorption units can change reported CO2 concentrations by N100 ppm. Absolute error, and detection limits for inclusion are thus higher than for matrix glasses. Again, standard deviations for replicate glass FTIR analyses are included on figures where appropriate. 3. Results 3.1. Deposit characteristics The deposit, which has a probable volume on the order of 10 5 to 10 6 m 3 (see Section 5.1), was examined in a ~22-meter thick section of pyroclastic material, with a base at approximately 1170 mbsl (Fig. 1C). Our stratigraphic observations are broadly in line with those of Clague et al. (2003), who also examined this deposit (their “Deposit D”) during an earlier dive series. Deposit character changes upsection through four zones (Fig. 2). (1) The basal few meters (not shown) are characterized by decimeter-thick beds of coarse ash separated by laterally discontinuous centimeter-thick layers of lapilli-dominated material. Samples collected from this sequence (162-1-A, 162-1-B, and 162-1-C) lacked coarse lapilli suitable for full textural analysis. (2) Above this basal sequence, pillow lava lobes ~20 to 30 cm in diameter, can be seen in the outcrop (Fig. 2B), revealing a mixed zone of pillows and lapilli-dominated, unconsolidated volcaniclastic material. The pillow lobes are extremely sparse, but visually striking in contrast to the surrounding clastic material. Pillow surfaces appear smooth, not broken or brecciated except where they are truncated in the scarp face. Volcaniclastic material in this section is dominated by vesicular lapilli in weakly grainsize-defined layers (Fig. 2C). (3) Upsection, no flow lobes were observed, and the deposit coarsens to thicker lapilli-dominant layers that are variably bedded (Fig. 2E) or structureless. In some cases, what appear at first to be pillow lavas are in fact massive, coarse clast-supported volcaniclastic units forming bulbous surface exposures (Fig. 2D,F). (4) The upper ~ 12 m of the deposit is unconsolidated, lapilli-rich volcaniclastic material in crude beds defined by layers of coarser fragments a few clasts thick, with some beds containing dispersed small angular blocks. Throughout the deposit, bedding is vague (e.g., Fig. 2C,E), but some units can be traced laterally over several 10s of meters. 3.2. Physical characteristics of pyroclasts 3.2.1. Density and vesicularity Over the entire ~22 m section sampled, median clast densities range from 1780 to 1990 kg/m 3 (30 to 37% vesicularity). Most samples have broad modal peaks corresponding to a vesicularity range of 30 to 40% (Fig. 2). For the purposes of description, we describe the lapilli according to three bulk density categories; low (LD), modal (MD), and high density (HD), with vesicularities of N~40%, ~30–40%, and b~30%, respectively. 3.2.2. Lapilli textures Lapilli surface textures vary between smooth vitreous areas with open vesicles (Fig. 3A), and dull gray zones with a rough texture (Fig. 3B). Both surfaces are often present on single clasts. Cavities and irregularities on lapilli surfaces often contain accumulations of sideromelane ash (Fig. 3C). LD clasts have generally sub-spherical vesicles, set in a groundmass that is light brown, microlite-free, sideromelane, opaque clinopyroxene microlite-rich tachylite, or transitional material. The relatively low density of most of these clasts can be attributed to a visually prominent population of large, sub-spherical vesicles (e.g., Fig. 3D).

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Fig. 2. Stratigraphy and clast density. Numbered sections correlate with sequence descriptions in Section 3.1. A: Reconstructed mosaic of video framegrabs from the Pisces IV cameras as collected during multiple upsection transects, showing the deposit in four sections. Placement of sample numbers shows where each was collected. Due to distortion in the reconstruction, depth scale is approximate and includes several breaks. B–F: Close-up video framegrabs of specific deposit features described in the text. Density/vesicularity histograms for seven samples shown in relative stratigraphic positions. Sample numbers and number of clasts measured form each given in each histogram panel. Shaded bar across histograms shows 30–40% modal vesicularity range.

MD clasts (Fig. 3E) have variable vesicle populations within groundmass that grades irregularly between sideromelane and tachylite. Groundmass types are usually not distributed symmetrically

with respect to clast edges. Some modal density clasts have sideromelane groundmass throughout, but few have tachylite groundmass throughout. Vesicles in MD clasts are generally sub-spherical, but

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Fig. 3. Textural features. A–B: Typical lapilli in hand sample, showing A) vitreous, and B) dull, rough surface textures. C: Random collection of sideromelane ash particles extracted from cavities in surface of clasts. D–E: Representative LD, MD, and HD clasts in thin section, common scale. G–H: Transitions in microcrystallinity, with pseudo-spherulitic microlites shown in transmitted light (G), and the sideromelane–tachylite transition shown with an overview SEM (H). Red box in panel H highlights deformed vesicle impinged upon by microlite network. I–J: Photomicrographs of olivine-hosted glass inclusions. Microlite-free sideromelane appears as brown, transparent glass, whereas microlite-rich tachylite groundmass appears as black, opaque regions. Note that (I) within olivines set in sideromelane, inclusions are always also sideromelane, but that (J) many of the olivines set in microcrystalline tachylite host inclusions that are also microcrystalline. Only microlite-free inclusions (as in G) were analyzed in this study.

in tachylite have irregular walls due to interaction with surrounding microlites (e.g., Fig. 3H), and in sideromelane are sometimes slightly elongated or in polylobate forms indicative of coalescence.

The walls of vesicles set in tachylitic groundmass of many LD and MD clasts are often penetrated by asperities of sideromelane. These enigmatic features are similar to an embryonic form of segregation

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vesicle imagined (but not observed) by Smith (1967, his Fig. 2A), and are the subject of continued study (Schipper et al., submitted for publication). All HD clasts (Fig. 3F) observed have tachylite-dominated groundmass. HD clasts do contain abundant, sub-spherical vesicles, but these are very small and thus do not substantially contribute to clast vesicularity. The microlites in tachylite are exclusively clinopyroxene. Microlite habits are variable, ranging from pseudo-spherulitic clusters that are overgrown in densely microcrystalline regions and become steadily more widely spaced moving from tachylite into sideromelane (Fig. 3G); to feathery, dendritic forms that also steadily decrease in abundance through irregular, but gradational, tachylite–sideromelane transitions (Fig. 3H). Olivine is generally euhedral, and is the only macroscopic crystal phase observed in the lapilli. Except for the HD clast having the highest crystallinity (~ 28%; Table 1), phenocryst content does not correlate with vesicularity or clast density. Olivine crystals range from ~ 0.25 to 9 mm on their longest 2-D sectioned axes. Glass inclusions in crystals are usually sub-spherical, and often contain Cr-spinel crystals and vapor bubbles. Inclusions range from ~ 20 to 140 μm radius. Vapor bubbles are proportional in size to their inclusions, indicating that they are the result of differential contraction of the host crystal and inclusion during cooling (Roedder, 1979, 1984). Vapor bubbles make up from 0 to 15% of the total inclusion volume; however, since the thermal volume expansion coefficient of basalt through the glass transition (Ryan and Sammis, 1981) is slightly less than an order of magnitude greater than that of highly magnesian olivine (Deer et al., 1992), bubble volumes larger than ~ 10% have probably been overestimated during our 2-D measurements. Where host olivines are set in transitional or sideromelane groundmass glass, their inclusions are always composed of clear glass (Fig. 3G); however, where set in highly microcrystalline tachylite, their inclusions are often also rich in microlites (Fig. 3H). Only inclusions composed of microlite-free glass (e.g., Fig. 3G) were analyzed in this study.

3.2.3. Vesicle size distributions Melt-corrected vesicle number densities (Nv = number of vesicles per unit volume of melt only) range over an order of magnitude, from 1.6 × 10 4 to 9.5 × 10 5 cm − 3; maximum vesicle sizes, expressed as equivalent diameters (EqD) range from 1.2 to 5.2 mm; median EqD ranges from 0.33 to 2.4 mm (Table 1). Quantitative vesicle parameters Nv and median EqD inversely and directly correlate with bulk vesicularity, respectively (Fig. 4, General).

