What Makes Hydromagmatic Eruptions Violent? Some Insights from the Keanaka¯Ko’I Ash, Kı¯Lauea Volcano, Hawai’I$ Larry G

Journal of Volcanology and Geothermal Research 137 (2004) 15–31 www.elsevier.com/locate/jvolgeores

What makes hydromagmatic eruptions violent? Some insights from the Keanaka¯ko’i Ash, Kı¯lauea Volcano, Hawai’i$ Larry G. Mastina,*, Robert L. Christiansenb, Carl Thornbera, Jacob Lowensternb, Melvin Beeson1

a U.S. Geological Survey, Cascades Volcano Observatory, 1300 SE Cardinal Court, Building 10, Suite 100, Vancouver, WA 98683, USA b U.S. Geological Survey, MS 910, 345 Middlefield Road, Menlo Park, CA 94025, USA

Abstract

Volcanic eruptions at the summit of Kı¯lauea volcano, Hawai’i, are of two dramatically contrasting types: (1) benign lava flows and lava fountains; and (2) violent, mostly prehistoric eruptions that dispersed tephra over hundreds of square kilometers. The violence of the latter eruptions has been attributed to mixing of water and magma within a wet summit caldera; however, magma injection into water at other volcanoes does not consistently produce widespread tephras. To identify other factors that may have contributed to the violence of these eruptions, we sampled tephra from the Keanaka¯ko’i Ash, the most recent large hydromagmatic deposit, and measured vesicularity, bubble-number density and dissolved volatile content of juvenile matrix glass to constrain magma ascent rate and degree of degassing at the time of quenching. Bubble-number densities (9 104 – 1 107 cm 3) of tephra fragments exceed those of most historically erupted Kı¯lauean tephras (3 103 –1.8 105 cm 3), and suggest exceptionally high magma effusion rates. Dissolved sulfur (average = 330 ppm) and water (0.15–0.45 wt.%) concentrations exceed equilibrium-saturation values at 1 atm pressure (100–150 ppm and f 0.09%, respectively), suggesting that clasts quenched before equilibrating to atmospheric pressure. We interpret these results to suggest rapid magma injection into a wet crater, perhaps similar to continuous-uprush jets at Surtsey. Estimates of Reynolds number suggest that the erupting magma was turbulent and would have mixed with surrounding water in vortices ranging downward in size to centimeters. Such fine-scale mixing would have ensured rapid heat exchange and extensive magma fragmentation, maximizing the violence of these eruptions. Published by Elsevier B.V.

Keywords: Kı¯lauea; Phreatomagmatism; explosive eruptions; vesicular texture; turbulence

1. Introduction hydromagmatic tephra deposits erupted from the summit of Kı¯lauea volcano, Hawai’i, since the late Pleis- The Keanaka¯ko’i Ash (Fig. 1) is the most recent, tocene. Originally thought to have been produced best exposed and most thoroughly studied of several during a single eruption around A.D. 1790 (Dana, 1888), the number of eruptions represented by the $ Supplementary data associated with this article can be found, deposit has been repeatedly revised (Stone, 1926; in the online version, at doi: 10.1016/j.jvolgeores.2004.05.015. Powers, 1948; Christiansen, 1979; McPhie et al., * Corresponding author. Tel.: +1-360-696-7518; fax: +1-360- 1990). It is currently thought to have resulted from 993-8980. E-mail address: [email protected] (L.G. Mastin). at least a few eruptions over a period of perhaps a few 1 formerly U.S. Geological Survey, MS 910, 345 Middlefield centuries that ended around 1790 AD (Swanson et al., Road, Menlo Park, CA 94025, USA. 1998; D. Swanson, written comm., 2001). The most

0377-0273/$ - see front matter. Published by Elsevier B.V. doi:10.1016/j.jvolgeores.2004.05.015 16 L.G. Mastin et al. / Journal of Volcanology and Geothermal Research 137 (2004) 15–31

(loc. 1, Fig. 1a; Easton, 1987; Clague et al., 1995a) and may have originally covered thousands of square kilometers. The violence of eruptions that produced these deposits is generally attributed to interaction of magma with water, as inferred from their fine grain size, wide range in vesicularity, cross-bedding, accretionary lapilli, and lithic components in the ejecta (Swanson and Christiansen, 1973; Decker and Christiansen, 1984; McPhie et al., 1990; Dzurisin et al., 1995; Clague et al., 1995a). Mastin (1997) hypothesized that water influx was caused by subsidence of the caldera floor below the water table (currently at f 490 m depth; Hurwitz et al., 2003). Since the early 1800s, the caldera floor has fluctuated in elevation by hundreds of meters and has come within less than f 100 m of the water table on at least three occasions (Mastin, 1997).

