Earth and Planetary Science Letters 299 (2010) 126–137
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Earth and Planetary Science Letters
j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / e p s l
Time-and temperature-dependent conduit wall porosity: A key control on degassing
and explosivity at Tarawera volcano, New Zealand
B.M. Kennedy a,b,⁎, A.M. Jellinek b, J.K. Russell b, A.R.L. Nichols c, N. Vigouroux d
a
Geological Sciences, University of Canterbury, Private Bag 4800, Christchurch, 8140, New Zealand
Department of Earth and Ocean Sciences, University of British Columbia, 6339 Stores Road, Vancouver, BC, Canada V6T 1Z4
Institute for Research on Earth Evolution (IFREE), Japan Agency for Marine Earth Science and Technology (JAMSTEC), 2-15 Natsushima-cho, Yokosuka, Kanagawa 237-0061, Japan
d
Department of Earth Science, Simon Fraser University, Burnaby, British Columbia, Canada
b
c
a r t i c l e
i n f o
Article history:
Received 14 February 2010
Received in revised form 18 August 2010
Accepted 23 August 2010
Available online 21 September 2010
Editor: R.W. Carlson
Keywords:
degassing
conduit
magma
volcano
explosive
lava dome
a b s t r a c t
The permeability of volcanic conduit walls and overlying plug can govern the degassing and explosivity of
eruptions. At volcanoes characterized by a protracted history of episodic volcanism, conduit walls are
commonly constructed of quenched magma. During each successive eruptive phase, reheating by ascending
magma can modify the porosity, permeability and H2O content of the conduit wall rocks and overlying plug.
We investigate whether the unusual explosivity of the 1886 basaltic eruption at Tarawera volcano is related to
the heating and degassing of the AD1314 Kaharoa rhyolitic rocks, through which it erupted. We heat cores of
perlitic Tarawera dome rhyolite to 300 °C–1200 °C for 30 min to 3 days at atmospheric pressure. We
characterize time (t)- and temperature (T)-dependent variations in porosity, volatile content and texture
through SEM image analyses. We also directly measure pre- and post-experimental connected and isolated
porosity and water content. We identify four textural/outgassing regimes: Regime 1 (T ≤ 800 °C, t ≤ 2 h), with
negligible textural changes and a significant loss of meteoric water (1.4–0.72 wt.% H2O); Regime 2
(800 ≤ T ≤ 1100 °C, t ≤ 6 h), with cracking and vesicle growth and a 5–10% increase in connected porosity;
Regime 3 (800 ≤ T ≤ 1200 °C, t ≥ 30 min), with healed cracks, coalesced and collapsed vesicles, and overall
reduced porosity; and Regime 4 (T ≥ 1200 °C, t N 30 min), with a collapse of all connected porosity. These
regimes are governed by the temperature of the event (T) relative to the glass transition temperature (Tg) and
the time scale of the event (t) relative to a critical relaxation time for structural failure of the melt (τr). We
identify a quantitative transition from predominantly brittle behavior such as cracking, which enhances
connected porosity and permeability, to viscous processes including crack healing and vesicle collapse, which
act to reduce connected porosity. Applied to the 1886 basalt eruption at Tarawera, we show that progressive
heat transfer ultimately reduced the open porosity and permeability of the conduit walls, thereby partially
sealing the conduit and reducing volatile loss. We argue that this mechanism was an underlying reason for the
exceptional explosivity of the 1886 eruption. We further suggest that textural changes associated with
reheating could explain some of the cyclic deformation and degassing observed at many lava domes
preceding explosive eruptions.
© 2010 Elsevier B.V. All rights reserved.
1. Introduction
Individual volcanic eruptions can shift rapidly in style from
relatively quiescent lava dome extrusion to explosive eruptions (e.g.
Sparks, 2003; Voight et al., 1999). These shifts are generally attributed
to variations in physical properties such as gas content, viscosity,
vesicularity, wall rock permeability, and crystallinity (e.g. Gonnermann and Manga, 2007; Jaupart, 1998; Melnik and Sparks, 2002). To
date experimental and theoretical studies have focussed on variation
of these properties as magma rises and decompresses (e.g. Baker et al.,
⁎ Corresponding author. Geological Sciences, University of Canterbury, Private Bag
4800, Christchurch, 8140, New Zealand.
E-mail address: [email protected] (B.M. Kennedy).
0012-821X/$ – see front matter © 2010 Elsevier B.V. All rights reserved.
doi:10.1016/j.epsl.2010.08.028
2006; Gardner, 2007; Hammer and Rutherford, 2002; Larsen et al.,
2004; Proussevitch et al., 1993; Takeuchi et al., 2009; Yoshimura and
Nakamura, 2008) or is sheared (Gonnerman and Manga, 2003;
Lavallee et al., 2007, 2008; Okamura et al., 2010; Smith et al., 2009;
Tuffen et al., 2003, 2008). Natural pumice, dome rocks and
experimentally decompressed glasses show huge textural variations
in their permeable vesicle and crack networks (Jaupart, 1998; Michaut
and Sparks, 2009; Mueller et al., 2008; Rust and Cashman, 2004; Saar
and Manga, 1999; Takeuchi et al., 2008; Westrich and Eichelberger,
1994; Wright et al., 2009; Yoshimura and Nakamura, in press). Other
experimental studies investigate the effects of temperature on water
speciation, solubility, and magma viscosity (Stolper, 1989; Yamashita,
1999; Zhang et al., 2007). Rocks from conduit walls also exhibit
variation in porosity (Kennedy et al., 2005; Rust et al., 2004; Stasiuk et
B.M. Kennedy et al. / Earth and Planetary Science Letters 299 (2010) 126–137
al., 1996). Yet, no studies have addressed the isobaric time-dependent
textural changes of conduit walls in response to reheating. Here we
address two key questions: 1. How is degassing of the ascending
magma affected by changes in permeability of the conduit walls
during reheating? 2. To what extent can the release of volatiles from
the reheated wall rock contribute towards the eruption?