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Vesicularity was measured across clast thin section areas, but a qualitative note is that the largest vesicles in MD lapilli tend to be in clasts with significant proportions of tachylite groundmass. LD clasts show polymodal vesicle volume distributions, with the largest vesicles contributing disproportionately to overall vesicularity. Truncations of the LD clasts' vesicle volume histograms toward the large vesicle sizes indicates that the magma contained still larger vesicles that are not represented within the domains of melt preserved as lapilli. Vesicle volume histograms in MD clasts are highly variable, but are generally truncated toward larger vesicle sizes, and show a greater diversity of vesicle sizes than do those of LD clasts. The HD clast yields a low, broad, vesicle volume histogram. Vesicle size distributions (VSD) for all clasts can be relatively well described by power laws of the form Nv(V)∝V − d (Fig. 4; Table 1), with R 2 values ranging from 0.89 to 0.96, taken across all vesicle size classes. The power law exponents (d) range from 1.3 to 2.5, and show a rough inverse correlation with bulk clast vesicularity. 3.3. Geochemistry 3.3.1. Whole rock geochemistry Lapilli have tholeiitic whole rock compositions (Table 2; Fig. 5A), with high MgO (18.7 to 22.6 wt.%), reflecting their high olivine content. Published MgO ranges for Lō`ihi whole-rock samples span 3 to 25 wt.% (Frey and Clague, 1983; Hawkins and Melchior, 1983; Garcia et al., 1995; 1998; Schipper et al., 2010b); the low-MgO samples are aphyric, and the high-MgO samples contain accumulated olivine (Garcia et al., 2006). 3.3.2. Olivine compositions Inclusion-bearing olivine crystals represent two different compositions. Most of the analyzed crystals are highly magnesian, and essentially unzoned, with core compositions ranging from Fo88.4 to Fo89.7 and rims from Fo87.8 to Fo89.9 (Table 4). Three sub-millimeter, less primitive, analyzed crystals have core compositions of Fo86.4 to Fo86.5. CaO content of olivine crystals (data not shown) ranges from 0.18 to 0.40 wt.%, indicating that they crystallized at crustal depths and are not mantle xenocrysts (Ren et al., 2005; Garcia et al., 2006). 3.3.3. Matrix glasses Glasses from Lō`ihi Seamount (Fig. 5A) have been studied extensively (Moore et al., 1982; Byers et al., 1985; Garcia et al., 1989; 1995; 1998; Kent et al., 1999a,b; Clague et al., 2000; Dixon and Clague, 2001; Clague et al., 2003; Schipper et al., 2010c,b) because of their widely ranging compositions, and because they better

Table 1 Vesicle and phenocryst quantification. Clast

161-4-A(7) 164-1-B(30) 164-1-F(3) 164-1-A(4) 164-1-B(9) 164-1-B(11) 164-1-C(1) 164-1-D(1) 164-1-E(6) 164-1-G(5) 164-1-G(8) 164-1-D(16)

Density class

LD LD LD MD MD MD MD MD MD MD MD HD

Vesicles

Olivine phenocrysts

Ves. (%)

Nv (cm–3)

Max EqD (mm)

Median EqD (mm)

45 47 42 38 34 34 34 31 29 27 28 18

1.55E + 04 5.88E + 04 9.96E + 04 2.48E + 05 2.37E + 05 3.21E + 05 1.31E + 05 8.20E + 04 1.46E + 05 2.81E + 05 1.34E + 05 9.49E + 05

4.11 2.45 4.12 3.27 5.19 3.92 2.03 1.62 4.16 2.18 3.29 1.19

2.40 1.05 1.10 0.62 1.03 0.63 1.00 0.58 0.90 0.65 0.63 0.33

Power law R2

d

0.92 0.89 0.96 0.94 0.93 0.96 0.94 0.93 0.94 0.94 0.95 0.95

1.3 1.8 1.7 2.4 2.0 2.3 1.5 2.3 1.9 2.2 2.5 2.4

Cryst. (%)

NA (cm–2)

Max EqD (mm)

Median EqD (mm)

23 10 23 15 14 23 17 20 16 14 14 28

31 60 13 22 23 17 20 49 14 26 33 68

3.39 2.36 4.43 2.15 3.76 3.09 2.13 2.52 3.85 4.32 4.3 4.65

0.41 0.33 0.89 0.67 0.61 1.1 0.66 0.43 0.99 0.58 0.45 0.32

Density classes: LD = low density, MD = modal density, HD = high density; Nv = number density of vesicles per unit volume of melt; EqD = equivalent diameter; d = power law exponent; NA = number density of crystals per bulk unit area.

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Fig. 4. Vesicle size distributions. For each clast, vesicle volume distribution (gray bars), and cumulative vesicle size distribution (VSD; Nv(V) N V; circles) shown on common (left side) vertical axis, against vesicles size bins shown as vesicle equivalent diameter of a sphere (EqD). Parameters of R2 and power-law exponent (d) for each VSD are shown, taken through all data points.

represent melt compositions than do whole rock analyses. Lapilli matrix glasses from this deposit (Table 3) are tholeiitic (Fig. 5A), with MgO from 6.13 to 8.43 wt.%, matching the compositions reported by Clague et al. (2003) for this same deposit. All matrix glasses listed in Table 3 are sideromelane, except for P4-164-1-B(30) and P4-164-1-F (3), which are tachylite over which the EMP beam was rasterized to yield similar compositions to the other glasses. Sulfur concentrations in lapilli matrix glasses range from 509 to 1025 ppm, all higher than the ~ 250 ppm upper limit inferred for subaerially erupted Hawaiian basalts (Moore and Thomas, 1988; Davis et al., 1991; Clague et al., 2000). Chlorine varies widely, from 371 to 1857 ppm. Cl is not expected to degas from basaltic melt until relatively low (b10–20 MPa) pressures (Dixon et al., 1991; Metrich and Wallace, 2008), but may also be an indicator of melt contamination by seawater-derived components (Michael and Schilling, 1989; Kent et al., 1999a,b). Matrix glass H2O concentrations range from 0.41 to 0.82 wt.% (average of 0.60 wt.%), and CO2 concentrations range from below the ~20 ppm

lower limit of FTIR detection, to 40 ppm. There is no correlation between clast vesicularity and measured volatile contents of matrix glasses. Matrix glass compositions are in equilibrium with olivine compositions from Fo76.1 to Fo81.8, and indicate eruption temperatures (Terupt) of 1093 to 1193 °C (Ford et al., 1983; Falloon and Danyushevsky, 2000; Danyushevsky, 2001). 3.3.4. Glass inclusions Major elements in glass inclusions (Table 4) are more variable than in matrix glasses, and do not, in their raw state, appear to define a single parental composition (Fig. 5). Raw EMP analyses show the inclusions to be as, or more, differentiated than the matrix glasses, with MgO ranging from 3.57 to 8.52 wt.%. The low MgO, low but widely ranging FeO T (7.51 to 12.0 wt.%) as well as high SiO2 (up to 53.3 wt.%) and CaO (up to 14.3 wt.%) in the inclusions indicates that some of them have experienced significant post-entrapment

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2002). Uncorrected H2O concentrations in inclusions range from 0.52 to 0.98 wt.% and CO2 concentrations are generally low, ranging from 29 to 481 ppm.

Fig. 5. Geochemistry. A: Total alkali-silica diagram. Gray field shows the range of previously reported Lō`ihi matrix glass compositions (Garcia et al., 2006). Field for compositions reported by Clague et al. (2003) for their “Deposit D” (not shown) corresponds exactly to the matrix glass compositions reported here. B: CaO/Al2O3 vs. MgO. Arrow shows trend of 30% olivine (Ol) fractionation in 5% increments from composition of least-differentiated inclusion after correction for PEC + Fe-loss. Slight trend of clinopyroxene (Cpx) fractionation is apparent in some matrix glasses.

modification (Sobolev, 1996; Danyushevsky et al., 2000; Geaetani and Watson, 2000; Danyushevsky et al., 2002a; Norman et al., 2002; Kent, 2008 and references therein). Sulfur in inclusions ranges more widely than in matrix glasses, from 367 to 2269 ppm. Raw Cl concentrations also range widely, from below detection to 3233 ppm. Even the highest measured Cl concentrations are modest compared to some of the extreme enrichments observed in other Lō`ihi inclusions (e.g., 1.11 wt.%; Hauri,

Table 2 Whole rock geochemistry. Clast

SiO2 TiO2 Al2O3 FeOT MnO MgO CaO Na2O K2O P2O5 LOI Total FoEqm (%)

P4-164

P4-164

P4-164

P4-164

1-C(5)

1-E(8)

1-C(15)

1-E(3)

46.3 1.97 8.66 11.4 0.17 18.7 8.13 1.24 0.52 0.06 0.06 97.2 91.5

45.7 1.85 8.15 11.7 0.17 20.6 7.64 1.09 0.48 0.06 − 0.21 97.3 92.1

45.3 1.81 7.95 11.8 0.17 21.4 7.48 1.07 0.48 0.20 − 0.17 97.6 92.4

45.2 1.70 7.49 11.7 0.17 22.6 7.02 0.94 0.45 0.07 − 0.29 97.3 92.8

Analyses by XRF. All oxides as wt.%. Total Fe as FeOT. FoEqm is the olivine composition in equilibrium with the measured compositions (Ford et al., 1983; Falloon and Danyushevsky, 2000; Danyushevsky, 2001) , using oxygen fugacity of QFM +0.5 (Green et al., 2001).