1.1. Explosivity

Subsidence of the eruptive vent below the water table would allow water to accumulate within the crater and would guarantee water-magma mixing during summit eruptions; It would not, however, guarantee that such eruptions be large and explosive. Along Kı¯lauea’s coast, lava commonly enters water without extensive fragmentation or dispersal of tephra (Mattox and Mangan, 1997). Non-violent effusion of lava into water was also observed at Mauna Loa in 1950 (Finch and Macdonald, 1953), at Soufrie`re of St. Vincent in 1972 (Shepherd and Sigurdsson, 1982) and Kick-’em Jenny in the West Indies in 1974–1978 (Devine and Sigurdsson, 1995), to name a few examples. We surmise that violence (or non-violence) of these Fig. 1. (a) Index map of Kı¯lauea showing the areal extent of the and other hydromagmatic eruptions is affected not Keanaka¯ko’i Ash (heavy dotted line; from R. Christiansen, only by the injection of magma into water but also the unpublished data), (b) close-up of summit showing sample circumstances under which they mixed. At least, two locations (numbered). Coordinates of numbered locations are given mechanisms for explosive magma-water mixing have in Table 2. been described during central-vent eruptions: (1) drawdown of magma in a conduit below the water extensive Keanaka¯ko’i units that likely resulted from table, followed by influx of water, and (2) jetting of a single eruption (e.g. unit 2 of McPhie et al., 1990) magma through surface water or through a wet crater cover a few hundred square kilometers (R. Christian- into which water is entering through the porous walls. sen, unpublished data, 2002), similar in area to the Mechanism (1) was noted at Ukinrek Maars, Alaska, deposits of moderately large historical basaltic hydro- when a hydromagmatic explosion, observed from the magmatic events (e.g. Taal, 1965; Moore et al., 1966). air on 3 April, 1977, was preceded a few minutes Late Pleistocene tephras at Kı¯lauea, however, are earlier by the draining of a lava lake and collapse of meters thick in exposures 10 km south of the caldera the crater walls (Kienle et al., 1980). The sequence L.G. Mastin et al. / Journal of Volcanology and Geothermal Research 137 (2004) 15–31 17 was similar to well-described non-juvenile steam In this study, we examine vesiculation textures and explosions at Ha¯lema’uma’u Crater in 1924, which dissolved-volatile concentrations in the lower unit of took place after a long-lived lava lake drained and hot the Keanaka¯ko’i Ash to constrain magma-water con- conduit walls collapsed (Stearns, 1925; Decker and ditions at the time of quenching. The Keanaka¯ko’i Christiansen, 1984). Ash is especially suited to this analysis because the The jetting of magma through water was described solubility and pre-eruptive volatile content of during continuous-uprush phases of Surtsey Volcano, Kı¯lauean basalt are well constrained (e.g. Gerlach, Iceland (Thorarinsson, 1964; Moore, 1985). These 1986; Dixon et al., 1991, 1995; Wallace and Ander- eruptions were characterized by cylindrical columns son, 1998) and vesiculation textures have been exten- 100–250 m in diameter at their base (Moore, 1985), sively characterized (Mangan et al., 1993; Cashman et that blasted fragmental debris upward at velocities of al., 1994; Mangan and Cashman, 1996). f 100 m/s and persisted for hours. They occurred when seawater access to the vent was ‘‘partially or wholly blocked2’’ (Thorarinsson, 1964),andwere 2. The Keanaka¯ko’i deposit generally preceded by intermittent hydromagmatic explosions that repeated with increasing frequency The Keanaka¯ko’i Ash consists of three major units until jetting became continuous. (The cause of inter- and several minor ones (Decker and Christiansen, mittent explosions has been the subject of disagree- 1984; McPhie et al., 1990). The major units include ment (Kokelaar, 1983) and may be a variant on (1) well-bedded ash and lapilli of inferred hydro- mechanism (1)). Water was apparently supplied to magmatic origin containing primarily fresh, juvenile the jet by seepage over and through the tephra pile sideromelane glass (the ‘‘lower, juvenile-rich beds’’ of that lined the vent crater. Continuous-uprush phases McPhie et al., 1990); (2) massive and cross-bedded produced much more tephra and dispersed it more ash and lapilli, also of inferred hydromagmatic origin, widely than Surtsey’s intermittent explosions (Thor- containing both juvenile glass and lithic fragments arinsson, 1964, p. 44). (the ‘‘middle, mixed beds’’); and (3) block fall and In principle, it should be possible to distinguish cross-bedded surge deposits composed almost exclu- mechanisms (1) and (2) using vesiculation textures sively of lithic debris (the ‘‘upper lithic beds’’). These and dissolved-volatile contents of juvenile clasts. units are named II, III and V, respectively, in the Under mechanism (1), if magma existed within a nomenclature of Decker and Christiansen and 0–4, subaerial lava lake prior to the explosions, it should 6–10, and 11–16 in the nomenclature of McPhie et contain dissolved volatiles in equilibrium with 1 atm al. (1990). The units are underlain by a massive pressure. Magma quenched during withdrawal should reticulite bed (unit I of Decker and Christiansen), contain low vesicularity and low bubble-number den- which in turn is locally underlain by a gray vitric- sities that typify low ascent rates or stagnation (Cash- lithic ash (D. Swanson, written comm. 2002). The man and Mangan, 1994). Under mechanism (2), three units are overlain on the northwest and south- juvenile clasts may contain dissolved volatiles equal west sides of the caldera by a pumice layer a few tens to or exceeding equilibrium values at 1 atmosphere of centimeters thick (unit VI of Decker and Christian- depending on the pressure and degree of equilibration sen; the ‘‘Golden Pumice’’ of Sharp et al., 1987). of the clasts at the time of quenching. Magma that Southwest of the caldera, along the contact between quenched during rapid ascent should exhibit abun- units III and V of Decker and Christiansen are dant, fine vesicles with high bubble-number densities scattered mats of Pele’s hair and tears (unit IV of similar to lava-fountain tephras in subaerial environ- Decker and Christiansen), inferred to have erupted ments (Cashman and Mangan, 1994; Mangan and from a fissure whose lavas are intercalated with the Cashman, 1996). Keanaka¯ko’i Ash (unit 1790f of Neal and Lockwood, in press). In addition, at the base of the middle, mixed 2 Water could not truly have been ‘‘wholly’’ blocked as beds (according to McPhie et al., 1990), or in the Thorarinsson describes, since he also characterizes these events as upper part of unit II (according to Decker and Chris- hydromagmatic. tiansen, 1984), is a scoria bed up to a few tens of 18 L.G. Mastin et al. / Journal of Volcanology and Geothermal Research 137 (2004) 15–31 centimeters thick (unit 6 of McPhie et al.) whose axis ash beds4 with sparse lapilli horizons (unit IIC; Figs. 2 of dispersal extends to the southeast, in contrast to the and 3). Conformable bedding within subunits IIA and southwestward dispersal of most units in the Keana- IIB suggests that each was deposited during a single ka¯ko’i deposit. eruptive pulse or rapid series of pulses with no significant pauses. The contacts at the top of units 2.1. Evidence for depositional breaks IIA and IIB are sharp and truncate underlying beds at a low angle (Fig. 3b), suggesting some centimeters of In one of the first written discussions of the erosion, perhaps by surges or wind, between deposi- deposit, Dana (1888) assumed it resulted from a single tional pulses. The upper contact of unit IIA lacks eruptionin1790ADthatwasknownfromoral recognizable evidence of water erosion, suggesting accounts (Ellis, 1827; Dibble, 1843). Early 20th the pause was short and without rain. The upper century authors attributed only the upper lithic unit contact of unit IIB exhibits scarce water-erosion to the eruption in 1790 and lower units to earlier rivulets and locally overlies cross-bedded and appar- eruptions (Hitchcock, 1909; Powers, 1916; Stone, ently wind-reworked debris from underlying units 1926; Stearns and Clark, 1930; Wentworth, 1938; (McPhie et al., 1990), suggesting a longer pause of Finch, 1947; Powers, 1948). Later, Christiansen unknown duration. (1979) and Decker and Christiansen (1984) proposed McPhie et al. (1990, p. 351) speculated that that all members between the basal reticulite and the juvenile-rich beds resulted from repeated explosions Golden Pumice erupted during a single sequence in or that began with vesiculation-driven magma ascent around 1790 AD. McPhie et al. (1990) concurred but and ended with drainback of degassed magma. noted three horizons of erosion and reworking that Mastin (1997), used quantitative vesicularity and suggested pauses of unknown and possibly long bubble-number density measurements to argue that duration. Most recently, Swanson et al. (1998) cited some horizons in units IIA2 and IIB2 involved organic horizons, water-erosion features, archaeolog- sustained, high magma discharge. Whether other ical artifacts and carbon dates as evidence for at least beds involved sustained high discharge was not few large eruptive sequences, beginning shortly after investigated. appearance of the present-day caldera in the late 15th century and ending around 1790 AD (D. Swanson, written comm., 2001). Horizons that show the stron- 3. Sample collection and analysis gest evidence for non-depositional periods lie immediately above the basal reticulite, slightly below and We collected samples by trowel from the outcrop above unit 6 of McPhie et al., and at the base of the (Tables 1 and 2) and measured the volume percent of upper lithic beds (McPhie et al., 1990). fresh glass, altered glass, olivine and lithic fragments by point-counting grains mounted in thin sections 2.2. Field locations (f 400 grains per section; Table 3). We analyzed dissolved-volatile concentrations using the electron Our study concentrates on unit II of Decker and microprobe for sulfur and Fourier transform infrared Christiansen, using samples collected in Sand Wash (FTIR) spectrometry for water and CO2. Juvenile (locations 61 and 84, Fig. 1b) and at a fissure a glass in the deposit ranges in vesicularity from 0% kilometer north of Sand Wash (location 62, Fig. 1b). to >80%, requiring us to prepare samples for FTIR At these locations, unit II contains three subunits: two analysis using two methods: (1) slightly-to-moderate- coarsening-upward sequences3 of well-bedded ash ly vesicular clasts, 1–2 mm in diameter were embed- and lapilli (units IIA and IIB) and a sequence of fine ded in a 2.5-cm diameter wafer of epoxy that was then ground and polished, flipped, remounted on a glass

