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Journal of Volcanology and Geothermal Research 137 (2004) 15–31

What makes hydromagmatic eruptions violent? Some insights

from the Keanakako’i Ash, Kılauea Volcano, Hawai’i$

Larry G. Mastina,*, Robert L. Christiansenb, Carl Thornbera,Jacob Lowensternb, Melvin Beeson1

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

Abstract

Volcanic eruptions at the summit of Kılauea 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 Keanakako’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 Kılauean 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: Kılauea; Phreatomagmatism; explosive eruptions; vesicular texture; turbulence

1. Introduction hydromagmatic tephra deposits erupted from the sum-

The Keanakako’i Ash (Fig. 1) is the most recent,

best exposed and most thoroughly studied of several

0377-0273/$ - see front matter. Published by Elsevier B.V.

doi:10.1016/j.jvolgeores.2004.05.015

$ Supplementary data associated with this article can be found,

in the online version, at doi: 10.1016/j.jvolgeores.2004.05.015.

* Corresponding author. Tel.: +1-360-696-7518; fax: +1-360-

993-8980.

E-mail address: [email protected] (L.G. Mastin).1 formerly U.S. Geological Survey, MS 910, 345 Middlefield

Road, Menlo Park, CA 94025, USA.

mit of Kılauea volcano, Hawai’i, since the late Pleis-

tocene. Originally thought to have been produced

during a single eruption around A.D. 1790 (Dana,

1888), the number of eruptions represented by the

deposit has been repeatedly revised (Stone, 1926;

Powers, 1948; Christiansen, 1979; McPhie et al.,

1990). It is currently thought to have resulted from

at least a few eruptions over a period of perhaps a few

centuries that ended around 1790 AD (Swanson et al.,

1998; D. Swanson, written comm., 2001). The most

Fig. 1. (a) Index map of Kılauea showing the areal extent of the

Keanakako’i Ash (heavy dotted line; from R. Christiansen,

unpublished data), (b) close-up of summit showing sample

locations (numbered). Coordinates of numbered locations are given

in Table 2.

L.G. Mastin et al. / Journal of Volcanology and Geothermal Research 137 (2004) 15–3116

extensive Keanakako’i units that likely resulted from

a single eruption (e.g. unit 2 of McPhie et al., 1990)

cover a few hundred square kilometers (R. Christian-

sen, unpublished data, 2002), similar in area to the

deposits of moderately large historical basaltic hydro-

magmatic events (e.g. Taal, 1965; Moore et al., 1966).

Late Pleistocene tephras at Kılauea, however, are

meters thick in exposures 10 km south of the caldera

(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 lap-

illi, 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 atf 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 thanf 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 dur-

ing summit eruptions; It would not, however, guaran-

tee that such eruptions be large and explosive. Along

Kılauea’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 Soufriere 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

and other hydromagmatic eruptions is affected not

only by the injection of magma into water but also the

circumstances under which they mixed. At least, two

mechanisms for explosive magma-water mixing have

been described during central-vent eruptions: (1)

drawdown of magma in a conduit below the water

table, followed by influx of water, and (2) jetting of

magma through surface water or through a wet crater

into which water is entering through the porous walls.

Mechanism (1) was noted at Ukinrek Maars, Alaska,

when a hydromagmatic explosion, observed from the

air on 3 April, 1977, was preceded a few minutes

earlier by the draining of a lava lake and collapse of

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

explosions at Halema’uma’u Crater in 1924, which

took place after a long-lived lava lake drained and hot

conduit walls collapsed (Stearns, 1925; Decker and

Christiansen, 1984).

The jetting of magma through water was described

during continuous-uprush phases of Surtsey Volcano,

Iceland (Thorarinsson, 1964; Moore, 1985). These

eruptions were characterized by cylindrical columns

100–250 m in diameter at their base (Moore, 1985),

that blasted fragmental debris upward at velocities of

f 100 m/s and persisted for hours. They occurred

when seawater access to the vent was ‘‘partially or

wholly blocked2’’ (Thorarinsson, 1964), and were

generally preceded by intermittent hydromagmatic

explosions that repeated with increasing frequency

until jetting became continuous. (The cause of inter-

mittent explosions has been the subject of disagree-

ment (Kokelaar, 1983) and may be a variant on

mechanism (1)). Water was apparently supplied to

the jet by seepage over and through the tephra pile

that lined the vent crater. Continuous-uprush phases

produced much more tephra and dispersed it more

widely than Surtsey’s intermittent explosions (Thor-

arinsson, 1964, p. 44).

In principle, it should be possible to distinguish

mechanisms (1) and (2) using vesiculation textures

and dissolved-volatile contents of juvenile clasts.