3.3.5. Corrections for post-entrapment modifications Post-entrapment modification of inclusions, especially those in highly magnesian olivine, can proceed by several different processes that are controlled in large part by the host crystal's cooling history (e.g., Kent, 2008 and references therein). Post-entrapment crystallization (PEC) of olivine on inclusion walls is nearly ubiquitous in basalts (Kent, 2008). Inclusions may also suffer irreversible diffusive reequilibration (“Fe-loss”) with their host crystals (Sobolev, 1996; Danyushevsky et al., 2000; 2002a,b), and/or with the evolving bulk melt surrounding the host crystals (Gaetani and Watson, 2000), and/or may lose H2O by diffusion through the crystal lattice (Danyushevsky et al., 2002b; Portnyagin et al., 2008). Inclusions corrected for PEC alone (data not shown) define a linear trend in FeO T-MgO space (Fig. 6A), strongly suggesting that they have experienced Fe-loss. We have corrected for combined PEC + Feloss following Danyushevsky et al. (2000; 2002b), using the FE_EQ2 (Danyushevsky et al., 2000) and Fe-loss (Danyushevsky et al., 2002b) software applications provided by those authors. The correction requires the user to define the melt oxidation state, and estimate the Fe concentrations in the inclusions at the time of entrapment, both of which may mask true variation in the inclusion compositions. We used constant Fe 2+/Fe 3+ of 8.5, consistent with an oxidation state of QFM +0.5 for Lō`ihi melts (Green et al., 2001), and defined FeO T as 12 wt.%, consistent with the inclusion compositions following FeO T versus MgO trends defined by whole rocks and matrix glasses (Fig. 6A). The degree of Fe-loss experienced by individual inclusions depends on their size and the residence time of their host crystals in cooler bulk melt (Danyushevsky et al., 2000; 2002a), and in this study ranges from 0 to 72% (see caption in Fig. 6). Calculated entrapment temperatures range from 1205 to 1314 °C (hydrous; Table 4; Fig. 6C). The time over which the inclusions re-equilibrated while their host crystals were in contact with the 1136 ± 50 °C bulk melt ranges from 0 days (for lower-Mg phenocrysts) to ~ 100 days (Table 4; Fig. 6D), as calculated by the graphical procedure described by Danyushevsky et al. (2002b). For these calculations, we used the calibrated re-equilibration time curves for Siquieros MORB, which is compositionally the closest proxy to our Lō`ihi tholeiite for which data is available (Danyushevsky et al., 2002b). Calculated residence times are minima, since inclusions were measured in 2-D, and their diameters will have been underestimated in cases where they were not perfectly spherical, and/or their largest cross-sectional area was not parallel within the sectioning plane. Corrected concentrations of S, Cl, H2O, and CO2 are calculated as a single-step reverse fractionation of olivine, according to the PEC % calculated in the correction procedure (Danyushevsky et al., 2000). Because of the strong degrees of calculated PEC + Fe-loss, this significantly reduces the corrected concentrations of all volatile elements, with post-correction S ranging from 258 to 1612 ppm, Cl from below detection to 2506 ppm, H2O from 0.36 to 0.74 wt.%, and CO2 from 26 to 296 ppm. Since H2O is in most cases lower in inclusions than in matrix glasses, and since inclusions are free of magnetite dust (Kent, 2008) and transparent rims indicative of H2O depletion (Danyushevsky et al., 2002b) we infer that they have not experienced post-entrapment diffusive H-loss. There is a temptation to correct inclusion volatile contents for the presence of vapor bubbles, which may partly represent volatile species lost from entrapped melt, and unaccounted for in analysis of glass inclusions (e.g., Anderson et al., 1989). In this study we have not used such a correction, because: (1) there is not yet a standardized method by which to make the correction; and (2) it is questionable to assume that vapor bubble formation is necessarily accompanied by

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C.I. Schipper et al. / Journal of Volcanology and Geothermal Research 207 (2011) 16–32

Table 3 Matrix glasses. Clast

Depth (mbsl) Ves. % EMP SiO2 TiO2 Al2O3 FeOT MnO MgO CaO Na2O K2O P2O5 S Cl Total n FTIR Thickness (μm) H2Ot H2Om CO2 n FoEqm (%) Terupt Psat CO2v

P4-161

P4-161

P4-161

P4-161

P4-161

P4-164

P4-164

P4-164

P4-164

P4-164

P4-164

P4-164

P4-164

4-A(3)

4-A(4)

4-A(7)

4-A(10)

4-A(25)

1-A(4)

1-A(7)

1-A(20)

1-A(27)

1-B(6)

1-B(9)

1-B(30)

1-C(1)

1165 38

1165 30

1165 45

1165 22

1165 37

1168 38

1168 28

1168 29

1168 30

1167 25

1167 34

1167 47

1166 34

48.8 2.93 12.6 11.9 0.15 7.24 11.9 1.97 0.79 0.29 877 391 98.6 5

48.0 2.87 12.4 11.7 0.20 7.41 11.5 2.10 0.84 0.45 939 645 97.6 5

147 0.63 0.06 35 3 79.2 1111 11.2 68

139 0.77 0.08 – 4 80 1116 9.9 44

49.4 2.87 12.9 11.6 0.16 7.13 12.2 2.35 0.80 0.25 908 469 99.8 10

79.6 1172

48.6 2.86 12.5 11.8 0.17 7.49 11.6 2.27 0.79 0.30 759 704 98.6 5

48.6 2.74 12.1 11.4 0.13 7.39 12.3 1.85 0.69 0.27 728 528 97.6 4

48.8 2.80 12.3 12.0 0.17 8.05 11.5 1.97 0.69 0.33 839 375 98.7 5

48.2 2.83 12.4 12.0 0.20 7.99 11.3 2.08 0.68 0.26 720 489 98.0 5

48.0 2.80 12.3 12.2 0.17 7.87 11.1 2.08 0.73 0.35 1010 606 97.7 5

48.5 2.80 12.5 11.6 0.21 7.84 12.0 2.11 0.68 0.27 728 1329 98.7 5

48.6 2.86 12.6 12.3 0.21 7.63 11.7 2.18 0.70 0.29 1025 1036 99.3 5

49.5 2.79 13.3 11.4 0.16 7.36 11.5 2.38 0.75 0.22 866 770 99.5 8

102 0.66 0.05 25 4 80.1 1122 9.4 58

147 0.65 0.07 24 3 80.2 1113 9.1 57

155 0.72 0.05 22 18 80.1 1132 9.6 50

80 0.71 0.04 24 2 80.1 1134 9.9 53

102 0.41 0.03 – 2 80.2 1144 6.1 72

96 0.46 0.03 22 2 81.1 1135 6.9 69

140 0.65 0.05 – 2 79.7 1123 8.2 53

133 0.62 0.04 – 4 80.3 1119 7.8 55

49.2 3.07 13.5 12.2 0.13 6.35 11.4 2.40 0.96 0.29 538 537 99.6 4

76.5 1152

48.6 2.81 12.3 11.8 0.17 8.05 11.4 1.92 0.70 0.37 686 461 98.2 6

232 0.47 0.05 40 5 81 1141 10.9 79

Clast vesicularity (ves.) measured by Archimedes' principal (Houghton and Wilson, 1989). Major elements by EMP, volatiles by FTIR. All oxides in wt. %, S, Cl, and CO2 in ppm. Compositions are average of n analyses. FTIR thicknesses are average for n analysis spots. H2Ot and H2Om are total, and molecular water, respectively. Saturation pressure (Psat) in MPa, and equilibrium vapor composition (CO2v) in mol. % calculated using VolcatileCalc (Ford et al., 1983; Newman and Lowenstern, 2002; Danyushevsky, 2004). FoEqm and Terupt are the composition of the olivine in equilibrium with the melt, and the estimated temperature (°C) of the melt during eruption, respectively (Ford et al., 1983; Falloon and Danyushevsky, 2000; Danyushevsky, 2001, 2004).

equilibrium diffusive mass transfer of volatiles out of the included melt, meaning that corrections based on this assumption risk systematically overestimating concentrations of volatiles trapped in the inclusion. For the corrected inclusion compositions, absolute uncertainties are large, and not accurately quantifiable. They include add-on errors at each step of the analytical and correction procedure, discussed by Danyushevsky et al. (2000, 2002a). We attempt to deal with this difficulty by presenting inclusion data at several stages of the correction processes; showing raw, PEC-corrected, and Fe-loss + PECcorrected data as appropriate.