3 slide, then ground to a thickness of 70–165 Am, and Unpublished grain-size analyses give mdf = 0.64 to 3.5, rf = 1.22–1.86 for ash in the lower half of these sequences and 4 mdf = 0.38–0.62, rf = 1.51–2.10 for ash and lapilli in the upper mdf = 0.67–3.71, rf = 1.31–2.53 based on unpublished half. grain-size analyses. L.G. Mastin et al. / Journal of Volcanology and Geothermal Research 137 (2004) 15–31 19 polished on the other side; (2) highly vesicular lapilli spectrometer with an attached SpectraTechR Analyti- were ground with a mortar and pestle and flakes f 1 cal-IR microscope that utilizes a liquid-N2-cooled mm in diameter were handpicked, mounted in epoxy MCT-A detector. We collected 512 scans per analysis and doubly polished as described above. on each sample and on backgrounds following each sample analysis. The measured absorbance of infrared 3.1. Dissolved volatile analysis radiation is directly proportional to the concentration of H2O in the glass, adjusted for sample density and Before removing the polished grain mounts from thickness. Weight percent H2O was calculated accord- their glass slides, each fragment was analyzed for ing to Eq. (15) of Ihinger et al. (1994) for the OH sulfur by electron microprobe using the JEOL JXA- stretch at 3570 cm 1, assuming a glass density of 8900 Electron Probe Microanalyzer at the USGS 2700 F 100 g l 1, an extinction coefficient (e)of Menlo Park Laboratory. These samples were a subset 63 F 3 l mol 1 cm 1 and sample thickness as mea- of more than 23 thin sections containing tephra frag- sured by a MitutoyoR Digital Micrometer. Uncertainty ments analyzed for sulfur during 6 microprobe ses- for e, sample thickness and density are all about 5%, sions of 1–2 days each. For reference, we also so that propagated errors are all between 10% and analyzed clasts from non-hydromagmatic units I, IV, 15% relative. VI and the 1959 Kı¯lauea Iki tephra (Table 1). Values We detected no molecular H2O (peak at 1630 1 1 of S and K2OinTable 4 and the major-element data cm )orCO3 (1515 and 1430 cm ); we analyzed (see supplementary data in the online version) repre- two clasts in sample 517B twice, collecting 1024 sent averages of two to four analyses per clast (exact scans during the second analysis, to verify the absence numbers are listed online). During each session, we of these peaks. The detection limit of dissolved used an acceleration voltage of 15 keV, a beam current CO3 is f 50 ppm, suggesting that the melt degassed of 25 nA, a beam diameter of 10–15 Am, a barite to less than f 5 MPa (using solubility relations of standard for sulfur, and a peak counting interval of 80 Dixon et al., 1995), or about 600 m depth (assuming a s. Results for elements other than sulfur, and standards pressure gradient of 25 MPa/km). The absence of a used, are provided in the online supplement. Of 343 molecular water peak is consistent with the experi- analyses, 330 gave totals between 97% and 101%. mentally measured speciation of dissolved water in The precision of sulfur concentrations (ppm) is basaltic glass at this concentration (Dixon et al., estimated from counting statistics and reproducibility 1995). Meteoric water adsorbed into the clasts at of USNM basaltic glass standard VG-2 (Jarosewich et magmatic temperature could also speciate of OH, al., 1979), which is similar in composition to the but water absorbed at lower temperature would likely analyzed glasses. The standard deviation in sulfur remain in molecular form (P. Wallace and J. Dixon, values of VG-2 glass measured during individual written comm., 2001). If water were absorbed into the microprobe sessions is 57 ppm (the half-width of clasts during or after the eruption, we would expect it the error bar in Fig. 2); however, mean VG-2 sulfur to be absorbed over a range of temperatures and to analyses for a given session range from 1080 to 1350 leave at least some water in molecular form. The ppm, suggesting some drift between sessions. Sulfur complete absence of molecular H2O in these clasts calibration numbers were therefore adjusted so that therefore suggests to us that all dissolved water is the average VG-2 analysis in each session equals magmatic. 1340 ppm, the value obtained by Dixon et al. (1991). After microprobe analysis, the grain mounts were 3.2. Vesicularity and bubble-number density removed from their glass slides and bubble-and crystal- free regions of glass were targeted for FTIR analysis To obtain bubble-number densities, all of the using apertures above and below the sample. Measure- dense, hand-picked clasts analyzed by FTIR were ments were performed using a Nicolet5 Magna 750 photographed in thin section under reflected light. Printed photomicrographs were overlain with trans- 5 Use of trade names in this document does not imply parencies and the outlines of the clasts, and of internal endorsement by the USGS. bubbles, were traced in ink. The inked transparencies 20 L.G. Mastin et al. / Journal of Volcanology and Geothermal Research 137 (2004) 15–31 L.G. Mastin et al. / Journal of Volcanology and Geothermal Research 137 (2004) 15–31 21 were then digitally scanned and the clast diameter, tages of 0.5, 1, 1.5, 3, 5, 7, 10, 15, 10, 25, 30, 35, 40, area, and the number and size distribution of bubbles 45 and 50. We consider those intervals to indicate the were computed using the software Scion ImageR. The resolution of the estimation technique. Results of number of bubbles per clast ranged from zero to these measurements, as well as microprobe data and several dozen; too few to derive meaningful size- lapilli clast density, are posted online. distribution statistics but enough to estimate the order-of-magnitude bubble-number density. The highly vesicular clasts were destroyed when 4. Results ground by mortar and pestle and could not be analyzed for bubble-number density. Prior to grinding, Sulfur values in matrix glass from unit II (Fig. 2) however, their vesicularity was measured by weighing average 330 ppm with a standard deviation of 80 the samples in and out of water following the method ppm—significantly above the 100–150 ppm typically of Houghton and Wilson (1989). We calculated clast measured for Kı¯lauean glass that has equilibrated to 3 density (q, kg/m ) using the formula q = qw(wa/ atmospheric pressure (Swanson and Fabbi, 1973; (wa ww)), where qw is the density of water (1000 Mangan and Cashman, 1996; Thornber, 2001). Glass 3 kg/m ), and wa and ww are the measured weights in air in non-hydromagmatic units I, IV, VI and the Kı¯lauea and in water, respectively. The volume-fraction gas Iki tephra, all contain dissolved sulfur values consis- (vg)—i.e. the fraction of the clast volume composed of tent with equilibrium degassing to 1 atm pressure. bubbles—was calculated using the formula vg =1 q/ Dissolved sulfur in glass inclusions in olivine ranges qg, where qg is the density of basaltic glass, taken as from about 250 to 1530 ppm, with most around 2700 kg/m3. We calculated the equivalent radius (r)of 1000–1200 ppm (Fig. 2). Considering the middle of each lapillus (assuming a spherical shape) from the this range to represent undegassed inclusions and 100 1/3 formula r =(3wa/(4pq)) . Densities are generally ppm to be totally degassed, most hydromagmatic repeatable to F 10%. glass fragments are about 70–80% degassed in sulfur. During the course of our analysis, we observed that The concentration of dissolved water in juvenile vesicularity varied systematically from one strati- glass also measurably exceeds equilibrium saturation graphic unit to the next and from ash to lapilli-sized under atmospheric conditions (Fig. 2).Sulfurand clasts within a given unit. To quantify these variations, water values (Table 4) are moderately correlated 2 we compare vg from 3000 clast density measurements, (r = 0.69) supporting the view that the water is published earlier (Mastin, 1997), with vg from f 3000 magmatic in origin. Measured H2O concentrations ash-sized grains, estimated visually in thin section by correspond to quench pressures (under equilibrium comparing the abundance of bubbles to diagrams of H2O saturation) of f 0.3–2.3 MPa (Fig. 2), equiva- modal percentage (Compton, 1985, Appendix 3). lent to 10–100 m depth at lithostatic pressure (as- Diagrams in Compton (1985) are given for percen- suming a pressure gradient of 25 MPa/km), or f 25–