Under mechanism (1), if magma existed within a

subaerial lava lake prior to the explosions, it should

contain dissolved volatiles in equilibrium with 1 atm

pressure. Magma quenched during withdrawal should

contain low vesicularity and low bubble-number den-

sities that typify low ascent rates or stagnation (Cash-

man and Mangan, 1994). Under mechanism (2),

juvenile clasts may contain dissolved volatiles equal

to or exceeding equilibrium values at 1 atmosphere

depending on the pressure and degree of equilibration

of the clasts at the time of quenching. Magma that

quenched during rapid ascent should exhibit abun-

dant, fine vesicles with high bubble-number densities

similar to lava-fountain tephras in subaerial environ-

ments (Cashman and Mangan, 1994; Mangan and

Cashman, 1996).

2 Water could not truly have been ‘‘wholly’’ blocked as

Thorarinsson describes, since he also characterizes these events as

hydromagmatic.

In this study, we examine vesiculation textures and

dissolved-volatile concentrations in the lower unit of

the Keanakako’i Ash to constrain magma-water con-

ditions at the time of quenching. The Keanakako’i

Ash is especially suited to this analysis because the

solubility and pre-eruptive volatile content of

Kılauean basalt are well constrained (e.g. Gerlach,

1986; Dixon et al., 1991, 1995; Wallace and Ander-

son, 1998) and vesiculation textures have been exten-

sively characterized (Mangan et al., 1993; Cashman et

al., 1994; Mangan and Cashman, 1996).

2. The Keanakako’i deposit

The Keanakako’i Ash consists of three major units

and several minor ones (Decker and Christiansen,

1984; McPhie et al., 1990). The major units include

(1) well-bedded ash and lapilli of inferred hydro-

magmatic origin containing primarily fresh, juvenile

sideromelane glass (the ‘‘lower, juvenile-rich beds’’ of

McPhie et al., 1990); (2) massive and cross-bedded

ash and lapilli, also of inferred hydromagmatic origin,

containing both juvenile glass and lithic fragments

(the ‘‘middle, mixed beds’’); and (3) block fall and

cross-bedded surge deposits composed almost exclu-

sively of lithic debris (the ‘‘upper lithic beds’’). These

units are named II, III and V, respectively, in the

nomenclature of Decker and Christiansen and 0–4,

6–10, and 11–16 in the nomenclature of McPhie et

al. (1990). The units are underlain by a massive

reticulite bed (unit I of Decker and Christiansen),

which in turn is locally underlain by a gray vitric-

lithic ash (D. Swanson, written comm. 2002). The

three units are overlain on the northwest and south-

west sides of the caldera by a pumice layer a few tens

of centimeters thick (unit VI of Decker and Christian-

sen; the ‘‘Golden Pumice’’ of Sharp et al., 1987).

Southwest of the caldera, along the contact between

units III and V of Decker and Christiansen are

scattered mats of Pele’s hair and tears (unit IV of

Decker and Christiansen), inferred to have erupted

from a fissure whose lavas are intercalated with the

Keanakako’i Ash (unit 1790f of Neal and Lockwood,

in press). In addition, at the base of the middle, mixed

beds (according to McPhie et al., 1990), or in the

upper part of unit II (according to Decker and Chris-

tiansen, 1984), is a scoria bed up to a few tens of


centimeters thick (unit 6 of McPhie et al.) whose axis

of dispersal extends to the southeast, in contrast to the

southwestward dispersal of most units in the Keana-

kako’i deposit.

2.1. Evidence for depositional breaks

In one of the first written discussions of the

deposit, Dana (1888) assumed it resulted from a single

eruption in 1790 AD that was known from oral

accounts (Ellis, 1827; Dibble, 1843). Early 20th

century authors attributed only the upper lithic unit

to the eruption in 1790 and lower units to earlier

eruptions (Hitchcock, 1909; Powers, 1916; Stone,

1926; Stearns and Clark, 1930; Wentworth, 1938;

Finch, 1947; Powers, 1948). Later, Christiansen

(1979) and Decker and Christiansen (1984) proposed

that all members between the basal reticulite and the

Golden Pumice erupted during a single sequence in or

around 1790 AD. McPhie et al. (1990) concurred but

noted three horizons of erosion and reworking that

suggested pauses of unknown and possibly long

duration. Most recently, Swanson et al. (1998) cited

organic horizons, water-erosion features, archaeolog-

ical artifacts and carbon dates as evidence for at least

few large eruptive sequences, beginning shortly after

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-

gest evidence for non-depositional periods lie imme-

diately above the basal reticulite, slightly below and

above unit 6 of McPhie et al., and at the base of the

upper lithic beds (McPhie et al., 1990).

2.2. Field locations

Our study concentrates on unit II of Decker and

Christiansen, using samples collected in Sand Wash

(locations 61 and 84, Fig. 1b) and at a fissure a

kilometer north of Sand Wash (location 62, Fig. 1b).