4. Interpretation 4.1. Interpretation of deposit characteristics The well-exposed stratigraphy of the deposit allows a first-order assessment of the eruptive sequence, with the caveat that the vent location is not known. The ~ 22 m succession of volcaniclastic material is a thick accumulation of coarse material, and is therefore likely proximal to its single, or multiple but closely related, source(s). The nearly identical matrix glass composition throughout the deposit (Fig. 5) is strong, although not unequivocal, evidence that the deposit represents the products of a single, co-evolved, batch of magma. This interpretation owes much to the fact that a wide diversity of glass compositions is present in surficial deposits on Lō`ihi, with the full spectrum of alkalic to tholeiitic compositions often represented within the suite of rocks collected in single dredge hauls (Moore et al., 1982), or in the suite of particles within other thick volcaniclastic deposits (Clague et al., 2003).

Variations in the total deposit thickness and bedding dip probably result from infilling of paleotopography. The few pillow lava lobes observed in the lower, mixed zone of the deposit (Fig. 2) are the only indication of effusive activity in the immediate area, and the lavas have morphologies similar to submarine lava flow termini (Batiza and White, 2000). Above the lava lobes, there is no evidence for later effusive activity accompanying the semi-continuous emplacement of coarse volcaniclastic material. It is possible that isolated, low-volume flow lobes (as in Fig. 2B) were co-eruptively extruded through the base of the growing accumulation of unconsolidated volcaniclastic material (Valentine et al., 2007; Pioli et al., 2008). The crude layering in upper sections does not have waveforms or other indicators of transport in density currents. Discontinuous layering and lack of bimodal clast density distribution is evidence against deposition having been from any significant eruptive plume, in which case clasts would be sorted according to their hydrodynamic properties (e.g., Cashman and Fiske, 1991). We suggest that the lapilli-sized ejecta were in general too coarse to be entrained upward in- and laterally transported by- a buoyant plume, and instead were emplaced pseudoballistically; that is, ejected forcefully from the vent, but deposited by settling through the water column, with little convective (plume) entrainment. Fine sideromelane ash cached in cavities on the surface of lapilli (Fig. 3C), however, suggests contemporaneous ejection of ash and lapilli-sized material. No horizons of clay, very fine ash, pelagic sediments, or other time break indicators were observed in the stratigraphy, marking a key difference between this deposit and another thick, fault-exposed volcaniclastic sequence to the south, which has been the subject of several studies (Clague et al., 2003; Clague, 2009; Schipper and White, 2009). Together, the above interpretations suggest that the sequence was emplaced relatively rapidly, possibly from a single sustained eruption,

C.I. Schipper et al. / Journal of Volcanology and Geothermal Research 207 (2011) 16–32

25

P4-164

P4-164

P4-164

P4-164

P4-164

P4-164

P4-164

P4-164

P4-164

P4-164

P4-164

P4-164

P4-164

P4-164

1-C(2)

1-C(55)

1-C(96)

1-D(1)

1-D(3)

1-D(7)

1-E(6)

1-E(16)

1-E(45)

1-F(3)

1-F(15)

1-F(20)

1-F(24)

1-G(22)

1166 44

1166 29

1166 13

1163 31

1163 33

1163 30

1162 29

1162 33

1162 19

1158 42

1158 50

1158 32

1158 21

1148 18

48.3 2.81 12.3 12.0 0.15 8.08 11.3 2.14 0.69 0.28 571 371 98.2 9

48.3 2.77 12.3 12.1 0.18 8.01 11.3 2.06 0.68 0.29 838 645 98.2 5

48.0 2.76 12.3 12.0 0.18 8.04 11.2 2.03 0.79 0.37 773 880 97.7 4

48.6 2.78 12.4 12.0 0.16 8.03 11.4 2.16 0.68 0.27 814 391 98.6 5

49.4 2.66 12.6 11.5 0.17 7.76 11.9 2.02 0.70 0.38 886 1008 99.2 5

77 0.56 0.05 28 6 81 1143 9.2 66

116 0.58 0.03 37 3 80.6 1139 11.3 71

108 0.67 0.02 29 3 80.1 1138 10.4 60

145 0.67 0.05 30 10 80.1 1136 10.6 61

87 0.60 0.04 23 7 80.8 1125 8.3 61

49.2 2.80 12.7 11.7 0.21 7.91 11.5 2.06 0.69 0.37 839 435 99.2 5

80.9 1193

48.8 2.79 12.6 11.8 0.20 7.96 11.5 2.31 0.71 0.23 848 749 99.0 6

48.3 2.84 12.4 11.8 0.15 7.54 11.4 2.21 0.70 0.27 978 606 97.8 5

48.3 2.78 12.1 12.0 0.20 8.43 11.3 2.00 0.69 0.25 767 371 98.2 4

230 0.49 0.03 32 6 81.1 1142 9.4 74

113 0.76 0.05 28 3 80.1 1122 11.5 53

99 0.42 0.03 33 2 81.6 1157 9.0 80

or from a series of successive eruptions closely spaced in time. Using bed thickness and clast size as proxies for eruption intensity/discharge rate, the period of highest discharge is likely to correspond to the sequence of thick structureless volcaniclastic beds from ~1165 to 1169 mbsl (Fig. 2). 4.2. Interpretation of textural features 4.2.1. Bulk vesicularity The modest range in clast vesicularity in deposits near the bottom and top of the sequence (Fig. 2) appears to represent periods when heterogeneously vesiculated portions of magma were co-ejected from the vent(s). This heterogeneity could be a consequence of variable durations and conditions of vesicle nucleation and growth, coexisting at any given depth/time in the conduit (e.g., cross-conduit variability). The most uniform clast vesicularity is in samples taken from the massive volcaniclastic beds (Fig. 2), which is consistent with this portion of the sequence representing the period of most powerful and continuous magma discharge and clast deposition. 4.2.2. Vesicle textures There is an overall impression of vesicle maturity in lapilli of all density classes; that is to say that the textures are not monodisperse populations of small sub-spherical vesicles. It is well documented that in many pyroclastic rocks, vesicle populations can evolve for significant times after fragmentation. This is seen in rim-to-core coarsening of vesicle populations in: subaerial breadcrust bombs (e.g., Wright et al., 2007), Hawaiian scoria (e.g., Stovall et al., 2010), and submarine bomb fragments (e.g., Schipper et al., 2010a). In these cases, the “younger” textures in quenched rims are typically

49.3 2.80 12.9 11.4 0.15 7.04 11.8 2.31 0.78 0.23 816 572 98.8 6

79.6 1172

48.6 2.98 13.4 11.9 0.17 6.13 10.9 2.36 0.72 0.31 509 528 97.6 5

48.7 2.82 12.5 11.5 0.15 7.64 11.8 2.20 0.73 0.32 955 469 98.6 5

47.8 2.79 12.2 11.7 0.14 8.22 11.2 2.32 0.66 0.31 751 1857 97.6 3

48.8 2.93 12.6 11.9 0.15 7.24 11.9 1.97 0.79 0.29 877 391 98.6 5

114 0.42 0.02 29 3 76.1 1093 8.1 78

63 0.52 0.04 27 3 80.8 1130 8.6 69

70 0.68 0.07 27 3 81.8 1146 10.1 58

153 0.61 0.05 – 3 79.2 1111 7.7 56

characterized by vesicle shapes tending toward sub-spherical, as well as by having higher Nv and lower EqD. Contrasting clast rims are certainly not universal across pyroclast types, but the important point is that in the above examples, the more mature textures in clast interiors can be shown to have developed from the younger rim textures (Schipper et al., 2010a; Stovall et al., 2010), through postfragmentation bubble interaction and volatile exsolution. Such clear gradations are not, in general, exhibited by lapilli in this study. The presumably rapidly-quenched Lō`ihi lapilli presented here do not have “youthful” bubble populations at clast rims when viewed in thin section. It appears instead, that the bubbles in all the magma at the time of fragmentation had already matured through extensive, but varying (e.g., producing a wide observed range of vesicle textures; Fig. 3D–F) steps of nucleation, growth, coalescence, and partial loss that characterize the life cycle of bubbles in magma. Such maturation may have occurred during periods of magma stagnation and residence in the shallow subsurface below the vent between eruption events or pulses. A semi-quantitative strengthening of the above interpretation is provided by the power-law form of the VSDs (Fig. 4). Power law VSDs may develop from exponential distributions by additional nucleation of new bubbles (e.g., continuous nucleation, modeled as multiple nucleation events by Blower et al., 2001, 2002), or by coalescence (e.g., Gaonac'h et al., 1996). These different interpretations notwithstanding, power-law distributions indicate relatively mature bubble populations; those that have experienced processes beyond short-lived single-stage nucleation and subsequent free growth. For the relatively low viscosity basaltic magma being examined in this study, extensive bubble coalescence was almost certainly important in controlling the textural evolution of the erupted magma.