Fig. 2. (left) Keanaka¯ko’i Ash, showing stratigraphic units exposed at Sand Wash plus units I and VI of Decker and Christiansen (1984); (left center) total dissolved sulfur in glass from microprobe measurements; (right center) total dissolved water in matrix glass, measured by FTIR; (far right) components of these units from point counts (Table 3). To the left of the stratigraphic section are unit names according to McPhie et al.

(1990) (left) and Decker and Christiansen (1984) (right). Tick marks at top of the H2O plot indicate the equilibrium dissolved water in MORB glass at saturation, calculated using a best-fit exponential curve through six experimental measurements from Dixon et al. (1995), using samples that contained no dissolved CO2 and were tested at pressures < 200 MPa. One extra data point was added to this data set: the well-known equilibrium concentration (0.09 wt.%) at p = 0.1013 MPa (e.g. Wallace and Anderson, 1998; Mangan et al., 1993; Cashman et al., 1994). The .515 2 best-fit curve has the form H2O = 0.295p (r = 0.9998), where H2O is in weight percent and p is pressure in megapascals (standard error = 0.003 wt.% H2Oatp = 0.1013 MPa, 0.029 wt.% at 2.5 MPa). This section contains one revision to the stratigraphy of Decker and Christiansen (1984): unit 6 of McPhie et al., which is thickest near Keanaka¯ko’i Crater, was thought by Decker and Christiansen (1984) to correlate with unit IIC2 at Sand Wash. However, the isopach map of unit 6 in McPhie et al. suggests that it does not extend this far west. Moreover, these units differ petrologically (IIC2 at Sand Wash contains pumice and lithics, whereas unit 6 at Keanaka¯ko’i Crater contains juvenile scoria). Scattered fragments of scoriaceous debris at the base of unit IIC1 at Sand Wash (hollow diamonds in the sulfur plot) may be remnants of unit 6. Further refinements to this stratigraphy are underway. No unit names corresponding to unit IIC of Decker and Christiansen (1984) are shown in the stratigraphy of McPhie et al. because no units of McPhie et al. appear to correlate with them. 22 L.G. Mastin et al. / Journal of Volcanology and Geothermal Research 137 (2004) 15–31

Fig. 3. Photos of the Keanaka¯ko’i Ash at Sand Wash (location 61, Fig. 1b) showing stratigraphy and locations of some of the samples collected there. (a) Upper part of section and (b) lower part of section. Three-digit numerals on photos indicate sample numbers. Stratigraphic names of Decker and Christiansen are given in white lettering on the left side of photos; names from McPhie et al. are on the right side.