At these locations, unit II contains three subunits: two

coarsening-upward sequences3 of well-bedded ash

and lapilli (units IIA and IIB) and a sequence of fine

3 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

mdf =� 0.38–0.62, rf = 1.51–2.10 for ash and lapilli in the upper

half.

ash beds4 with sparse lapilli horizons (unit IIC; Figs. 2

and 3). Conformable bedding within subunits IIA and

IIB suggests that each was deposited during a single

eruptive pulse or rapid series of pulses with no

significant pauses. The contacts at the top of units

IIA and IIB are sharp and truncate underlying beds at

a low angle (Fig. 3b), suggesting some centimeters of

erosion, perhaps by surges or wind, between deposi-

tional pulses. The upper contact of unit IIA lacks

recognizable evidence of water erosion, suggesting

the pause was short and without rain. The upper

contact of unit IIB exhibits scarce water-erosion

rivulets and locally overlies cross-bedded and appar-

ently wind-reworked debris from underlying units

(McPhie et al., 1990), suggesting a longer pause of

unknown duration.

McPhie et al. (1990, p. 351) speculated that

juvenile-rich beds resulted from repeated explosions

that began with vesiculation-driven magma ascent

and ended with drainback of degassed magma.

Mastin (1997), used quantitative vesicularity and

bubble-number density measurements to argue that

some horizons in units IIA2 and IIB2 involved

sustained, high magma discharge. Whether other

beds involved sustained high discharge was not

investigated.

3. Sample collection and analysis

We collected samples by trowel from the outcrop

(Tables 1 and 2) and measured the volume percent of

fresh glass, altered glass, olivine and lithic fragments

by point-counting grains mounted in thin sections

(f 400 grains per section; Table 3). We analyzed

dissolved-volatile concentrations using the electron

microprobe for sulfur and Fourier transform infrared

(FTIR) spectrometry for water and CO2. Juvenile

glass in the deposit ranges in vesicularity from 0%

to >80%, requiring us to prepare samples for FTIR

analysis using two methods: (1) slightly-to-moderate-

ly vesicular clasts, 1–2 mm in diameter were embed-

ded in a 2.5-cm diameter wafer of epoxy that was then

ground and polished, flipped, remounted on a glass

slide, then ground to a thickness of 70–165 Am, and

4 mdf =� 0.67–3.71, rf = 1.31–2.53 based on unpublished

grain-size analyses.


polished on the other side; (2) highly vesicular lapilli

were ground with a mortar and pestle and flakes f 1

mm in diameter were handpicked, mounted in epoxy

and doubly polished as described above.

3.1. Dissolved volatile analysis

Before removing the polished grain mounts from

their glass slides, each fragment was analyzed for

sulfur by electron microprobe using the JEOL JXA-

8900 Electron Probe Microanalyzer at the USGS

Menlo Park Laboratory. These samples were a subset

of more than 23 thin sections containing tephra frag-

ments analyzed for sulfur during 6 microprobe ses-

sions of 1–2 days each. For reference, we also

analyzed clasts from non-hydromagmatic units I, IV,

VI and the 1959 Kılauea Iki tephra (Table 1). Values

of S and K2O in Table 4 and the major-element data

(see supplementary data in the online version) repre-

sent averages of two to four analyses per clast (exact

numbers are listed online). During each session, we

used an acceleration voltage of 15 keV, a beam current

of 25 nA, a beam diameter of 10–15 Am, a barite

standard for sulfur, and a peak counting interval of 80

s. Results for elements other than sulfur, and standards

used, are provided in the online supplement. Of 343

analyses, 330 gave totals between 97% and 101%.

The precision of sulfur concentrations (ppm) is

estimated from counting statistics and reproducibility

of USNM basaltic glass standard VG-2 (Jarosewich et

al., 1979), which is similar in composition to the

analyzed glasses. The standard deviation in sulfur

values of VG-2 glass measured during individual

microprobe sessions is 57 ppm (the half-width of

the error bar in Fig. 2); however, mean VG-2 sulfur

analyses for a given session range from 1080 to 1350

ppm, suggesting some drift between sessions. Sulfur

calibration numbers were therefore adjusted so that

the average VG-2 analysis in each session equals

1340 ppm, the value obtained by Dixon et al. (1991).