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Table 4 Glass inclusions. Crystal

L8–09

L8–10

L8–11

*L8–12

L8–13

L8–14

L8–15

L8–16

Inclusion

A

A

A

A

A

A

A

A

Radius (μm) % vap. EMP SiO2 TiO2 Al2O3 FeOT MnO MgO CaO Na2O K2O P2O5 S Cl Total n

66 8

69 6

51.0 50.4 2.31 2.66 12.1 13.6 9.51 8.68 0.22 0.13 7.60 7.51 12.7 12.1 2.22 2.28 0.43 0.74 0.23 0.43 367 1201 238 676 98.4 98.7 3 3

29 2

114 11

75 0

67 8

144 1

L8–17 B

46 9

51.9 48.9 2.88 2.48 12.8 12.7 8.23 11.5 0.04 0.21 8.27 8.09 13.2 11.7 2.26 2.13 0.21 0.36 0.00 0.14 774 1509 40 795 99.8 98.4 3 3

51.0 50.2 49.7 2.56 2.64 2.68 12.9 12.7 12.7 9.91 8.70 10.2 0.07 0.09 0.11 6.91 8.37 8.21 11.9 11.6 11.9 2.24 2.51 2.40 0.30 0.33 0.38 0.32 0.32 0.33 908 1135 1308 119 477 199 98.1 97.6 98.7 3 3 3

49.3 3.23 12.6 9.85 0.13 7.05 13.3 1.99 0.45 0.69 1762 715 98.8 3

49.1 3.16 12.6 10.6 0.17 7.72 12.1 2.34 0.64 0.39 1562 477 99.0 3

L8–22

L8–23

A

A

A

49.2 3.22 12.8 8.95 0.14 8.27 11.8 2.43 0.48 0.64 1308 199 98.1 3

50.0 2.90 12.8 8.95 0.08 8.48 11.8 2.45 0.52 0.56 1215 238 98.7 3

48.6 2.85 12.5 12.0 0.20 7.55 12.0 2.15 0.59 0.48 975 119 98.9 3

50.5 2.72 12.8 8.69 0.18 7.37 13.0 2.18 0.42 0.97 881 358 99.0 3

23 0.60 225 2

24 0.73 199 3

68 0.68 80 3

34 0.73 192 2

25 0.59 140 2

45 0.72 185 2

25 0.94 396 2

86.4

89.6 89.6

89.3 89.9

89.6 89.4

89.6

89.4

26 0.52 300 2

37 0.88 156 4

32 0.66 219 3

45 0.89 157 3

50 0.68 87 3

35 0.72 190 2

21 0.89 331 1

Host Olivine Core Fo.% Rim Fo.%

89.7 89.6

89.4 89.5

89.4 89.1

88.4 87.8

89.6 89.6

89.6 89.8

89.4 89.8

89.6 89.5

89.4 89.6

26.8 3

L8–21

A

73 4

33 0.64 236 1

47.3 2.50 9.98 12.0 0.17 15.5 9.59 1.85 0.51 0.31 1143 349

L8–20

A

98 2

77 0.70 56 3

47.3 2.48 9.65 12.0 0.13 15.5 10.3 1.53 0.35 0.53 1225 497 0.50 132 30.5 35 1303 6 19.0 86

L8–19

A

B

FTIR Thickness (μm) H2Ot CO2 n

Compositions corrected for PEC and Fe–loss SiO2 48.7 47.9 48.6 48.2 48.8 48.3 47.9 TiO2 1.79 2.04 2.20 2.14 1.97 2.07 2.15 Al2O3 9.31 10.4 9.75 10.9 9.94 9.97 10.2 T FeO 12.0 12.0 12.0 12.0 12.0 12.0 12.0 MnO 0.20 0.13 0.06 0.21 0.09 0.10 0.12 MgO 15.6 15.1 15.1 13.7 15.5 15.5 15.2 CaO 9.90 9.38 10.1 10.2 9.22 9.17 9.64 Na2O 1.72 1.75 1.72 1.84 1.73 1.97 1.93 K2O 0.34 0.57 0.16 0.31 0.23 0.26 0.30 P2O5 0.18 0.33 0 0.12 0.25 0.25 0.26 S 258 842 552 1247 630 813 982 Cl 167 473 28 657 83 342 149 H2O 0.49 0.45 0.37 0.73 0.46 0.64 0.51 CO2 39 165 214 129 152 113 66 %PEC 29.8 30.0 28.7 17.3 30.6 28.3 24.9 %Fe-loss 41 53 60 8 34 54 29 Ttrap 1309 1304 1304 1263 1306 1305 1303 tcool (days) 12 25 11 4 13 31 32 Psat 9.9 26.1 40.6 25.7 32.3 22.5 12.3 CO2v 75 92 96 80 93 83 78

*L8–18

A

47.5 2.55 10.14 12.0 0.14 15.2 9.39 1.93 0.38 0.51 957 145 0.65 242 26.9 49 1298 59 35.4 88

28 0

80 0

47 4

121 1

129 0

61 6

53.3 51.9 50.9 50.8 2.82 2.48 2.37 2.68 14.2 13.8 13.4 13.2 7.86 9.08 8.98 7.51 0.21 0.22 0.09 0.26 3.57 4.92 7.90 7.05 13.9 14.2 12.0 12.9 2.36 1.67 2.47 2.41 0.63 0.61 0.89 0.64 0.11 0.06 0.11 0.04 2269 1348 1001 1215 100 67 233 67 99.2 99.1 99.2 97.6 3 3 3 3

47.8 48.2 47.8 48.9 48.4 48.2 48.3 2.29 2.59 2.05 1.95 1.78 1.82 2.02 10.13 11.35 9.62 9.76 9.85 10.2 9.93 12.0 12.0 12.0 12.0 12.0 12.0 12.0 0.10 0.20 0.17 0.18 0.20 0.10 0.23 15.2 11.5 15.5 15.1 15.5 15.5 15.2 9.37 10.9 9.87 9.67 10.3 9.25 9.78 1.93 1.95 1.64 1.63 1.20 1.89 1.82 0.41 0.54 0.31 0.43 0.43 0.68 0.48 0.44 0.43 0.73 0.07 0.05 0.08 0.03 893 868 600 1277 825 704 810 175 106 243 56 41 164 44 0.44 0.65 0.46 0.41 0.36 0.51 0.63 165 177 55 108 85 130 264 26.5 10.9 31.9 43.8 38.8 29.7 33.3 7 0 54 66 47 49 72 1307 1219 1304 1303 1306 1314 1300 2 0 71 25 81 62 100 25.3 32.2 12.7 24.1 16.0 23.4 47.0 92 87 83 93 91 88 92

Inclusion radius and vapor bubble volume measured using Image J software. * Inclusions in lower-Mg phenocrysts. Major elements by EMP, volatiles by FTIR. All oxides in wt. %, S, Cl, and CO2 in ppm. Compositions are average of n analyses. FTIR thicknesses are average for n analysis spots. Saturation pressure (Psat) in MPa, and equilibrium vapor composition (CO2v) in mol. % calculated using VolcatileCalc (Newman and Lowenstern, 2002). Corrections for PEC + Fe-loss calculated with FEO_EQ2 (Danyushevsky et al., 2000) and Fe-loss (Danyushevsky et al., 2002b) software, with time of partial reequilibration (tcool) calculated using the graphical method of Danyushevsky et al. (2002b). Entrapment temperature Ttrap calculated for hydrous conditions (Ford et al., 1983; Falloon and Danyushevsky, 2000; Danyushevsky, 2001), and not given for inclusions with no volatile data.