250 m at hydrostatic pressure (100 MPa/km). The K2O (0.44 F 0.05 wt.%) in clasts from hydromag- amount of H2O lost to degassing can be assessed from matic units imply pre-eruptive water content of the ratio of H2O/K2O(Table 4, Fig. 4). Both of these 0.57 F 0.06 wt.%. Measured H2O/K2O ratios (0.39– components are incompatible in a crystallizing 1.02) suggest that the melt had lost 25–80% of its Kı¯lauean melt and, in the absence of degassing, H2O by the time it quenched. their ratio should remain constant at about 1.3 (Dixon Fig. 5 shows an inverse relationship between et al., 1991; Wallace and Anderson, 1998). Average dissolved water and vg, which is consistent with a L.G. Mastin et al. / Journal of Volcanology and Geothermal Research 137 (2004) 15–31 23

Table 1 Table 2 Summary of samples collected and analyzed for this paper Sample locations shown in Fig. 1 Sample Unit Location Type of Number Easting Northing number D and C McPhie number analysis 1 259221 2146965 426 Iki Iki 21 mp 21 261979 2146795 3HK215Q VI 1 mp 61 259261 2145550 62 259240 2146443 602 IV 84 mp 423 IIC4 62 mp, pc 67 259715 2148958 525 IIC3 61 d, ves 84 259237 2145737 422 IIC3 62 mp, pc Easting and Northing are given in UTM Zone 5 coordinates. 407b IIC3 61 pc 523 IIC3 61 d, ves 521 IIC3 61 d, ves below the line that represents closed-system gas 421 IIC2 62 mp exsolution with total water (dissolved plus exsolved) 407a IIC2 61 mp, pc equal to 0.55 wt.%. Large, highly vesicular lapilli 520 IIC2 61 mp, FTIR, bnd, d appear to have suffered more gas loss than smaller, 519 IIC1 61 d, ves less vesicular ash. 518 IIC1 61 ves 420 IIC1 62 pc These data clearly show that hydromagmatic glass 517 IIC1 61 mp, FTIR, bnd did not fully degas in sulfur or water. In contrast, non- 516 IIC1 61 d, ves hydromagmatic clasts are completely degassed in 419 IIC1 62 pc sulfur. We did not measure dissolved water in non- 514 6? 61 mp, d hydromagmatic tephra, but tephra from phase 1 of the 513 IIB2 4 61 d, ves 418 IIB2 4 62 d Kı¯lauea Iki eruption, analyzed by Wallace and Ander- 417 IIB2 4 62 mp, pc son (1998, fig. 1), are in equilibrium with 1 atm 512 IIB2 4 61 d, ves pressure (phase 1 tephra erupted before degassed lava 406 IIB2 4 61 d 405 IIB2 4 61 mp, pc 511 IIB2 4 61 mp, FTIR, bnd, d, ves Table 3 416 IIB1 3 62 pc Summary of point count results 509 IIB1 3 61 d Sample Unit Components n 508 IIB1 3 61 mp, FTIR, bnd 403 IIB1 3 61 mp, pc D and C McPhie Altered Fresh Lithic Crystal 503 IIB1 3 61 mp, FTIR, bnd glass glass 413 IIA2 2 62 mp, pc 423 IIC4 17% 63% 12% 8% 400a 414 IIA2 2 62 d 422 IIC3 14% 80% 0% 6% 400 502 IIA2 2 61 mp, FTIR, bnd, d 407b, 1f IIC3 18% 76% 4% 2% 400 402 IIA2 2 61 d 407a, 1f IIC2 49% 43% 6% 2% 400 412 IIA2 2 62 mp, d 420 IIC1 11% 86% 0% 3% 401 401 IIA2 2 61 mp (top) 501 IIA2 2 61 d 419 IIC1 73% 24% 0% 3% 400 411 IIA1 1 62 mp, pc (base) 526 IIA1 1 61 ves 417 IIB2 4 6% 78% 15% 2% 400 409 IIA1 1 62 mp, FTIR, bnd (top) 404 IIA1 1 61 pc 405, 1f IIB2 4 11% 81% 5% 3% 390 432 I 67 mp, d 416 IIB1 3 9% 91% 0% 1% 400 Abbreviations in the right column include clast density measure- (top) ments (‘‘d’’), vesicularity measurements from thin section (‘‘ves’’), 403, 1f IIB1 3 3% 96% 1% 0% 404 microprobe glass analyses (‘‘mp’’), FTIR analyses (‘‘FTIR’’), 413 IIA2 2 8% 73% 9% 10% 400 bubble-number density (‘‘bnd’’) and point counts (‘‘pc’’). (top) 411 IIA1 1 12% 87% 1% 1% 400 (top) scenario in which clasts were quenched during active 404, 1f IIA1 1 7% 86% 7% 1% 400 vesiculation. Dissolved H2O is lowest in the highly a The standard error for point counts of 400 grains is less than vesicular clasts of unit IIB2. Nearly all data lie 4% (Van der Plas and Tobi, 1965). 24 L.G. Mastin et al. / Journal of Volcanology and Geothermal Research 137 (2004) 15–31