After microprobe analysis, the grain mounts were

removed from their glass slides and bubble-and crystal-

free regions of glass were targeted for FTIR analysis

using apertures above and below the sample. Measure-

ments were performed using a Nicolet5 Magna 750

5 Use of trade names in this document does not imply

endorsement by the USGS.

spectrometer with an attached SpectraTechR Analyti-

cal-IR microscope that utilizes a liquid-N2-cooled

MCT-A detector. We collected 512 scans per analysis

on each sample and on backgrounds following each

sample analysis. The measured absorbance of infrared

radiation is directly proportional to the concentration of

H2O in the glass, adjusted for sample density and

thickness. Weight percent H2O was calculated accord-

ing to Eq. (15) of Ihinger et al. (1994) for the OH�

stretch at 3570 cm� 1, assuming a glass density of

2700F 100 g l� 1, an extinction coefficient (e) of

63F 3 l mol� 1 cm� 1 and sample thickness as mea-

sured by a MitutoyoR Digital Micrometer. Uncertainty

for e, sample thickness and density are all about 5%,

so that propagated errors are all between 10% and

15% relative.

We detected no molecular H2O (peak at 1630

cm� 1) or CO3�(1515 and 1430 cm� 1); we analyzed

two clasts in sample 517B twice, collecting 1024

scans during the second analysis, to verify the absence

of these peaks. The detection limit of dissolved

CO3�is f 50 ppm, suggesting that the melt degassed

to less than f 5 MPa (using solubility relations of

Dixon et al., 1995), or about 600 m depth (assuming a

pressure gradient of 25 MPa/km). The absence of a

molecular water peak is consistent with the experi-

mentally measured speciation of dissolved water in

basaltic glass at this concentration (Dixon et al.,

1995). Meteoric water adsorbed into the clasts at

magmatic temperature could also speciate of OH�,

but water absorbed at lower temperature would likely

remain in molecular form (P. Wallace and J. Dixon,

written comm., 2001). If water were absorbed into the

clasts during or after the eruption, we would expect it

to be absorbed over a range of temperatures and to

leave at least some water in molecular form. The

complete absence of molecular H2O in these clasts

therefore suggests to us that all dissolved water is

magmatic.

3.2. Vesicularity and bubble-number density

To obtain bubble-number densities, all of the

dense, hand-picked clasts analyzed by FTIR were

photographed in thin section under reflected light.

Printed photomicrographs were overlain with trans-

parencies and the outlines of the clasts, and of internal

bubbles, were traced in ink. The inked transparencies


were then digitally scanned and the clast diameter,

area, and the number and size distribution of bubbles

were computed using the software Scion ImageR. Thenumber of bubbles per clast ranged from zero to

several dozen; too few to derive meaningful size-

distribution statistics but enough to estimate the or-

der-of-magnitude bubble-number density.

The highly vesicular clasts were destroyed when

ground by mortar and pestle and could not be ana-

lyzed for bubble-number density. Prior to grinding,

however, their vesicularity was measured by weighing

the samples in and out of water following the method

of Houghton and Wilson (1989). We calculated clast

density (q, kg/m3) using the formula q = qw(wa/

(wa�ww)), where qw is the density of water (1000

kg/m3), and wa and ww are the measured weights in air

and in water, respectively. The volume-fraction gas

(vg)—i.e. the fraction of the clast volume composed of

bubbles—was calculated using the formula vg = 1� q/qg, where qg is the density of basaltic glass, taken as

2700 kg/m3. We calculated the equivalent radius (r) of

each lapillus (assuming a spherical shape) from the

formula r = (3wa/(4pq))1/3. Densities are generally

repeatable to F 10%.

During the course of our analysis, we observed that

vesicularity varied systematically from one strati-

graphic unit to the next and from ash to lapilli-sized

clasts within a given unit. To quantify these variations,

we compare vg from 3000 clast density measurements,

published earlier (Mastin, 1997), with vg fromf 3000

ash-sized grains, estimated visually in thin section by

comparing the abundance of bubbles to diagrams of

modal percentage (Compton, 1985, Appendix 3).

Diagrams in Compton (1985) are given for percen-

Fig. 2. (left) Keanakako’i Ash, showing stratigraphic units exposed at San

center) total dissolved sulfur in glass from microprobe measurements; (right

right) components of these units from point counts (Table 3). To the left o

(1990) (left) and Decker and Christiansen (1984) (right). Tick marks at top

glass at saturation, calculated using a best-fit exponential curve through six

that contained no dissolved CO2 and were tested at pressures < 200 MPa

equilibrium concentration (0.09 wt.%) at p= 0.1013 MPa (e.g. Wallace an

best-fit curve has the form H2O= 0.295p.515 (r2 = 0.9998), where H2O

error = 0.003 wt.% H2O at p= 0.1013 MPa, 0.029 wt.% at 2.5 MPa). Th

Christiansen (1984): unit 6 of McPhie et al., which is thickest near Kean

correlate with unit IIC2 at Sand Wash. However, the isopach map of uni

Moreover, these units differ petrologically (IIC2 at Sand Wash contains

juvenile scoria). Scattered fragments of scoriaceous debris at the base of u

remnants of unit 6. Further refinements to this stratigraphy are underway. N

(1984) are shown in the stratigraphy of McPhie et al. because no units of

tages of 0.5, 1, 1.5, 3, 5, 7, 10, 15, 10, 25, 30, 35, 40,

45 and 50. We consider those intervals to indicate the

resolution of the estimation technique. Results of

these measurements, as well as microprobe data and

lapilli clast density, are posted online.