The vents for many explosive eruptions are often plugged by
fractured, vesicular lava domes or partially filled conduits (e.g.
Johnson and Lees, 2000; Voight et al., 1999). Surprisingly, almost no
attention has been given to the effects of the hot rising magma on the
behaviour (i.e. evolution of texture and volatile content) of the lava
that plugs the volcanic conduit. This is despite observations that show
temperature rises in older lava domes prior to eruption (Wooster and
Kaneko, 1997) that correlate with eruption style (Sahetapy-Engel and
Harris, 2009). We propose that reheating can influence the degassing
of both the older plug and the rising magma. Magmatic water and
resorbed meteoric water dissolved in glass within the old plug may be
available for degassing and vesiculation. This vesiculation in turn
affects the porosity and the permeability of the plug and the ability of
the rising hot magma to degas.
The 1886 Tarawera eruption (Cole, 1970; Nairn, 2002), is one of
only a few examples of basaltic plinian eruptions (Houghton et al.,
2004). At Tarawera, basalt erupts through a pre-existing dome
complex and silicic conduit system (Carey et al., 2007). This may be
a common occurrence at bimodal vents, however, descriptions of
bimodal vent exposures are absent in volcanological literature.
Detailed stratigraphic studies at Tarawera have tracked the shifting
eruption centres and fragmentation level and documented interaction
with groundwater and the pre-existing hydrothermal system (Carey
et al., 2007; Houghton et al., 2004; Sable et al., 2006, 2009). The effect
of rhyolitic conduit wall recycling during this eruption has also been
discussed (Rosseel et al., 2006).
Conduit wall permeability is an important variable in explosive
basaltic eruptions (Houghton and Gonnerman, 2008) but has not been
investigated experimentally. We use laboratory experiments on the
Tarawera rhyolitic lava to show that during an eruption the wall rock
permeability is both time- and temperature-dependent, as is the
release of volatiles from the wall rocks into the erupting magma. We
argue that a reduction in the permeability of the conduit walls as a
result of reheating hindered outgassing and increased the explosivity
of the 1886 Tarawera eruption.
127
2. Methodolgy
2.1. Sampling
We collected samples that contained rhyolite and basalt from the
proximal deposits of 1886 basaltic fissure eruption. The motivation for
this sampling was to collect samples that show evidence for heat
transfer between basalt and rhyolite. Enclave samples were collected
from along the rim of the fissure on the summit of Tarawera lava
dome (Fig. 1). We limited our samples to the crystal rich 1314 AD lava
dome xenoliths/enclaves and excluded the older crystal-poor xenoliths (Carey et al., 2007). We chose a single large homogeneous glassy
and perlitic sample of the lava dome to use for our experimental
starting material. From this sample we drilled cylindrical cores 1 cm in
diameter by 2 cm in length, which were then left in an oven overnight
at 100 °C to remove surface water from the samples. The precise
dimensions, density and porosity of samples were measured prior to,
and following, heating experiments. The volume of these porous
cylindrical cores was calculated from averages of replicate (n = 3)
measurements of diameter and length. This volume and the
sample mass were used to calculate the bulk density (ρbulk) of
the core. Skeletal (or framework) density (ρskeletal) is obtained by
measuring sample volume via helium pycnometry. Connected
porosity (Φconnected) (Table 1) was calculated from skeletal and bulk
density from the relationship: Φconnected = 1 − (ρbulk/ρskeletal). We
obtained values of dense rock equivalent (DRE) density for rock by
crushing three cores and performing pycnometry on the resulting
powders. All experimental cores have the same average DRE density
(2.34 g/cm3 ± 0.10 g/cm3). Using this average value for powder density
we compute total and isolated porosity (Table 1) as: Φtotal = 1 − (ρbulk/
ρpowder), and Φisolated = (ρbulk/ρskeletal) − (ρbulk/ρpowder) (Michol et al.,
2008). Porosity values have an uncertainty of up to 4% associated with
the largest porosities measured by pycnometer (Michol et al., 2008).
2.2. Experiments
We first heated the furnace to 300–1200 °C (± 10 °C) at
atmospheric pressure. This represents the expected temperature
range of conduit rocks in contact with the erupting basalt (Rosseel
et al., 2006). Cylindrical samples (described subsequently) in ceramic
crucibles were placed in the centre of the furnace for a specified time
Fig. 1. Tarawera volcanic edifice shown as: a) geological map illustrating sampling locations on the rim of the 1886 eruption fissure (marked by X), and b) aerial photograph of the
fissure and domes looking SW from NE of the Wahanga dome and taken by Lloyd Homer, GNS Science.
128
B.M. Kennedy et al. / Earth and Planetary Science Letters 299 (2010) 126–137
Table 1
Results of high temperature experiments performed on natural samples summarized as experimental conditions, and properties of starting materials and run products. The number
in brackets in the FTIR column represents the number of measurements the mean value is based upon.