Table 4 Summary of FTIR and bubble-number density measurements

Sample Unit Thickness vg Diameter Clast area Number of Number density S H2O K2OH2O/K2O (Am) (mm) (mm2) bubbles (cm 3) (wt.%) (wt.%) 520b IIC2 136 0.35 1.79 0.575 153 5.8E + 05 0.023 0.33 0.4585 0.7197 126 0.35 1.62 0.435 119 9.5E + 05 0.033 0.29 0.4427 0.6551 145 0.43 1.13 0.435 135 1.0E + 07 0.030 0.38 0.4383 0.8669 140 0.49 1.46 0.920 256 7.9E + 06 0.037 0.28 0.4387 0.6383 134 0.46 1.12 0.804 201 8.6E + 06 0.042 0.38 0.4400 0.8636 517b IIC1 152 0.39 1.63 1.536 161 2.1E + 06 0.042 0.37 0.4293 0.8618 132 0.42 1.65 1.755 136 1.0E + 06 0.042 0.33 0.4540 0.7269 125 0.44 1.21 1.027 53 1.1E + 06 0.033 0.36 0.4447 0.8096 136 0.43 1.4 1.157 69 1.0E + 06 0.042 0.45 0.4450 1.0112 128 0.42 1.35 1.027 105 1.9E + 06 0.035 0.29 0.4470 0.6488 140 0.42 1.31 0.934 55 8.3E + 05 0.034 0.34 0.4390 0.7745 147 0.44 1.23 0.048 0.41 0.4398 0.9323 511e IIB2 94 0.671 5.29 0.017 0.25 0.4002 0.6247 93 0.643 4.61 0.018 0.25 0.3958 0.6317 92 0.643 3.73 0.022 0.19 0.3903 0.4868 83 0.51 3.86 0.021 0.23 0.4052 0.5676 74 0.51 3.32 0.021 0.2 0.4053 0.4935 75 0.70 5.02 0.020 0.23 0.3977 0.5783 75 0.31 3.26 0.031 0.29 0.4103 0.7068 74 0.51 3.32 0.024 0.3 0.4021 0.7461 68 0.86 5.71 0.034 0.34 0.4165 0.8163 76 0.59 5.48 0.024 0.24 0.4168 0.5758 71 0.75 4.28 0.017 0.16 0.4101 0.3901 73 0.41 3.20 0.028 0.29 0.4009 0.7234 508f IIB1 93 0.67 0.49 0.1573 5 3.9E + 05 0.039 0.31 0.4273 0.7254 101 0.21 0.52 0.1591 3 9.5E + 04 0.034 0.33 0.4383 0.7529 90 0.66 0.56 0.1188 18 3.4E + 06 0.040 0.38 0.4283 0.8872 100 0.50 0.81 0.3902 13 3.2E + 05 0.041 0.32 0.4220 0.7583 90 0.32 0.66 0.2581 23 1.1E + 06 0.032 0.34 0.4227 0.8044 83 0.21 0.47 0.032 0.26 0.3950 0.6582 104 0.39 0.74 0.3237 18 7.2E + 05 0.027 0.36 0.4097 0.8788 100 0.52 0.41 0.048 0.33 0.4055 0.8138 502h IIA2 136 0.35 1.66 1.0158 200 5.5E + 06 0.041 0.3 0.4523 0.6632 143 0.44 1.26 1.0158 94 1.6E + 06 0.032 0.35 0.4270 0.8197 147 0.14 1.48 0.8732 79 1.1E + 06 0.027 0.42 0.4130 1.0169 148 0.24 1.04 0.6620 19 2.1E + 05 0.037 0.42 0.4373 0.9604 147 0.26 1.10 1.4885 36 1.4E + 05 0.047 0.4 0.4303 0.9295 409b IIA1 157 0.49 1.31 0.4768 35 9.2E + 05 0.041 0.34 0.4200 0.8095 167 0.31 1.71 1.2399 100 1.2E + 06 0.040 0.37 0.4303 0.8598 148 0.33 1.28 0.6712 139 4.5E + 06 0.041 0.37 0.4197 0.8817 152 0.46 1.16 0.5397 15 3.1E + 05 0.041 0.38 0.4050 0.9383 163 0.27 1.47 0.8240 30 2.8E + 05 0.035 0.4 0.4137 0.9670 pffiffiffiffiffiffiffiffi Unit names are those of Decker and Christiansen (1984). Clast diameter (D) is calculated as D ¼ 2 A=p, where A is clast area. Clasts in sample 511e were prepared by grinding lapilli in a mortar and pestle and handpicking fragments for analysis. Other clasts were prepared by handpicking small clasts from grab samples. Because of the difference in sample preparation, no number densities were measured for clasts in sample 511e. from lava ponds entered the vent). We therefore infer 4.1. Vesiculation textures that water, like sulfur, is elevated due to rapid quenching by external water. Whether these clasts quenched In general, pyroclasts of a given size are less at elevated pressure or at 1 atm before they could vesicular in units IIA1, IIB1, more vesicular in units degas to equilibrium is the subject of later discussion. IIA2 and IIB2. Within individual units, average vesic- L.G. Mastin et al. / Journal of Volcanology and Geothermal Research 137 (2004) 15–31 25

values exceed number densities obtained from vent samples of effusive eruptions (3000–5000 cm 3; Mangan et al., 1993) and most exceed values measured in lava-fountain tephra from Pu’uO’o (19,000– 180,000 cm 3; Mangan and Cashman, 1996). Unlike our Keanaka¯ko’i clasts, the lava-fountain tephra rose many seconds in a hot plume before cooling, raising the question of whether bubble coa- lescence and escape could account for their lower number densities. We do not think so. Such processes recognizably alter texture by increasing vesicularity in clasts interiors and reducing it on margins (Walker and Croasdale, 1972). Mangan and Cashman (1996) avoided this complication by selecting clasts that were texturally homogeneous. From plots of log bubble Fig. 4. Dissolved H2OversusK2O in matrix glass from number versus bubble diameter, they inferred that Keanaka¯ko’i Ash. Melt that has not lost H2O to degassing plots only about 13–40% of bubbles coalesced from two along the line H2O/K2O = 1.3, while completely degassed melt would plot along a line having zero slope and an intercept at or more smaller ones—not enough to account for the H2O = 0.09 wt.%. Intermediate degrees of degassing are represented < 100 lower number densities of lava-fountain by lines having intermediate slopes, intersecting the H2O/K2O = 1.3 clasts relative to Keanaka¯ko’i clasts. line at H2O = 0.09 wt.%. ularity increases with average grain size (Fig. 6). 5. Implications Although vg was estimated for the ash-and lapilli-sized clasts using two different methods, we do not think These high bubble-number densities of unit II that sample bias in one or both methods is responsible suggest rapid magma ascent and eruptive activity for the variation. Measurements made using both methods in the same size range (for example, in unit IIB2, 2–3 mm diameter clasts) yield the same range in vg. Moreover, estimates by visual inspection are similar to results from the more accurate image processing technique (triangles, units IIA2, IIB1 and IIC1). For clast diameters larger than a few tenths of a millimeter, lower values of vg do not appear to result simply from fragmentation at a scale smaller than the average bubble diameter. Many clasts of this size (e.g. Fig. 7) contain bubbles smaller than their own diameter. The lower average dissolved water concentration of the larger, more vesicular clasts (Fig. 5) and the incomplete degassing of all hydromagmatic clasts suggest that the smaller clasts were at an earlier stage of vesiculation. Bubble-number densities in unit II (Table 4, Fig. 8) Fig. 5. Weight percent dissolved water versus volume fraction gas in range from 9 104 to 1 107 cm 3 similar to those melt. Squares are measurements (symbol size is proportional to clast measured in unit IIB2 by Mastin (1997). Though the diameter). Lines give volume-fraction gas versus dissolved water for closed-system degassing for a total water content (dissolved plus earlier data were obtained only from highly vesicular exsolved) of (top to bottom) 0.55, 0.44, 0.33 and 0.22 wt.%, lapilli, the new data come from ash-sized clasts with respectively. The lines represent loss of 0%, 20%, 40% and 60% of vesicularity ranging down to about 20%. All these the original gas, respectively. 26 L.G. Mastin et al. / Journal of Volcanology and Geothermal Research 137 (2004) 15–31