4. Results

Sulfur values in matrix glass from unit II (Fig. 2)

average 330 ppm with a standard deviation of 80

ppm—significantly above the 100–150 ppm typically

measured for Kılauean glass that has equilibrated to

atmospheric pressure (Swanson and Fabbi, 1973;

Mangan and Cashman, 1996; Thornber, 2001). Glass

in non-hydromagmatic units I, IV, VI and the Kılauea

Iki tephra, all contain dissolved sulfur values consis-

tent with equilibrium degassing to 1 atm pressure.

Dissolved sulfur in glass inclusions in olivine ranges

from about 250 to 1530 ppm, with most around

1000–1200 ppm (Fig. 2). Considering the middle of

this range to represent undegassed inclusions and 100

ppm to be totally degassed, most hydromagmatic

glass fragments are about 70–80% degassed in sulfur.

The concentration of dissolved water in juvenile

glass also measurably exceeds equilibrium saturation

under atmospheric conditions (Fig. 2). Sulfur and

water values (Table 4) are moderately correlated

(r2 = 0.69) supporting the view that the water is

magmatic in origin. Measured H2O concentrations

correspond to quench pressures (under equilibrium

H2O saturation) of f 0.3–2.3 MPa (Fig. 2), equiva-

lent to 10–100 m depth at lithostatic pressure (as-

suming a pressure gradient of 25 MPa/km), or f 25–

d Wash plus units I and VI of Decker and Christiansen (1984); (left

center) total dissolved water in matrix glass, measured by FTIR; (far

f the stratigraphic section are unit names according to McPhie et al.

of the H2O plot indicate the equilibrium dissolved water in MORB

experimental measurements from Dixon et al. (1995), using samples

. One extra data point was added to this data set: the well-known

d Anderson, 1998; Mangan et al., 1993; Cashman et al., 1994). The

is in weight percent and p is pressure in megapascals (standard

is section contains one revision to the stratigraphy of Decker and

akako’i Crater, was thought by Decker and Christiansen (1984) to

t 6 in McPhie et al. suggests that it does not extend this far west.

pumice and lithics, whereas unit 6 at Keanakako’i Crater contains

nit IIC1 at Sand Wash (hollow diamonds in the sulfur plot) may be

o unit names corresponding to unit IIC of Decker and Christiansen

McPhie et al. appear to correlate with them.

Fig. 3. Photos of the Keanakako’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

amount of H2O lost to degassing can be assessed from

the ratio of H2O/K2O (Table 4, Fig. 4). Both of these

components are incompatible in a crystallizing

Kılauean melt and, in the absence of degassing,

their ratio should remain constant at about 1.3 (Dixon

et al., 1991; Wallace and Anderson, 1998). Average

K2O (0.44F 0.05 wt.%) in clasts from hydromag-

matic units imply pre-eruptive water content of

0.57F 0.06 wt.%. Measured H2O/K2O ratios (0.39–

1.02) suggest that the melt had lost 25–80% of its

H2O by the time it quenched.

Fig. 5 shows an inverse relationship between

dissolved water and vg, which is consistent with a

Table 2

Sample locations shown in Fig. 1

Number Easting Northing

1 259221 2146965

21 261979 2146795

61 259261 2145550

62 259240 2146443

67 259715 2148958

84 259237 2145737

Easting and Northing are given in UTM Zone 5 coordinates.