4.2.3. Microlite textures The abundance and patterns of tachylite and associated features also yield important, but not unequivocal, information about magma ascent conditions. The hindrance to making unequivocal interpretations from microlite textures is that microlites can form in a variety of different ways. They can result from devitrification of volcanic glass at temperatures below the glass transition (Lofgren, 1971, 1974; Kirkpatrick, 1978), thus providing no information about syn-eruptive processes, or they can be “primary,” developing pre- or syn-eruptively in liquid melt above the glass transition (Monecke et al., 2004). Primary microlite crystallization is itself complicated, since it can occur either: (1) due to “effective undercooling,” or suppression of crystalline phase liquid by H2O exsolution during decompression/ degassing (Hammer and Rutherford, 2002), or even by H2O exsolution

induced by the passage of CO2-rich fluids fluxing permeably or passing as decoupled slugs (Newman et al., 1988; Blundy et al., 2010); or (2) due to a temperature drop (undercooling) in the upper conduit (Taddeucci et al., 2004) or even in the eruptive plume or deposit (Szramek et al., 2010). In this case, we rule out devitrification because the presence of segregation vesicles in some lapilli (Schipper et al., submitted for publication) indicates that microlite formation occurred while the bulk melt was still highly fluid (Smith, 1967). We rule out primary microlite formation by “effective undercooling”/ H2O degassing, because: (1) microlite densities are often gradational (Fig. 3G,H) within lapilli and do not correlate with bulk vesicularity (e.g., LD and HD have common, and ubiquitous tachylitic groundmasses, respectively, Fig. 3D, F); (2) there is no evidence for H2O degassing from the bulk melt

C.I. Schipper et al. / Journal of Volcanology and Geothermal Research 207 (2011) 16–32

L8–24

L8–25

A

A

85 5

40 3

52.5 2.01 12.6 8.39 0.17 8.29 12.1 2.48 0.33 0.00 1001 0 98.9 3

49.5 2.84 13.3 9.12 0.15 4.59 14.1 2.38 0.44 0.26 1482 133 96.8 3

B

49.9 2.99 13.7 9.36 0.22 4.11 13.7 2.26 0.89 0.06 1629 167 97.3 3

L8–26

L8–27

A

A

B

64 9

82 7

97 0

51.3 2.51 13.2 8.67 0.28 7.10 13.1 2.36 0.50 0.02 934 300 99.2 3

51.3 2.32 12.4 9.68 0.26 8.27 12.0 2.34 0.46 0.05 1642 0 99.2 3

51.8 2.42 12.5 8.54 0.27 8.06 12.3 2.45 0.48 0.05 1375 133 99.0 3

53 0.61 131 2

77 0.89 71 3

12 0.75 481 2

36 0.62 156 2

53 0.71 97 2

89.5

89.5 89.6

89.2 88.8

89.5 89.7

49.4 1.55 9.71 12.0 0.16 15.3 9.40 1.92 0.25 0.00 713 0 0.64 50 28.7 58 1300 80 14.6 74

47.6 2.09 9.78 12.0 0.15 15.3 10.5 1.75 0.32 0.19 913 82 0.46 296 38.4 46 1306 8 41.5 94

48.4 1.91 10.0 12.0 0.25 14.9 10.1 1.80 0.38 0.02 651 209 0.43 108 30.4 52 1300 30 20.4 90

48.8 1.83 9.77 12.0 0.23 15.3 9.48 1.85 0.37 0.04 1207 0 0.52 71 26.5 37 1305 27 16.9 84

47.7 2.17 9.95 12.0 0.19 15.3 10.1 1.64 0.65 0.05 983 101

39.7 6

48.8 1.86 9.60 12.0 0.24 15.3 9.54 1.88 0.37 0.04 971 94 0.43 92 29.4 8 1309 3 20.2 91

27

L8–28

L8–29

L8–30

P01

P02

P02

*P07

P08

P09

P11

P13

*P14

A

A

A

A

A

B

A

A

A

A

A

A

75 5

70 2

98 11

48 0

34 4

28 15

51.8 2.72 13.7 8.26 0.25 5.97 13.7 2.25 0.53 0.07 921 33 99.3 3

47.8 3.06 12.3 10.8 0.30 7.06 13.8 1.82 0.86 0.04 2136 867 98.2 3

48.7 2.85 12.6 10.9 0.26 7.15 13.6 1.93 0.40 0.02 574 3233 98.9 3

50.2 2.82 13.6 10.74 0.11 5.40 11.4 2.42 0.46 0.12 1055 367 97.5 3

50.4 2.96 12.7 8.47 0.11 6.56 13.8 2.21 0.35 0.34 1148 400 98.2 3

50.3 3.04 13.4 7.90 0.07 5.90 14.1 2.20 0.38 0.34 1522 367 98.0 3

48.9 2.74 12.5 11.8 0.14 7.99 11.8 2.27 0.45 0.23 988 467 99.1 3

48.3 2.68 12.0 9.11 0.13 7.74 13.8 2.14 0.12 1.85 908 233 98.1 3

48.0 2.09 12.7 10.1 0.09 7.50 14.4 2.23 0.27 0.43 1308 367 98.1 3

25 0.59 210 2

19 0.55 253 2

37 0.79 126 2

61 0.77 187 2

39 0.74 98 2

31 0.73 143 2

66 0.78 180 2

40 0.57 105 2

88.8 89.2

89.0 88.2

88.6 88.9

89.7

89.6

89.6

86.5

48.5 2.03 10.2 12.0 0.22 14.3 10.3 1.68 0.40 0.05 623 23 0.40 142 32.4 57 1287 39 26.8 94

46.8 2.49 10.0 12.0 0.28 14.5 11.3 1.48 0.70 0.03 1612 654 0.42 191 24.5 19 1294 5 22.9 91

47.5 2.35 10.4 12.0 0.24 14.1 11.2 1.59 0.33 0.01 445 2506 0.61 98 22.5 16 1274 8 16.3 78

48.4 2.14 10.3 12.0 0.11 15.7 8.72 1.84 0.35 0.09 694 241 0.50 123 34.2 21 1313 4 23.6 89

47.9 2.21 9.50 12.0 0.12 15.5 10.3 1.65 0.26 0.26 752 262 0.49 64 34.5 57 1306 17 11.8 79

47.7 2.22 9.75 12.0 0.09 15.5 10.4 1.61 0.28 0.25 949 229 0.45 89 37.6 66 1307 24 14.4 85

48.4 2.50 11.4 12.0 0.14 11.7 10.8 2.07 0.41 0.21 883 417 0.70 161 10.6 3 1205 0 32.1 85

(see Section 4.4); and (3) inclusions are assumed to be “pressure vessels” that do not decompress along with the surrounding bulk melt, yet inclusions within olivine crystals that are set in tachylite groundmass are commonly themselves microcrystalline (Fig. 3J). This latter point is supported also by the recognition that glass embayments (e.g., incompletely sealed inclusions) in olivine phenocrysts of many volcanic rocks often retain higher volatile contents than surrounding bulk melt, so that even “leaky” inclusions do not communicate fully with bulk magma and do not degas as completely during decompression (Metrich and Wallace, 2008). The remaining option is that the observed tachylite developed by heat loss (true undercooling) from the magma. The fact that microlite abundances do not increase systematically from rim-to-core of the clasts examined in this study suggests that they did not develop

58 7

24 4

21 1

116 6

32 4

69 0

49.2 2.86 12.3 10.8 0.13 8.52 11.8 2.26 0.54 0.24 1695 200 99.0 3

50.4 3.15 12.7 8.93 0.10 6.78 13.5 2.21 0.31 0.28 1362 500 98.8 3

48.2 2.81 12.4 11.7 0.15 7.01 12.2 2.28 0.47 0.14 1015 433 97.7 3

14 0.98 114 3

97 0.91 54 2

49 0.85 39 2

50 0.66 29 2

89.6

89.0

89.4

89.1

86.4

46.7 2.08 9.29 12.0 0.13 15.5 10.8 1.66 0.09 1.44 640 164 0.40 74 29.5 47 1302 11 18.7 81

46.9 1.70 10.3 12.0 0.10 14.6 11.7 1.81 0.22 0.35 984 276 0.74 86 24.8 30 1282 1 15.2 63

47.8 2.35 10.1 12.0 0.13 15.2 9.72 1.85 0.44 0.19 1312 155 0.71 41 22.6 20 1297 15 10.7 54

48.0 2.42 9.76 12.0 0.11 14.8 10.5 1.70 0.24 0.22 949 348 0.59 27 30.3 48 1288 10 7.6 53