Fig. 6. Volume fraction gas (vg) versus clast diameter for ash and lapilli fragments collected from most juvenile-bearing units in the Keanaka¯ko’i Ash. Dots represent vg estimated by visual comparison of clasts with modal area plots (e.g. Compton, 1985, pp. 366–367). Crosses are vg of lapilli published earlier (Mastin, 1997). Triangles are obtained from digital image processing of photomicrographs.

more akin to continuous-uprush events at Surtsey than to drainback-induced explosions of Ukinrek or Ha¯lema’uma’u. Moreover, bubble-number data suggest that all unit II subunits erupted rapidly, even those containing only moderately or poorly vesicular tephra. Whether all hydromagmatic phases during this period were violent is less clear; less violent phases may have failed to eject debris out of the caldera where it could be preserved. The cause of repeated, high-flux rate eruptive pulses is not known but could be related to caldera subsidence. High ascent rates could explain some but not all of Fig. 7. Moderately vesicular ash from unit IIA2. the fragmentation and dispersal that we associate with L.G. Mastin et al. / Journal of Volcanology and Geothermal Research 137 (2004) 15–31 27

5.1. The importance of turbulence

We hypothesize that the degree of fragmentation is controlled by the turbulence of mixing. Photographs of continuous-uprush phases at Surtsey (Thorarins- son, 1964) show fully turbulent jets with exit velocities exceeding 100 m/s (Moore, 1985). Turbulence occurs where perturbations along the jet margins are notdampedbyviscousforcesanddevelopinto eddies and vortices (White, 1991, p. 471). In circular jets, the jet margin becomes unstable at Reynolds numbers (Re) as low as 4 (Kaplan, 1964), though full turbulence at the vent is not ensured unless flow is turbulent in the upper conduit (i.e. Re> f 2300, where Re u quD/l, q = fluid density, u =velocity, Fig. 8. Bubble-number density of samples from the Keanaka¯ko’i D = conduit diameter and l is fluid viscosity. During Ash (dots), from lava-fountain tephras (triangles; data from Mangan basaltic lava-fountain eruptions, flow in the upper and Cashman, 1996) and from effusive lavas (crosses; data from conduit may be either laminar or turbulent (Fig. 10), Mangan et al., 1993). Inset is a photomicrograph of one clast (508f- 7), taken under reflected light, and the line-drawing interpretation of with laminar or unstable flow during less vigorous the number of bubbles in this clast. The number of bubbles is 18, events and full turbulence during more vigorous and the clast area (excluding bubbles) is 0.00324 cm2, giving a ones. number density of ((18 bubbles)/0.00324 cm2))3/2 =7 105 cm 3. Engineering tests have long established that jet turbulence increases rates of heat transfer and chem- explosivity. Dry Ki¯lauean high-flux eruptions (lava ical reaction between mixing fluids by orders of fountains) certainly produce more fragmental debris magnitude (e.g. Burmeister, 1983, p. 394). Higher than low-flux eruptions, but the fragmentation process rates of turbulent shear cause more efficient mixing in dry fountaining is primarily gas expansion. In the as eddies progress to ever smaller scales. The finest- Keanaka¯ko’i Ash, many juvenile clasts are poorly scale vortices dampen out at the so called Kolmogorov vesicular and blocky in shape, implying fragmentation length scale (j u l3d/q3U3)1/4, where d is boundary- by thermal fracture rather than bubble expansion. We layer thickness and U is mean velocity, Kolmogorov, surmise that thermal fractures develop on clast surfa- 1941; White, 1991, p. 471). For magma–steam–water ces when they contact water or stream. Fracture mixtures moving at tens of meters per second (based spacing, which determines the final grain size, is on observed lava-fountains; Mangan and Cashman, influenced by the temperature contrast at the interface. 1996) within a boundary layer a meter or two wide, the Rapidly formed surfaces are hotter and likely to host Kolmogorov scale would be decimeters or smaller more finely-spaced thermal cracks than surfaces depending on whether the viscosity and density of formed by slow, taffy-like stretching. The glassy rinds gas, water or magma are used in the calculation. The between these fractures probably shed as the interface scale of mixing is fully sufficient to incorporate strains and deforms (Fig. 9a); the more rapid the centimeter-to decimeter-sized volumes of water into deformation, the more extensive and fine the shedding melt as required to generate molten fuel-coolant inter- of debris. Fragmentation by this mechanism should be actions (Buettner and Zimanowski, 1998). finest and most extensive, and heat-transfer rates from magma greatest, when ascent rates are high and when 5.2. Water depth and volatile equilibration melt is violently torn and deformed as it enters a lake or the atmosphere. Fine particles are likely to originate Turbulent entrainment relations can be used to on the clast margins, which cool before clast interiors constrain water depth from the deposit characteristics. and should, as our results suggest, contain the highest Immediately above the jet exit (‘‘A’’, Fig. 9b), jets dissolved volatile concentration. entrain ambient fluid within a turbulent boundary layer 28 ..Msi ta./Junlo ocnlg n etemlRsac 3 20)15–31 (2004) 137 Research Geothermal and Volcanology of Journal / al. et Mastin L.G.