Table 1

Summary of samples collected and analyzed for this paper

Sample Unit Location Type of

numberD and C McPhie

number analysis

426 Iki Iki 21 mp

3HK215Q VI 1 mp

602 IV 84 mp

423 IIC4 62 mp, pc

525 IIC3 61 d, ves

422 IIC3 62 mp, pc

407b IIC3 61 pc

523 IIC3 61 d, ves

521 IIC3 61 d, ves

421 IIC2 62 mp

407a IIC2 61 mp, pc

520 IIC2 61 mp, FTIR, bnd, d

519 IIC1 61 d, ves

518 IIC1 61 ves

420 IIC1 62 pc

517 IIC1 61 mp, FTIR, bnd

516 IIC1 61 d, ves

419 IIC1 62 pc

514 6? 61 mp, d

513 IIB2 4 61 d, ves

418 IIB2 4 62 d

417 IIB2 4 62 mp, pc

512 IIB2 4 61 d, ves

406 IIB2 4 61 d

405 IIB2 4 61 mp, pc

511 IIB2 4 61 mp, FTIR, bnd, d, ves

416 IIB1 3 62 pc

509 IIB1 3 61 d

508 IIB1 3 61 mp, FTIR, bnd

403 IIB1 3 61 mp, pc

503 IIB1 3 61 mp, FTIR, bnd

413 IIA2 2 62 mp, pc

414 IIA2 2 62 d

502 IIA2 2 61 mp, FTIR, bnd, d

402 IIA2 2 61 d

412 IIA2 2 62 mp, d

401 IIA2 2 61 mp

501 IIA2 2 61 d

411 IIA1 1 62 mp, pc

526 IIA1 1 61 ves

409 IIA1 1 62 mp, FTIR, bnd

404 IIA1 1 61 pc

432 I 67 mp, d

Abbreviations in the right column include clast density measure-

ments (‘‘d’’), vesicularity measurements from thin section (‘‘ves’’),

microprobe glass analyses (‘‘mp’’), FTIR analyses (‘‘FTIR’’),

bubble-number density (‘‘bnd’’) and point counts (‘‘pc’’).

Table 3

Summary of point count results

Sample Unit Components

D and C McPhie Altered

glass

Fresh

glass

Lithic Crystal

n

423 IIC4 17% 63% 12% 8% 400a

422 IIC3 14% 80% 0% 6% 400

407b, 1f IIC3 18% 76% 4% 2% 400

407a, 1f IIC2 49% 43% 6% 2% 400

420 IIC1

(top)

11% 86% 0% 3% 401

419 IIC1

(base)

73% 24% 0% 3% 400

417 IIB2

(top)

4 6% 78% 15% 2% 400

405, 1f IIB2 4 11% 81% 5% 3% 390

416 IIB1

(top)

3 9% 91% 0% 1% 400

403, 1f IIB1 3 3% 96% 1% 0% 404

413 IIA2

(top)

2 8% 73% 9% 10% 400

411 IIA1

(top)

1 12% 87% 1% 1% 400

404, 1f IIA1 1 7% 86% 7% 1% 400

a The standard error for point counts of 400 grains is less than

4% (Van der Plas and Tobi, 1965).


scenario in which clasts were quenched during active

vesiculation. Dissolved H2O is lowest in the highly

vesicular clasts of unit IIB2. Nearly all data lie

below the line that represents closed-system gas

exsolution with total water (dissolved plus exsolved)

equal to 0.55 wt.%. Large, highly vesicular lapilli

appear to have suffered more gas loss than smaller,

less vesicular ash.

These data clearly show that hydromagmatic glass

did not fully degas in sulfur or water. In contrast, non-

hydromagmatic clasts are completely degassed in

sulfur. We did not measure dissolved water in non-

hydromagmatic tephra, but tephra from phase 1 of the

Kılauea Iki eruption, analyzed by Wallace and Ander-

son (1998, fig. 1), are in equilibrium with 1 atm

pressure (phase 1 tephra erupted before degassed lava

Table 4

Summary of FTIR and bubble-number density measurements

Sample Unit Thickness

(Am)

vg Diameter

(mm)

Clast area

(mm2)

Number of

bubbles

Number density

(cm� 3)

S

(wt.%)

H2O

(wt.%)

K2O H2O/K2O

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

Unit names are those of Decker and Christiansen (1984). Clast diameter (D) is calculated as D ¼ 2ffiffiffiffiffiffiffiffiA=p

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

that water, like sulfur, is elevated due to rapid quench-

ing by external water. Whether these clasts quenched

at elevated pressure or at 1 atm before they could

degas to equilibrium is the subject of later discussion.

4.1. Vesiculation textures

In general, pyroclasts of a given size are less

vesicular in units IIA1, IIB1, more vesicular in units

IIA2 and IIB2. Within individual units, average vesic-

Fig. 4. Dissolved H2O versus K2O in matrix glass from

Keanakako’i Ash. Melt that has not lost H2O to degassing plots

along the line H2O/K2O= 1.3, while completely degassed melt

would plot along a line having zero slope and an intercept at

H2O= 0.09 wt.%. Intermediate degrees of degassing are represented

by lines having intermediate slopes, intersecting the H2O/K2O= 1.3

line at H2O= 0.09 wt.%.

Fig. 5. Weight percent dissolved water versus volume fraction gas in

melt. Squares are measurements (symbol size is proportional to clast

diameter). Lines give volume-fraction gas versus dissolved water

for closed-system degassing for a total water content (dissolved plus

exsolved) of (top to bottom) 0.55, 0.44, 0.33 and 0.22 wt.%,

respectively. The lines represent loss of 0%, 20%, 40% and 60% of

the original gas, respectively.


ularity increases with average grain size (Fig. 6).