48.4 2.56 11.3 12.0 0.16 11.6 11.1 2.08 0.43 0.12 887 379 0.58 26 12.6 4 1226 0 7.8 57

during post-fragmentation diffusive cooling (i.e., by cooling of a small volume of melt completely surrounded by water), as observed in other Lō`ihi submarine scoria (Schipper et al., 2010a). Our interpretation is that in the current case, microlites formed in response to comparatively slow, but variably-moving cooling fronts in the upper conduit (subsurface just below the vent(s)), with heat being lost primarily to the conduit margins in a process analagous to that recognized in the ejecta from classical subaerial Strombolian and violent Strombolian eruptions (Taddeucci et al., 2004; Polacci et al., 2006; Pioli et al., 2008; Andronico et al., 2009; Cimarelli et al., 2010). In subaerial cases, portions of magma with higher microlite contents are associated with longer conduit residence times (more cooling). The submarine location of the vent(s) that produced this deposit may have facilitated sub-surface magma cooling, with seawater infiltration

28

C.I. Schipper et al. / Journal of Volcanology and Geothermal Research 207 (2011) 16–32

Fig. 6. PEC + Fe-loss corrections. A: FeOT vs. MgO, showing the effects of different major-element corrections (data not shown for GiPEC). B: “Fe-loss triangle” showing degree of diffusive reequilibration each inclusion constructed following Danyushevsky et al. (2000). Upper limit (gray field) defined by fractionation trend of whole rock and matrix glass compositions. Trend I shows the maximum possible Fe-loss from inclusions entrapped at different stages of fractionation, defined at the low-Fo end by the intersection of the fractionation trend and FoEqm of the erupted melt, and at the high-Fo end by the modeled 100% Fe-loss from the inclusion in the most magnesian olivine in the suite. This same most primitive olivine has the potential for the greatest degree of re-equilibration, thus vertical Trend II is defined by its Fo%. Inclusion compositions that had experienced full (100%) reequilibration (none in this suite) would lie along Trend I. See Danyushevsky et al. (2000) for full discussion of triangle construction. C: Eruption and entrapment temperatures for matrix glasses and inclusions. D: Histogram of inclusion re-equilibration times.

through the saturated volcanic pile, particularly given the typically high, and isotropic, hydraulic conductivity of scoria cones (Josnin et al., 2007).

~3 months). Had the eruption been characterized by direct, uninterrupted ascent from the crystal cumulate zone (which is the most likely source of the high-Mg olivines) to the vent, we would in theory expect reequilibration timescales to be more uniform.

4.3. Olivine: crystallization, entrainment, and magma residence 4.4. Degassing of H2O and CO2 The olivine crystals are dominantly high-Mg (NFo88) phenocrysts, most of which contain diffusively reequilibrated inclusions entrapped at high temperatures (Fig. 6C), with only three analyzed lower-Mg (bFo86.5) phenocrysts containing inclusions entrapped at lower temperatures. The inclusions in high-Mg phenocrysts cannot be used to directly infer continuous processes from magma storage to vent, but may be used cautiously to infer the state of the magma at some point earlier and deeper in the magmatic system. Inclusions in lower-Mg phenocrysts represent the most probable parental compositions for the erupted bulk melt, at timeframes closer to the eruption. A histogram, on logarithmic timescale, of inclusion reequilibration times is shown in Fig. 6D. Within the errors and assumptions of the reequilibration calculation, we consider times less than about 1 week to be essentially equivalent to “no reequilibration.” This nonreequilibrated subset then includes inclusions in the three lowerMg phenocrysts as well as some in high-Mg phenocrysts. The time for reequilibration starts when each crystal is entrained into comparatively cool bulk melt, and ends at the time of quenching, but the physical location of the cooling process is not specifically known, and is likely to include time spent traveling up the magmatic system from depth, as well as static time in the shallow subsurface. Given the textural evidence for significant periods of shallow magma stagnation and evolution, it is not surprising that the high-Mg phenocrysts appear to have reequilibrated over a wide range of times (nil to

Major element heterogeneity in the high-Mg olivine inclusions (Fig. 5B) is common in suites of Hawaiian olivine-hosted inclusions, and is thought by some workers to represent a complex, fine-scale zonation of the Hawaiian plume that is not apparent in bulk glass or whole rock analyses of eruption products (Ren et al., 2005). Therefore, with the caveat that we estimated the initial FeO t of the inclusions using a correction for Fe-loss, we infer that the resulting inclusion compositions (GIcorr) do represent a proxy for parental tholeiite from which the bulk melt could eventually have evolved (Fig. 5B) by olivine fractionation. H2O can generally be assumed to drive ascent, vesiculation, and eruption at subaerial vents, but this assumption is not valid for deep submarine eruptions of Hawaiian basalt. Solubility studies indicate that pure H2O is soluble to ~ 1 wt.% in Hawaiian tholeiite (~49 wt.% SiO2) at 10 MPa unless a CO2-rich vapor phase remains in equilibrium with the melt (Dixon et al., 1995; Dixon and Stolper, 1995; Newman and Lowenstern, 2002), and the parental H2O content in Lō`ihi magmas is estimated to be no greater than 0.60 wt.% for more differentiated alkalic magmas (Dixon and Clague, 2001; Hauri, 2002). Strong, eruption style-determining, exsolution of H2O from water-undersaturated Lō`ihi magmas can be achieved only if a CO2-rich vapor phase remains in physical contact and equilibrium with the melt (Schipper et al., 2010c, b), either as isolated bubbles that remain coupled to the

C.I. Schipper et al. / Journal of Volcanology and Geothermal Research 207 (2011) 16–32

magma from which they have grown, or possibly as a decoupled vapor phase fluxing through permeable pathways in the magma (Newman et al., 1988). Because of the barriers to strong H2O exsolution at submarine vents, but low solubility of CO2, submarine explosivity is often inferred to be driven primarily by magmatic CO2 (Clague et al., 2008; Sohn et al., 2008). Lapilli matrix glasses at this site are nearly completely degassed of CO2, but retain widely varying H2O contents (Table 3, Fig. 7). The vesicularity of individual lapilli is not inversely correlated with melt H2O, as would be expected if a melt originally homogenous in H2O had vesiculated in a closed, volatile-coupled system (Schipper et al., 2010b). CO2 versus H2O is shown in Fig. 7A, illustrating the variation in volatile contents of both the inclusions and matrix glasses. H2O in matrix glasses extends to higher concentrations than in any of the corrected inclusions. Increases in H2O obviously cannot be achieved by degassing, but may partially be explained by the incompatible behavior of H2O during crystallization. Fig. 7B shows H2O versus MgO, with olivine fractionation contours starting from the lowest-, average-, and highest-H2O inclusions measured. This suggests that, if the range of H2O in inclusions represents true heterogeneity in parental melts, fractionation of olivine during ascent and differentiation could generate a bulk melt with similar short length-scale H2O heterogeneity of the same magnitude. We note, however, that there is a paradox in this oversimplified interpretation: any major element

29

heterogeneity in inclusion compositions appears to have been homogenized during evolution of the bulk melt, whereas centimeter-scale heterogeneity in H2O appears to have been preserved. Perhaps at some time between crystallization of high-Mg olivine and eruption, an intermediate version of the magma was subjected to volatile enrichment by assimilation of H2O-saturated country rocks or other hydrous components (e.g., Kent et al., 1999a,b). This was observed in the analysis of degassing at the northern cone (Fig. 1B) deposit of Lō`ihi (Schipper et al., 2010b), where “low-Fe” inclusions (which were not corrected for Fe-loss) also had low H2O, but “highFe” inclusions that were direct parents of the bulk melt had significantly elevated H2O from which closed-system degassing paths could be followed to generate the erupted melts (Schipper et al., 2010b). Furthermore, experimental work of Portnyagin et al. (2008), who showed that H2O can often diffuse in- or out- of olivine-hosted inclusions that cool over timescales of a few days, suggests that the H2O concentrations in the longer re-equilibrated inclusions may be underrepresented. The direct evidence against either of the above being issues here is the modest H2O (0.58 to 0.70 wt.%) in the three low-Mg phenocryst-hosted inclusions (Table 4). This, together with the fact that the average H2O content of the inclusions (0.52 wt.%) is close to the estimated parental H2O content for Lō`ihi tholeiites (0.48 wt.%; Dixon and Clague, 2001), there is no reason to suspect that at any point there was higher H2O in this particular magmatic subsystem of Lō`ihi. There are several inherently heterogeneous processes that may have generated H2O heterogeneity in the lapilli matrix glasses late in magma's ascent history. Anhydrous microlite formation in significant portions of the magma would create localized enrichments of H2O in periods of pre-fragmentation, upper conduit stagnation. Also, degassing of H2O from submarine pyroclasts lofted to different heights above explosive submarine vents may assist in the development of clast-toclast heterogeneity (Fouquet et al., 1998; Schipper et al., 2010c, b), as might the recycling of clasts back into the vent (Deardorff et al., 2011). Furthermore, any small degree of CO2 fluxing-induced degassing of H2O (Newman et al., 1988; Blundy et al., 2010) would likely be restricted to portions of magma adjacent to permeable pathways or ascending gas slugs, and would not affect the magma as a whole. In any of these processes, the fast diffusion of H2O in basaltic melts (Zhang and Ni, 2010) is an important consideration; it means both that heterogeneity can develop very quickly in response to spatial enrichments (e.g., around growing crystals) or depletions (e.g. around growing bubbles), but also that heterogeneities will be quickly homogenized. Within the context of the current work, the most important feature of the data is that H2O, from a bulk perspective, is higher in the matrix glasses than in glass inclusions in phenocrysts at all degrees of fractionation. There is thus no evidence for significant exsolution of H2O from the bulk melt between storage and eruption. 4.5. Eruption style