Sample name
Time
(t)
T/Tg
(h)
Original g
Original h
Original i
Original j
mean original
BMK06T300
BKM06T400 2
BKM06T500 2
BKM06T600 2
BKM06T700 2
BK06T800 0.5
BK06T800 2
BK06T800 6
BK06T800 17
BK06T800 24
BK06T800 72
BK06T900 0.5
BK06T900 2
BK06T900 4
BK06T900 6
BK06T900 17
BK06T900 24
BK06T900 72
BK06T1000 0.5
BK06T1000 2
BK06T1000 4
BK06T1000 6
BK06T1000 17
BK06T1000 24
BK06T1100 0.5
BK06T1100 2
BK06T1100 4
BK06T1100 6
BK06T1100 17
BK06T1100 24
BK06T1200 0.5
BK06T1200 24
Minimum
Maximum
2
2
2
2
2
0.5
2
6
17
24
72
0.5
2
4
6
17
24
72
0.5
2
4
6
17
24
0.5
2
4
6
17
24
0.5
24
0.5
72
0.41
0.54
0.68
0.81
0.95
1.08
1.08
1.08
1.08
1.08
1.08
1.22
1.22
1.22
1.22
1.22
1.22
1.22
1.35
1.35
1.35
1.35
1.35
1.35
1.49
1.49
1.49
1.49
1.49
1.49
1.62
1.62
Temp.
(T)
Relax.
time
(oC)
(h)
(frac)
(frac)
(frac)
(wt.%)
FTIR (4520 +
5230
peaks)
total H2O
(wt.%)
20
0.28
0.00
0.28
1.62
0.81 (4)
20
20
20
20
300
400
500
600
700
800
800
800
800
800
800
900
900
900
900
900
900
900
1000
1000
1000
1000
1000
1000
1100
1100
1100
1100
1100
1100
1200
1200
300
1200
0.27
0.21
0.25
0.25
0.24
0.23
0.16
0.21
0.21
0.27
0.28
0.39
0.29
0.28
0.32
0.32
0.31
0.36
0.35
0.37
0.33
0.33
0.34
0.46
0.40
0.41
0.43
0.37
0.43
0.38
0.32
0.28
0.28
0.23
0.34
0.04
0.08
0.04
0.04
0.07
0.05
0.12
0.10
0.08
0.04
0.03
0.05
0.02
0.04
0.03
0.06
0.03
0.03
0.06
0.03
0.06
0.05
0.11
0.05
0.05
0.08
0.04
0.08
0.05
0.04
0.04
0.10
0.04
0.09
0.06
0.30
0.29
0.29
0.29
0.31
0.28
0.28
0.31
0.29
0.31
0.31
0.45
0.32
0.32
0.34
0.38
0.34
0.39
0.41
0.40
0.39
0.38
0.45
0.51
0.45
0.50
0.47
0.45
0.49
0.42
0.36
0.37
0.32
0.32
0.40
1.72
1.61
1.69
1.660
0.76
0.43
0.27
0.22
0.2
0.39
0.3
0.16
0.46
0.00
0.12
0.28
0.51
0.20
1.72
30
30
30
30
30
30
1
1
1
1
1
1
1
0.08
0.08
0.08
0.08
0.08
0.08
0
0
0
0
0
0
0
0
1.4 × 106
t/τr
0.5
15
60
180
510
720
2160
476
1905
3810
5714
1.7 × 104
2.3 × 104
6.8 × 104
1.5 × 104
6 × 104
1.2 × 105
1.8 × 105
5.1 × 105
7.2 × 105
1.9 × 105
7.5 × 105
1.5 × 106
2.3 × 106
6.3 × 106
9.0 × 106
1.5 × 106
7.2 × 107
log
visc.
Connected
porosity
Isolated
porosity
Total
porosity
Leco
indution
H2O
(Pa s)
12.4
10.3
10.3
10.3
10.3
10.3
10.3
8.8
8.8
8.8
8.8
8.8
8.8
8.8
7.3
7.3
7.3
7.3
7.3
7.3
6.2
6.2
6.2
6.2
6.2
6.2
5.3
5.3
(30 min–72 h). This timescale was constrained by (1) the 5 h eruption
duration (Keam, 1988) and (2) a total magma rise time of a ~1 day
(Houghton et al., 2004). We characterized time and temperature
dependent changes in texture and water content between 800 and
1200 °C. For temperatures 300–700 °C, experiments were run for 2 h to
characterize temperature dependent degassing below the calculated
theoretical glass transition (Giordano et al., 2008). Immediately
following each experiment samples were removed, cooled in air,
reweighed and their porosities remeasured with a helium pycnometer.
2.3. S.E.M.
Natural samples and samples from each experiment were coated
in carbon and examined using the Philips XL30 electron microscope
(SEM) (Beam power 15 kV and setting “Spot 6”) at the Earth and
Ocean Sciences Department of UBC. We compare the pre- and postexperimental textures of the samples. When imaging both natural and
cored samples we avoided surfaces affected by the rock saw and corer.
2.4. Water content
We ground and sieved samples, including crystals, to 100 μm for
bulk water analysis from experiments lasting 30 min, 2, 4, 17.5 and
FTIR
(3570
peak)
total H2O
(wt.%)
FTIR
(4520
peak)
OH−
(wt.%)
FTIR
(5230
peak)
mol. H2O
(wt.%)
FTIR
image
mol.
H2O
(wt.%)
0.78 (2)
0.09
(4)
0.73 (4)
0.16–
1.30
0.058 (1)
0.32
0.33
0.41
0.35
0.29
0.26
0.3
0.37
0.33
0.26
0.044 (2)
0.32
0.31
0.29
0.28
0.29
0.012 (4)
0.26
0.25
0.29
36 h. These samples were analyzed at ALS Actlabs using a Leco
induction furnace combined with spectral analysis of the emitted
volatiles. The sample (~ 0.3 g) is thermally decomposed in a resistance
furnace (ELTRA CW-800) in a pure nitrogen environment at 1000 °C,
causing release of volatiles, including both H2O− and H2O+.