Although vg was estimated for the ash-and lapilli-sized

clasts using two different methods, we do not think

that sample bias in one or both methods is responsible

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 sim-

ilar 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 diam-

eter. 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)

range from 9� 104 to 1�107 cm� 3 similar to those

measured in unit IIB2 by Mastin (1997). Though the

earlier data were obtained only from highly vesicular

lapilli, the new data come from ash-sized clasts with

vesicularity ranging down to about 20%. All these

values exceed number densities obtained from vent

samples of effusive eruptions (3000–5000 cm� 3;

Mangan et al., 1993) and most exceed values mea-

sured in lava-fountain tephra from Pu’uO’o (19,000–

180,000 cm� 3; Mangan and Cashman, 1996).

Unlike our Keanakako’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

number versus bubble diameter, they inferred that

only about 13–40% of bubbles coalesced from two

or more smaller ones—not enough to account for the

< 100� lower number densities of lava-fountain

clasts relative to Keanakako’i clasts.

5. Implications

These high bubble-number densities of unit II

suggest rapid magma ascent and eruptive activity

Fig. 7. Moderately vesicular ash from unit IIA2.

Fig. 6. Volume fraction gas (vg) versus clast diameter for ash and lapilli fragments collected from most juvenile-bearing units in the Keanakako’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 Hale-

ma’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

the fragmentation and dispersal that we associate with

Fig. 8. Bubble-number density of samples from the Keanakako’i

Ash (dots), from lava-fountain tephras (triangles; data from Mangan

and Cashman, 1996) and from effusive lavas (crosses; data from

Mangan et al., 1993). Inset is a photomicrograph of one clast (508f-

7), taken under reflected light, and the line-drawing interpretation of

the number of bubbles in this clast. The number of bubbles is 18,

and the clast area (excluding bubbles) is 0.00324 cm2, giving a

number density of ((18 bubbles)/0.00324 cm2))3/2 = 7� 105 cm� 3.


explosivity. Dry Kilauean high-flux eruptions (lava

fountains) certainly produce more fragmental debris

than low-flux eruptions, but the fragmentation process

in dry fountaining is primarily gas expansion. In the

Keanakako’i Ash, many juvenile clasts are poorly

vesicular and blocky in shape, implying fragmentation

by thermal fracture rather than bubble expansion. We

surmise that thermal fractures develop on clast surfa-

ces when they contact water or stream. Fracture

spacing, which determines the final grain size, is

influenced by the temperature contrast at the interface.

Rapidly formed surfaces are hotter and likely to host

more finely-spaced thermal cracks than surfaces

formed by slow, taffy-like stretching. The glassy rinds

between these fractures probably shed as the interface

strains and deforms (Fig. 9a); the more rapid the

deformation, the more extensive and fine the shedding

of debris. Fragmentation by this mechanism should be

finest and most extensive, and heat-transfer rates from

magma greatest, when ascent rates are high and when

melt is violently torn and deformed as it enters a lake

or the atmosphere. Fine particles are likely to originate

on the clast margins, which cool before clast interiors

and should, as our results suggest, contain the highest

dissolved volatile concentration.

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 veloc-

ities exceeding 100 m/s (Moore, 1985). Turbulence

occurs where perturbations along the jet margins are

not damped by viscous forces and develop into

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 Reu quD/l, q = fluid density, u = velocity,

D = conduit diameter and l is fluid viscosity. During

basaltic lava-fountain eruptions, flow in the upper

conduit may be either laminar or turbulent (Fig. 10),

with laminar or unstable flow during less vigorous

events and full turbulence during more vigorous

ones.

Engineering tests have long established that jet

turbulence increases rates of heat transfer and chem-

ical reaction between mixing fluids by orders of

magnitude (e.g. Burmeister, 1983, p. 394). Higher

rates of turbulent shear cause more efficient mixing

as eddies progress to ever smaller scales. The finest-

scale vortices dampen out at the so called Kolmogorov

length scale (ju l3d/q3U3)1/4, where d is boundary-

layer thickness and U is mean velocity, Kolmogorov,

1941; White, 1991, p. 471). For magma–steam–water

mixtures moving at tens of meters per second (based

on observed lava-fountains; Mangan and Cashman,

1996) within a boundary layer a meter or two wide, the

Kolmogorov scale would be decimeters or smaller

depending on whether the viscosity and density of

gas, water or magma are used in the calculation. The

scale of mixing is fully sufficient to incorporate

centimeter-to decimeter-sized volumes of water into

melt as required to generate molten fuel-coolant inter-

actions (Buettner and Zimanowski, 1998).

5.2. Water depth and volatile equilibration

Turbulent entrainment relations can be used to

constrain water depth from the deposit characteristics.