Fig. 7. Considerations for degassing of H2O and CO2. A: CO2 versus H2O, with isobars and constant vapor composition isopleths calculated with VolatileCalc (Ford et al., 1983; Newman and Lowenstern, 2002). B: H2O versus MgO, with contours of 25% olivine fractionation in 5% increments, from inclusion with variable starting H2O contents. See Section 4.4 for discussion of H2O variability.

To determine submarine eruption styles from deposit and sample characteristics is challenging, because there is a very small catalog to which each type of data can be compared. None of the types of data we present in this work independently proves what the eruption(s) in question would have been like. It is only in the combination of all available data types that we can begin to make inferences about the conditions of magma ascent, degassing, and eruption that produced this deposit and its components. The deposit examined in this study is significantly different from others examined during the same dive series, which were determined to result from “Poseidic” eruptions driven in part by strongly volatilecoupled degassing of CO2 and H2O, during rapid ascent to their vents (Schipper et al., 2010b,c). For this deposit, in contrast, textural features and diffusive re-equilibration of inclusions in high-Mg olivines indicate that the ejected magma resided for significant periods of time

30

C.I. Schipper et al. / Journal of Volcanology and Geothermal Research 207 (2011) 16–32

in the shallow subsurface. Furthermore, exsolution of magmatic H2O does not appear to have played any role in explosive ejection of the magma from the vent. Thus, the question of “how and why did this magma erupt explosively” still remains. We consider the most likely scenario to involve an eruption style sharing with subaerial Strombolian eruptions the ejection of magma driven by passage of large slugs of deep-exsolved magmatic CO2. This interpretation, although invoked by many authors to explain abyssal explosive activity (Chadwick et al., 2008; Clague et al., 2008; Sohn et al., 2008), is not directly supported by this dataset. The CO2 concentrations in inclusions are low, and require that there was an additional CO2 supply to produce the decoupled slugs. If this interpretation is correct, there must be a process of CO2 exsolution and accumulation from deeper in the magmatic system, before the inclusions were entrapped.

that although the majority of magmatic gas discharged at Stromboli volcano is characterized by high H2O/CO2 ratios, and is emitted during passive outgassing, the gasses in eruption-driving slugs have considerably lower H2O/CO2 ratios, and thus are sourced much deeper in the volcanic plumbing system. Similarly at Kilauea volcano, Edmonds and Gerlach (2007) demonstrated that deep-derived CO2-rich slugs can rise through the conduit to drive gas-pistoning events. Also, at Kilauea, the passage of these slugs does not deplete the surrounding melt in H2O, permitting (at atmospheric pressure) H2O-driven fountaining events to be interspersed with gas pistoning (Edmonds and Gerlach, 2007). Similarly here, the eruption appears to have been driven by deepderived slugs of CO2, which passed upward through overlying magma without stripping H2O during passage.

5. Discussion

We have used deposit characteristics, clast textures, geochemistry, and volatile concentrations to infer the style of submarine explosive eruption that produced a thick pyroclastic sequence at ~ 1100 mbsl on Lō`ihi Seamount, Hawaii. We find that deposit characteristics and textural data point to an eruption characterized by the ejection of magma that had significant time to stagnate, cool, and mature texturally in the shallow subsurface before being ejected. This inference broadly applies to all stratigraphic levels of the deposit, leading us to conclude that the eruption(s) were pulsatory, and likely most similar to classical subaerial Strombolian activity. Dissolved volatile data shows that magmatic water did not significantly exsolve from the magma that fed this eruption sequence, and thus did not drive the eruption(s). Theoretical considerations suggest that the eruptions were instead driven by buoyant slugs of magmatic CO2, exsolved from a larger volume of magma than was erupted. Our conclusions are consistent with many recent interpretations of submarine explosive eruption dynamics. We recognize, however, that the data in this case only indirectly shows that CO2-driven Strombolian eruptions are possible at Lō`ihi. Even with abundant data of diverse types, we run up against the fact that the role of CO2 from unsampled reservoirs is difficult, for any eruption lacking gas flux measurements, to directly assess analytically. Because the full range of eruption styles that are possible in the deep ocean is not yet known, it is only through the concurrent consideration of many different types of data that inferences about submarine eruption styles can be made. We stress that despite significant recent advances in the study of submarine volcanism, our understanding of submarine eruption styles is still “leagues” behind our understanding of subaerial events.

5.1. Estimating the CO2 budget required for Strombolian eruption Recent work suggests that some MORB magmas may include melt much richer in CO2 than generally thought (Helo et al., 2011); however, the classic utility of the submarine CO2-Strombolian model lies in the fact that it allows the typical “problem” of low parental volatile contents in explosively erupted OIB and MORB magmas to be circumvented, by having the eruption-driving volatiles accumulate from larger magma bodies than are actually erupted. In the following, we make a rough estimate of the volume of magma required to supply CO2 in sufficient quantity to produce the ~ 22 m thick pyroclastic sequence examined in this work. We consider the deposit to have originally been a cone, with height in the range of 25 to 50 m, and angle of repose for loosely consolidated lapilli of 25°. This yields an original deposit volume of 8 × 10 4 to 6 × 10 5 m 3. Using average lapilli density of ~ 1600 kg/m 3 (estimated as 70% lapilli with average density of 1800 kg/m 3 plus interstitial water) this translates to 10 8 to 10 9 kg of dense rock equivalent magma. The gas/magma ratio of discharges during subaerial Strombolian eruptions is both high and variable. To bracket possible CO2 budgets required to emplace the deposit, we use the range of gas:magma mass ratios determined for small (15.8:1) and large (2.4:1), lapilliproducing, eruptive bursts at Stromboli volcano, as estimated by Chouet et al. (1974). Taking all gas as CO2, we then estimate total CO2slug gas masses as 10 8 (25 m cone, low gas:magma mass ratio) to 10 10 kg (50 m cone, high gas:magma mass ratio). Dixon and Clague (2001) estimated tholeiitic parental CO2 concentrations at Lō`ihi to be 1.5 wt.%. The ~0.04 wt.% that must be retained to account for the highest concentration measured in the inclusions in high-Mg phenocrysts is insignificant at the precision of these calculations. The total slug CO2 could be derived by complete CO2 exsolution from 10 10 to 10 12 kg, or ~ 10 7 to 10 8 m 3 of dense magma (density of 2900 kg/m 3). For visualization purposes, this is equivalent to a spherical magma body with diameter of ~250 to 850 m. This rough calculation suggests that only a modest volume of magma, releasing all of its CO2 over the course of the eruption, but from below the storage zone where melt was trapped to form glass inclusions, would be needed to provide the CO2 required to produce this deposit through Strombolian-like activity.

6. Conclusions

Acknowledgments This work was partially funded by the ERC grant 202844 under the European FP7. Dives were funded by the National Oceanic and Atmospheric Administration (NOAA) Hawai`i Undersea Research Laboratory (HURL). JDLW acknowledges support from the FRST contract C05X0804 via subcontract to GNS science. The authors thank L. Danyushevsky and H. Wright for thorough and thoughtful reviews that helped to significantly refine the original manuscript. CIS especially thanks R.B. and M. Stewart for open access to the Massey University FTIR laboratories and for accommodating his perigrinations. References

5.2. Excess volatiles and shallow vs. deep degassing regimes “Excess degassing,” where the volcanic gas emissions at the surface of a volcano appear to be derived from excessively large volumes of magma at depth, is a common feature of many volcanoes (e.g., Shinohara, 2008). In specific case studies, distinct shallow and deep degassing regimes have been documented. Burton et al. (2007) showed

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