Some natural and post-experimental samples were thinned and
polished on both sides to make wafers suitable for analysis of volatiles
with the FTIR. FTIR analysis was undertaken at the University of
Oregon following the technique described in Wright et al. (2007).
During analysis, we made every attempt to avoid areas containing
crystals and vesicles. Total H2O contents and the speciation of H2O in
the glass were determined using the absorbances of combination
bands at 5230 cm− 1 (molecular H2O) and 4520 cm− 1 (OH−). At low
total H2O contents (b0.5 wt.%; Stolper, 1982), molecular H2O is not
detectable, in which case total water content was measured using the
3570 cm− 1 band, representing the fundamental –OH stretching
vibration. Absorbances were converted to concentrations using the
Beer–Lambert law and absorption coefficients at 5230 cm− 1 and
4520 cm− 1 from Zhang et al. (1997) and at 3570 cm− 1 from Stolper
(1982). A glass density of 2.34 g/cm3 was obtained from pycnometry.
Sample thickness was determined using a micrometer and varied
between 300 and 900 μm. Results from the 3570 cm− 1 band and
totals from the 5230 cm− 1 and 4520 cm− 1 bands (Table 1) agree to
B.M. Kennedy et al. / Earth and Planetary Science Letters 299 (2010) 126–137
within 0.004 wt.%. Multiple spots were analyzed where possible and
the number of spots per measurement is shown in Table 1 using
number in brackets after the mean measurement.
Additionally, we conducted FTIR spectroscopic imaging of a
fragment of the original dome material to investigate molecular H2O
distribution. This fragment broke naturally along perlitic cracks
during further thinning of the samples, which was carried out to
reduce the number of crystals and vesicles in the wafers. Images (each
350 × 350 μm) of the fragment were collected using a Varian Inc.
Lancer Focal Plane Array (FPA) camera attached to a Varian FTS
Stingray 700 Micro Image Analyser spectrometer and UMA 600
microscope at the Institute for Research on Earth Evolution (IFREE),
Japan Agency for Marine Earth Science and Technology (JAMSTEC).
For more detailed discussion of FTIR spectroscopic imaging see
Wysoczanski and Tani (2006). The thickness of the fragment was
determined using reflective light spectra and the wavelength of the
resulting interference fringes following the method of Wysoczanski
and Tani (2006) and Nichols and Wysoczanski (2007). A refractive
index of 1.50 was used for rhyolite (Liu et al., 2005; Long and
Friedman, 1968). From 89 reflection spectra across the image, the
average thickness was 62 μm (1σ = 1 μm), and this value was used for
all calculations. The glass density used was the same as for the spot
analyses. Variations in glass density as a result of variations in H2O
concentration result in a maximum error of ±0.01 wt.% on the
molecular H2O contents. Owing to the thinness of the sample, the
combination bands at 5230 and 4520 cm− 1 used to measure
molecular H2O and OH− were below detection (see Supplementary
Figure A1). As a result, the spectroscopic image for molecular H2O
represents the absorbance of the fundamental bending of molecular
H2O at ~ 1630 cm− 1. See Supplementary Figure A1 for additional
discussion of the use of this band. Total H2O contents were obtained
from the peak at 3570 cm− 1 as described earlier.
3. Description of natural samples
Our 30 natural samples vary in relative proportions of rhyolite and
basalt (Fig. 2). In contrast to Rosseel et al., 2006, we have chosen to
classify samples on the basis of vesicularity rather than bomb type.
The samples described subsequently are all from the AD1314 lavas or
dykes; wall rocks from older Tarawera lavas (Carey et al., 2007) were
not sampled. These young lithics can be split into four vesicularity
types. Type 1 are glassy samples of the Tarawera lava dome that we
use in the experiments (Fig. 2a). Types 2–4 are samples from the 1886
plinian basaltic fall deposit and contain both rhyolite from the AD
1314 Tarawera lava dome and basalt from the 1886 eruption. The
samples vary from rhyolite “lithics” coated in basaltic scoria to
completely remelted rhyolite enclaves entirely contained within
basaltic spatter (Rosseel et al., 2006) (Fig. 2b–d).
Type 1 samples are representative of the Tarawera lava dome; these
samples were collected from outcrops within the 1886 crater within the
Tarawera lava dome (Fig. 1). Samples are variably devitrified and
texturally perlitic and show a range of vesicularities. Vesicles show a
range of vesicle sizes and shapes (Fig. 2a). We chose a typical large glassy
block without obvious devitrification but with perlitic cracks (Fig. 2a)
for our experiments. All the drilled cores are from the one sample and
show a total porosity variation of 28–30 vol.%. The dome is 15–35%
crystalline (Cole, 1970), and contains phenocrysts of plagioclase, quartz,
biotite and minor amphibole, orthopyroxene and Fe–Ti oxides. SEM
images show curvilinear perlitic microcracks at 50 μm intervals in the
glass (Fig. 2a(ii)) and unknown secondary minerals on glass surfaces.
Water content of the starting material measured by FTIR is 0.8 wt.% and
by Leco induction is 1.6 wt.% (see Fig. 5b).
Type 2 samples were collected from the basaltic fall deposits
draping the Tarawera lava dome, exposed on the margin of the crater.
Samples are 1–10 cm chunks of vesicular rhyolite (Fig. 2c), partially
coated in (b10% by volume) basaltic scoria (b1 cm clast diameter).