Immediately above the jet exit (‘‘A’’, Fig. 9b), jets

entrain ambient fluid within a turbulent boundary layer

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

etal./JournalofVolca

nologyandGeotherm

alResea

rch137(2004)15–31

28

Fig. 10. Plot of magma viscosity versus mass flux (m), showing the

threshold for fully turbulent flow in a basaltic conduit. For a given

conduit diameter, diagonal lines separate regions of fully turbulent

flow (right) from laminar or unstable flow (left). The two dashed

vertical lines bound the range of m produced by lava-fountains at

Pu’u O’o (from volume-flow rates in Parfitt et al., 1995 assuming a

bubble-free magma density of 2500 kg/m3). Reynolds numbers are

calculated by substituting the equation for mass flux (m= pquD2/4)

into the Reynolds equation (Re= quD/l) to obtain Re= 4m/(pDl).Long tick marks give basalt viscosity at specified temperatures using

relations from Shaw (1972). The horizontal dotted line represents the

viscosity at temperatures of the Pu’u O’o melts (Thornber, 2001);

bubbles could increase viscosity up to a few times; Pal, 2003).

Holocene Kılauean basalt temperatures have ranged from about 1100

to 1300 jC (e.g. Clague et al., 1995b). Numerical calculations using

Conflow, a publicly available program (Mastin, 2002), indicate the

mass flow rates at Pu’uO’o could have been delivered through a

conduit a few to several meters in diameter.

6 Magma injection into water involves phase changes and

viscosity contrasts that may result different entrainment coefficient

than the a= 0.08, which is derived from jets of similar fluids.

Therefore, the implications of these relations are only presented

semi-quantitatively here. Experimental studies using magma and

water (or analogue fluids) are required for more quantitative

analysis.


whose thickness increases linearly until, at a height (h)

of several times the vent diameter (D), it has penetrated

into the potential core of the jet (‘‘B’’, Fig. 9b). Further

downstream (‘‘C’’, Fig. 9b), the jet diameter (d)

increases linearly and centerline velocity decreases

inversely with distance. Jet momentum remains con-

stant with distance but total mass flux increases as the

jet continues to entrain sorrounding fluid.

In experiments using jet and ambient fluids of

similar properties (e.g. water or saline solution

injected into water), the rate of entrainment follows

the relation (Eq. (6) of Wilson et al., 1995):

Mi ¼ pDhqaauj

ffiffiffiffiffiqj

qa

r

where Mi is the mass per unit time of ambient fluid

added to the jet between the jet exit and height h; D, qj

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

jet, Mj = pD2qjuj/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

(1000 kg/m3), qj (500–1500 kg/m3 assuming a mod-

erately to highly vesicular magma), uj (50–100 m/s

based on observed lava-fountain velocities; Mangan

and Cashman, 1996) and a (0.06–0.08), suggest that

the mass flux of water entrained will equal the mass

flux of the magma–gas mixture when water depth is a

few to several vent diameters. A magma:water ratio of

1:1 is much less than the optimal ratio (f 3–5.5:1)

for converting thermal to kinetic energy (Wohletz,

1986), and would leave most (60–80%) of the water

in liquid from (Mastin, 1995), very likely causing

deposits to remobilize into watery slurries. Optimal

magma:water ratios would likely require water depths

on the order of a single vent diameter, perhaps less.

Significantly, at such shallow depths, water would be

entrained only into the boundary layer and would mix

with magma in the core of the jet only after escaping

into the atmosphere. This inference is consistent with

observations from Surtsey, in which the margins of

continuous-uprush jets contained black ash and white

steam, but the jet core glowed a visible red color at

night (Thorarinsson, 1964; Moore, 1985).

Although accretionary lapilli are common in fine-

grained beds of unit II, coarse-grained beds of units

IIA2 and IIB2 contain none, and there is no evidence

for soft-sediment deformation or other remobilization.

We therefore suspect that that these coarse-grained

beds erupted through water depths on the order of a

vent diameter. Their elevated volatile concentrations


therefore probably reflect disequilibrum rather than

elevated quenching at pressure in a deep lake. Could

the disequilibrium have resulted from high ascent

rates alone? We suspect not, as non-hydromagmatic

lava-fountain tephras (Fig. 2) appear to have equili-

brated before cooling. High volatile concentrations in

these glasses reflect both rapid ascent and rapid

quenching by water at the surface. In this sense, these

deposits provide a freeze-fame of degassing at the

time of eruption that is unattainable in non-hydro-

magmatic tephra.

Acknowledgements

This paper has significantly benefited from dis-

cussions with Don Swanson, Richard Fiske, Tim

Rose, David Clague, Paul Wallace and Jacqueline

Dixon. Pete Friedman provided a helpful review of

the discussion of jet entrainment. We also wish to

acknowledge the assistance of Robert Oscarson for

microprobe assistance; Margaret Mangan and Kathy

Cashman for advice on collecting and analyzing

vesicle data; and Marvin Couchman for advice on

collecting clast density measurements.

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