Fig. 9. (a) Schematic illustration of the fragmentation sequence of a magma blob entrained in turbulent flow. Fragmentation is assumed to occur by extension and shear of the droplet, supplemented by growth of cooling fractures and shedding of the glassy rind on clast margins. (b) Velocity profiles and boundary layers around an axisymmetric turbulent jet exiting into ambient fluid. L.G. Mastin et al. / Journal of Volcanology and Geothermal Research 137 (2004) 15–31 29

and uj are the average jet diameter, density and velocity between the exit and h; qa is density of the ambient fluid; and a is an empirical entrainment coefficient, which is roughly 0.08 in the self-similar region (‘‘C’’, Fig. 9b) and somewhat less in the initial region (‘‘A’’).6 The mass influx calculated by this method may be compared with the mass flux of the 2 jet, Mj = pD qjuj/4. For a jet of magma and gas exiting into water, realistic values for D (f 5 m, based on conduit modeling of observed flow rates; Mastin, 2002), qa 3 3 (1000 kg/m ), qj (500–1500 kg/m assuming a moderately to highly vesicular magma), uj (50–100 m/s based on observed lava-fountain velocities; Mangan Fig. 10. Plot of magma viscosity versus mass flux (m), showing the and Cashman, 1996) and a (0.06–0.08), suggest that threshold for fully turbulent flow in a basaltic conduit. For a given the mass flux of water entrained will equal the mass conduit diameter, diagonal lines separate regions of fully turbulent flux of the magma–gas mixture when water depth is a flow (right) from laminar or unstable flow (left). The two dashed few to several vent diameters. A magma:water ratio of vertical lines bound the range of m produced by lava-fountains at f Pu’u O’o (from volume-flow rates in Parfitt et al., 1995 assuming a 1:1 is much less than the optimal ratio ( 3–5.5:1) bubble-free magma density of 2500 kg/m3). Reynolds numbers are for converting thermal to kinetic energy (Wohletz, calculated by substituting the equation for mass flux (m = pquD2/4) 1986), and would leave most (60–80%) of the water into the Reynolds equation (Re = quD/l) to obtain Re =4m/(pDl). in liquid from (Mastin, 1995), very likely causing Long tick marks give basalt viscosity at specified temperatures using deposits to remobilize into watery slurries. Optimal relations from Shaw (1972). The horizontal dotted line represents the viscosity at temperatures of the Pu’u O’o melts (Thornber, 2001); magma:water ratios would likely require water depths bubbles could increase viscosity up to a few times; Pal, 2003). on the order of a single vent diameter, perhaps less. Holocene Kı¯lauean basalt temperatures have ranged from about 1100 Significantly, at such shallow depths, water would be to 1300 jC (e.g. Clague et al., 1995b). Numerical calculations using entrained only into the boundary layer and would mix Conflow, a publicly available program (Mastin, 2002), indicate the with magma in the core of the jet only after escaping mass flow rates at Pu’uO’o could have been delivered through a conduit a few to several meters in diameter. into the atmosphere. This inference is consistent with observations from Surtsey, in which the margins of whose thickness increases linearly until, at a height (h) continuous-uprush jets contained black ash and white of several times the vent diameter (D), it has penetrated steam, but the jet core glowed a visible red color at into the potential core of the jet (‘‘B’’, Fig. 9b). Further night (Thorarinsson, 1964; Moore, 1985). downstream (‘‘C’’, Fig. 9b), the jet diameter (d) Although accretionary lapilli are common in fine- increases linearly and centerline velocity decreases grained beds of unit II, coarse-grained beds of units inversely with distance. Jet momentum remains con- IIA2 and IIB2 contain none, and there is no evidence stant with distance but total mass flux increases as the for soft-sediment deformation or other remobilization. jet continues to entrain sorrounding fluid. We therefore suspect that that these coarse-grained In experiments using jet and ambient fluids of beds erupted through water depths on the order of a similar properties (e.g. water or saline solution vent diameter. Their elevated volatile concentrations injected into water), the rate of entrainment follows the relation (Eq. (6) of Wilson et al., 1995): 6 rffiffiffiffiffi Magma injection into water involves phase changes and q viscosity contrasts that may result different entrainment coefficient ¯ j Mi ¼ pDhqaauj than the a = 0.08, which is derived from jets of similar fluids. qa Therefore, the implications of these relations are only presented semi-quantitatively here. Experimental studies using magma and where Mi is the mass per unit time of ambient fluid water (or analogue fluids) are required for more quantitative added to the jet between the jet exit and height h; D¯ , qj analysis. 30 L.G. Mastin et al. / Journal of Volcanology and Geothermal Research 137 (2004) 15–31 therefore probably reflect disequilibrum rather than Compton, R.L., 1985. Geology in the Field. Wiley, New York, elevated quenching at pressure in a deep lake. Could 398 pp. Dana, J.D., 1888. History of changes of the Mt. Loa craters. Amer- the disequilibrium have resulted from high ascent ican Journal of Science 36 (212), 81–112 (ser. 3). rates alone? We suspect not, as non-hydromagmatic Decker, R.W., Christiansen, R.L., 1984. Explosive eruptions of lava-fountain tephras (Fig. 2) appear to have equili- Kı¯lauea Volcano, Hawaii. In: National Research Council, et al., brated before cooling. High volatile concentrations in (Eds.), Explosive Volcanism: Inception, Evolution, and Hazards. these glasses reflect both rapid ascent and rapid National Academy Press, Washington, DC, pp. 122–132. Devine, J., Sigurdsson, H., 1995. Petrology and eruption styles of quenching by water at the surface. In this sense, these Kick ’em-Jenny submarine volcano, Lesser Antilles island arc. deposits provide a freeze-fame of degassing at the Journal of Volcanology and Geothermal Research 69, 35–58. time of eruption that is unattainable in non-hydro- Dibble, S., 1843. A history of the Sandwich Islands. T.G. Thrum, magmatic tephra. Lahainaluna, Hawaii, 451 pp. Dixon, J.E., Clague, D.A., Stolper, E.M., 1991. Degassing history of water, sulfur, and carbon in submarine lavas from Kı¯lauea Acknowledgements volcano, Hawaii, 1956–1983. Journal of Geology 99, 371–394. Dixon, J.E., Stolper, E.M., Holloway, J.R., 1995. An experimental

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