129
The basaltic scoria is weakly attached to the surface of the rhyolite.
Where basaltic scoria is absent, cracks are visible on the surfaces of the
rhyolite. The rhyolite has total porosity estimated to range between
40 and 60% (using proportion comparison charts), only one sample
was measured at 53% (Table 1). SEM images show almost no
microcracks and complex vesicle shapes with evidence of vesicle
wall retraction and vesicle collapse (Fig. 2c (ii)).
Type 3 samples were collected from the same basaltic fall deposits
as Type 2 and the talus slopes below this fall deposit. Pyroclasts
contain 10–90 vol.% basalt. Samples are up to 30 cm in diameter, and
visibly very vesicular N60%, one sample was measured at 73% total
porosity with individual vesicles up to 5 mm. The surfaces of these
samples are covered in basaltic scoria and spatter which is strongly
attached to the surface of the rhyolite (Fig. 2c). Adjacent to the basaltic
spatter, the outer few mm of the rhyolite is vesicle poor and appears
locally as black obsidian (Fig. 2c). SEM images show no microcracks
and in contrast to the Type 1 dome rock, the vesicles are spherical
(Fig. 2c (ii)).
Type 4 samples were collected from spatter-rich areas in the
basaltic fall. These pyroclasts are N90% basalt (Fig. 2d). These samples
contain small (b5 cm) dense enclaves of rhyolite within basaltic
spatter. The rhyolite blebs are generally too small to measure
porosities, however, we estimate the total porosity is consistently
b25%. SEM images show a range of textures, with prominent small
spherical and irregular vesicles (Fig. 2d (ii)).
4. Experimental results
4.1. Qualitative results: SEM
We present the results of the experimental heating of 32 Type 1
lava dome cores with the aim of understanding the potential
consequences of reheating conduit filling material. SEM images
showing time- and temperature-dependent changes in the microstructure of the samples are shown in Figure 3. Generally, the textures
progress from cracked vesicular glassy original (Type 1) textures, to
an increase in vesicles with cracked walls, to a connected network of
un-cracked vesicles, and finally to less vesicular isolated vesicles.
Experiments lasting 2 h at 300–700 °C show no noticeable difference
in texture from the original samples. However, the time series from
experiments at 800–1200 °C show remarkable differences in the
geometry and size of cracks and vesicles relative to the initial state.
At 800 °C, samples heated for ~6 h are characterized by open cracks.
In addition, the widest cracks are spatially associated with dome
structures (related to vesicle growth) (Mungall et al., 1996) on the
surface of the glass which are 100 μm in diameter (Fig. 3a). In samples
heated for 24 h cracks are still apparent and some cracks appear more
curved and crack edges appear slightly rounded (Fig. 3b).
At 900 °C, samples heated for ~6 h (Fig. 3c) are texturally similar to
the 800 °C runs (Fig. 3a). However, at 24 h, cracks are commonly
curved. The dome structures appear to have surfaces made of a thin
film of glass and relict cracks and rafts of original surfaces exist on
some domal structures (vesicles) (Fig. 3d), suggesting that films of
melt have allowed vesicles to grow.
In experiments conducted at 1000 °C, the edges of cracks are
distinctly rounded indicating (re-)melting (Fig. 3e). In the sample
heated for 24 h at 1000 °C most of the glass is smooth with angular
cracks only occurring in crystals. In addition, completely smooth
domes occur with only a vague hint of relict cracks (Fig. 3f). The
domes appear as connected vesicles, however these vesicles are not
bulbous, and some appear deflated (Fig. 3f).
At 1100 °C, a sample heated for 40 min shows similar textures to
the sample heated for 24 h at 1000 °C (Fig. 3f and g). The surface of the
sample is smooth with bulbous vesicle-like domes, and no relict
cracks (Fig. 3g). The surface of the sample heated for 6 h at 1100 °C is
completely smooth and cracks are no longer visible in the glass. Most
130
B.M. Kennedy et al. / Earth and Planetary Science Letters 299 (2010) 126–137
Fig. 2. (i) Photographs and (ii) corresponding scanning electron micrographs of rhyolitic samples from Tarawera illustrating 4 main types of vesicularity. (a) Type 1: (i) vesicularity in
vesicular lava dome used as the starting material for our experiments; (ii) S.E.M. image showing curvilinear micro cracks and variety of vesicle sizes and shapes. (b) Type 2: (i) vesicularity
within a rhyolitic pyroclast partially coated in basaltic scoria; (ii) S.E.M. image showing smooth surfaces, retracting vesicle walls, large connected vesicles and no cracks. (c) Type 3: (i) high
vesicularity in a rhyolitic pyroclast coated in basaltic spatter, in some areas 1–3 mm thick obsidian occurs at the boundary between rhyolite and basalt; (ii) S.E.M. image showing large
connected spherical vesicles and no cracks. (d) Type 4: (i) vesicularity in basaltic spatter containing dense enclaves of rhyolite; (ii) S.E.M. image showing a dense rhyolitic enclave with
some isolated irregular vesicles towards its margin.
of the glass is bulbous and appears to be connected vesicles (Fig. 3g).
The glass in the sample heated to 1100 °C for 24 h also is smooth and
has no cracks. Vesicles are not as apparent on the surface of the
sample (Fig. 3h).
Finally, the sample heated for 30 min for at 1200 °C (Fig. 3i) shows
similar textures to the sample heated for 6 h at 1100 °C (Fig. 3g). Glass
is smooth and unfractured and contains connected vesicles (Fig. 3i).
The sample heated at 1200 °C for 24 h has smooth surfaces similar to
that heated to 1100 °C (Fig. 3h), however fresh surfaces show some
isolated vesicles exist beneath the smooth surface (Fig. 3j). All original
textures were obliterated.
In summary, experiments at lower temperatures and/or for short
times generally produced samples with open cracks. Experiments at
higher temperatures and/or for longer times show evidence of
melting, crack annealing, the growth of films of melt, and the inflation
and deflation of vesicles.
B.M. Kennedy et al. / Earth and Planetary Science Letters 299 (2010) 126–137
131
Fig. 3. SEM images of experimental run products produced by heating of natural samples over fixed periods of time. a) Experimental run product for 800 °C over 6 h is largely
texturally unchanged from the original. However, cracks have begun to open especially in areas around domed glass associated with subsurface vesicles. b) Run product for 800 °C
over 24 h shows similar cracking associated with sub-surface vesicles; also shows convex areas and some cracks appear to be rounded. c) Run product for 900 °C over 6 h shows open
cracks associated with sub-surface vesicles. d) Heating at 900 °C for 24 h causes formation of new vesicles that feature films of melt that heal old fractures and create rafts of the old
vesicle wall on the surface of the new vesicle. e) Sample heated at 1000 °C for 6 h shows angular fracture edges rounding and annealing. g) Sample heated at 1000 °C for 24 h shows
almost no vestiges of the earlier cracked surfaces, however, existing dome-like structures appear deflated. g) Samples heated at 1100 °C for 6 h feature connected inflated vesicles
within the original vesicle wall septa, and no evidence of the original cracked glass. h) Heating at 1100 °C for 24 h produces a sample on which the surface shows no evidence of the
original bubbly and cracked surface. i) Run product for 1200 °C over 30 min is similar to g), vesicle walls contain inflated connected vesicles. j) Heating at 1200 °C for 24 h causes the
sample to collapse and flow; material had to be chipped out of the crucible and internally it shows spherical isolated vesicles.
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B.M. Kennedy et al. / Earth and Planetary Science Letters 299 (2010) 126–137
4.2. Porosity
4.4. Results summary
Time series of fractional porosity evolution during the experiments
are given in Table 1 and illustrated in Figure 4. Both total and
connected porosity initially increase with time at all experimental
temperatures. The time taken to reach a maximum in total porosity
decreases as the experimental temperature increases. This increase in
porosity is then followed by a reduction in porosity, where the specific
total vs. connected porosity path is dependent upon temperature.
Isolated porosity is illustrated in Figure 4c by the vertical distance
between any point and the dashed, fully connected porosity line. At
b700 °C, overall porosity and volume were indistinguishable within
error from the unheated samples (Tables 1 and Supplementary Table
A2).
At 800 °C the connected porosity reaches a maximum of 0.39 in the
experiment conducted for 6 h (Fig. 4a). Experimental times greater
than 6 h correspond with lower connected and total porosities.
Isolated porosity remains less than 0.05 in all experiments at 800 °C.
The time series at 900 °C also shows an initial increase in
connected and overall porosity, connected porosity then fluctuates
and reaches a maximum at 17.5 h of 0.37. Thereafter the connected
porosity drops monotonically to 0.33 at 36 h (Fig. 4b). Again, the
isolated porosity remains similar during this time series (Table 1).
The experiments at 1000 °C, have a maximum total and connected
porosity of 0.46 at 2 h. Longer experiments correspond with smaller
porosities (Fig. 4c). This time series shows higher isolated porosities
(up to 0.11) compared to 800 and 900 °C, particularly the experiments
for 2, 6 and 24 h.
At 1100 °C a maximum porosity of 0.43 occurs at 30 min.
Thereafter, the connected porosity drops sharply to 0.23 by 24 h.
Despite the consistent decrease in connected and overall porosity,
isolated porosity varies considerably and shows a substantial peak
(~0.1) at 6 h (Fig. 4d).
Samples heated at 1200 °C collapsed and lost their cylindrical form
(textural collapse), the sample heated for 30 min remained generally
cylindrical whilst the sample heated for 24 h collapsed completely.
In summary, at all experimental temperatures an initial increase in
total porosity is followed by a decrease.
Our observations and measurements of the rhyolite lava prior to
experimentation illustrate that the lava has an intricate network of
perlitic cracks (Fig. 2a) and connected vesicles (connected porosity
0.26 and isolated porosity 0.04). In addition, a significant portion of
the water contents of the original dome glass are not accounted for by
the water species measured by either spot or FTIR mapping. Water
contents that are effectively mapped by FTIR are shown to be strongly
enriched along the margins of the perlitic cracks (Fig. 5a). Our
experimental results show that our samples, where heated above
500 °C lose most of their water within 2 h. However, changes in total
porosity do not occur until 800 °C. All time series above 800 °C
showed a small but continued degassing and an initial increase in
connected porosity associated with cracking and vesicle growth,
followed by a decrease in porosity associated with vesicle collapse.
Generally, the hotter the experiment the earlier in the time series the
porosity decrease occurred. Textural changes are summarized in
Fig. 6, experiments are subdivided into samples with (1) no textural
changes, (2) cracking/ inflation, (3) vesicle collapse/ deflation, and (4)
textural obliteration. This subdivision is used to develop degassing
and deformation regimes which are discussed below.
4.3. Water content
The water content of the glass in the original sample (starting
material) varies between 1.6 and 0.8 wt.% as measured by Leco
induction furnace and by FTIR spot analyses (N = 4), respectively
(Fig. 5b, Table 1). These differences in measured water contents
between Leco Induction furnace and FTIR are absent at higher
temperatures and lower water contents. The FTIR spectroscopic
image of the absorbance for the molecular H2O band (Fig. 5a)
supports this, with absorbance increasing, by up to a factor of 7
compared to the interior of the fragment, at the perlitic margins.
Notwithstanding the qualifying statements of Zhang et al. (1997)
(described in the figure caption of the Supplementary Fig. A1) this
represents an increase in molecular H2O content from approximately
0.16 to 1.30 wt.%. All time series show that N75% of bulk water
measured by Leco induction was lost within 0.5 h. Over 2 h at 500 °C
the bulk water content shows a systematic decrease as temperature
increases from 1.6 wt.% to around 0.22 wt.%, which is close to the
detection limit of this instrument (Fig. 5b). Within the uncertainty of
the measurements, repeated experiments at temperatures above
500 °C showed no change in bulk water content. FTIR spot measurements using the 3550 cm− 1 band support the same general pattern of
rapid water loss and show a mean initial water content of 0.78 wt.%
was reduced to 0.06% after 6 h at 800 °C (Fig. 5b). Additionally, these
FTIR spot analysis support continued degassing to 0.01 wt.% after 6 h
at 1100 °C.
5. Discussion
5.1. Comparison between natural and experimental samples
By comparing the textures of the naturally heated samples and the
experimentally heated samples we constrain the porosity and, by
inference, the permeability history of the dome and conduit wall rocks
in response to reheating. In general, there is a striking similarity
between the textures of type 2–4 samples and the higher temperature
and long timescale experiments we performed (Figs. 2 and 3). The
textural similarities support our estimates of the timescales and
temperatures of reheating that were derived from the eruption
timescale (Keam, 1988). However, we do not imply that the thermally
driven textural changes in the natural samples occurred in-situ at the
conduit margin; most of these erupted as bombs and probably
followed the thermal history described in Rosseel et al. (2006). The
open cracks and cracked vesicle walls of lower temperature and
shorter timescale experiments were not seen in our suite of naturally
heated samples, although this may be due to sampling bias.
The type 2 samples (Fig. 2b) have textures and porosity similar to
experiments for 24 h at 1000 °C (Fig. 3f). In both these rocks, glass
surfaces are smooth, cracks have healed but vesicles remain small and
only partially inflated.
The type 3 samples (Fig. 2c), appear similar to experiments above
1100 °C (Fig. 3g and h) and show no evidence of the original textures
of the type 1 rock (Fig. 2a). However, the experiments do not exhibit
the large vesicle sizes and high porosities of the natural type 3
samples (Fig. 2d). In our experiments above 1100 °C, at timescales
longer than 40 min, vesicles coalesced, connected with the ambient
pressure and collapsed. The large size of the natural samples allowed
more vesicle growth and coalescence, and the developments of a
larger isolated porosity before depressurizing to ambient pressure.
Such an internal pressure could prevent collapse of the sample and
loss of porosity (Yoshimura and Nakamura, 2008). Alternatively,
vesiculation of these samples could have been aided by decompression as they erupted.
The type 4 samples are similar to experiments at 1200 °C for
timescales greater than 30 min, no original vesicle textures resembling the type 1 sample remain. The amoeboid outer shapes and dense
textures of the type 4 samples imply that they fully melted. Vesicles
are generally small, implying that larger vesicles coalesced, collapsed
or escaped by migration through the melt.
B.M. Kennedy et al. / Earth and Planetary Science Letters 299 (2010) 126–137
133
Fig. 4. Experimental results summarized as total porosity vs. connected porosity and labelled by experimental times. Open headed arrows follow a temporal evolution of increasing
experimental time. A dotted 1:1 line indicating equality between connected and total porosity is shown on all plots; the vertical distance between each data point and this line
measures the isolated porosity within each sample. a) Time series at 800 °C; inset of solid headed arrows shows how porosity is affected by various relevant processes. b) Time series
at 900 °C. c) Time series at 1000 °C. d) Time series at 1100 °C.
The similarity of textures shared by the experimental and natural
rocks implies that reheating drove vesiculation and vesicle collapse in
the natural samples. Natural type 2 samples found in the basaltic
scoria reached temperatures of 1000 °C for at least 24 h, whereas type
3 and 4 samples found in spatter imply higher temperatures and/or
longer hot residence times (Carey et al., 2008; Rosseel et al., 2006;
Sable et al., 2009). In summary, the comparison of Figures 2 and 3
show that the temperatures and timescales of reheating during our
experiments were appropriate for the eruption, and that large
variations in conduit wall porosity occurred during the eruption. An
ongoing study of the scoria and spatter filled basaltic dyke margins
exposed in the base of the fissure support this porosity variation.
Additionally, this fieldwork constrains the thermal impact of the
basalt to be between a few millimetres to tens of centimetres, in some
areas we identified a distinct 5 cm thick welded portion; and these
detailed field descriptions will be the subject of a future publication.
5.2. Degassing, and deformation regimes
Degassing and deformation proceeded differently in our experiments relative to other experimental studies involving hydration of
melts (Baker et al., 2006; Gardner, 2007; Larsen et al., 2004; Takeuchi
et al., 2009; Yoshimura and Nakamura, 2008). Our samples were
composed predominantly of glass naturally hydrated by water along
perlitic cracks at temperatures well below Tg (Denton et al., 2009). As
a result this water can be outgassed at temperatures well below Tg
(Tuffen et al., 2010). Isotopic studies indicate that this water is likely
to be meteoric in origin (DeGroat-Nelson et al., 2001; Friedman and
Smith, 1958; Friedman et al., 1966; Shane and Ingram, 2002). An
additional contrast to previous experiments is that our experimental
sample already had a high proportion of connected porosity through
cracks and vesicles. This distinction is important because the
relationship between permeability and porosity is different during
vesicle growth and vesicle collapse (Michaut and Sparks, 2009; Rust
and Cashman, 2004). A sample that has undergone vesicle collapse
will have a higher ratio of permeability to porosity than a sample of
similar porosity that has not undergone vesicle collapse (Michaut and
Sparks, 2009; Mueller et al., 2008). The complexities of this
relationship are most apparent when illustrated by up to 6-fold
increase in permeability over the porosity range of 30–40% (Wright
et al., 2009). The permeability/porosity relationship is further
complicated by the presence of cracks with the ability to open and
to heal (Yoshimura and Nakamura, in press). However, all our
experiments had the same starting texture and we are confident that
incremental changes in open porosity correlate with changes in
permeability.
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B.M. Kennedy et al. / Earth and Planetary Science Letters 299 (2010) 126–137
Fig. 5. a) FTIR spectroscopic images of a fragment of unheated original sample that broke along perlitic cracks: (i) photomicrograph of fragment; (ii) planar spectroscopic map of
absorbance of the molecular H2O band at 1630 cm− 1 (see Supplementary Fig. A1) across the fragment; the fragment is resting on a H2O-free IR-invisible KBr disk (absorbance = 0);
(iii) three-dimensional view of same image indicating that there is some absorbance in the interior of the fragment. Note that one of the edges of the fragment is not enriched in
molecular H2O, suggesting that this edge did not form along a perlitic crack. Colour scale applies to both images. Molecular H2O contents are estimated using a constant molar
absorption coefficient and an average thickness and density for the whole fragment (also refer to Supplementary Fig. A1). b) Weight percent water plotted against temperature for
experiments lasting 2 h. Black squares show water contents of glass measured with FTIR spectroscopy and black circles show bulk water content measured by Leco induction.
Vertical bars show the range in water content from multiple analyses and are due to variation within samples.
On reheating, our experiments imply that the porosity and
permeability structure of lava domes or plugs will pass through a
series of regimes, the nature of which depend on 1) the temperature
of the event; 2) the time spent at that temperature (Fig. 6a); and 3)
the volatile content and original vesicularity of the rock being heated.
Physically, the character of the change in porosity and, by inference
the permeability, depends on whether dome rocks respond in a
viscous or brittle way to heat transfer from the newly erupting
magma. Thus, the key parameter is the glass transition, which
depends on the temperature of the event relative to the glass
transition temperature Tg (Knoche et al., 1994) and the time scale of
the event relative to a critical relaxation time for structural failure of
the melt. Consistent with experimental results, we take this critical
time to be proportional to the viscous relaxation time τr = μ/G (Webb
and Dingwell, 1990) (Table 1). Here, G = 1010 Pa is the elastic
modulus, μ is the melt viscosity at a given temperature and water
content and C is a constant to be determined from our experiments.
We take Tg to be the temperature corresponding to a viscosity of
1011.4 Pa s which correlates well with the calorimetric Tg (Giordano
et al., 2008). The value of Tg is calculated to be 740 °C for this rhyolite
glass composition (Nairn et al., 2004) at 0.1 wt.% H2O using the
viscosity model of Giordano et al. (2008) (see Supplementary Table
A2). The transition from brittle to viscous behaviour (Regimes 1 to 3 in
Fig. 6) depends on the temperature (T) relative to Tg and the time
scale for the experiment (t) (or eruptive event).
To gain additional insight we replot the data in Figure 6a in terms
of T normalized to Tg and t normalized to τr (Fig. 6b). The data collapse
to a power law defining the transition from brittle to viscous
behaviour of the form T / Tg = C(t / τr)0.046, where C ~ 0.86. This result
shows that whether or not brittle processes govern the final porosity
of the sample depends on the response time of the melt, which is
governed by its strongly temperature-dependent viscosity, relative to
the time scale of the thermal forcing applied in the experiment.
We use our experimental degassing and deformation data to
propose four degassing and deformation regimes (Fig. 6). In Regime 1
(T ≤ 800, t ≤ 2 h; T/Tg N Ct/τ0.046
) degassing occurs without significant
r
deformation (Fig. 6). The rock maintains its vesicle structure
connected by microcracks and ruptured vesicle walls (Fig. 2a, b)
and consequently maintains a high proportion of connected to total
porosity (Fig. 4). No visible changes to this structure could be seen
with the SEM. Between 0 and 700 °C we attribute this initial H2O
degassing to (1) release of water captured in micropores and/or low
temperature alteration hydrous minerals (Denton et al., 2009), and
(2) diffusion of resorbed meteoric water out of the perlitic margins
(Fig. 5a) of the glass (Tuffen et al., 2010). 0.20–0.05 wt.% magmatic
water remains dissolved in the glass interior (Fig. 5). This initial
period of degassing had no impact on the textures of the rock as the
rock was not sufficiently above its calculated glass transition
temperature (Giordano et al., 2008) and ductile deformation did not
occur.
B.M. Kennedy et al. / Earth and Planetary Science Letters 299 (2010) 126–137
Fig. 6. a) Sample deformation style separated into Regime 1 no deformation, Regime 2
cracking/ inflation, Regime 3 vesicle collapse/ deflation, and Regime 4 textural
obliteration plotted in time and temperature space. The dashed lines mark approximate
boundaries between regimes. The arrow suggests a heating trajectory based on the
thickness of the remelted dyke margin measured in the field. The curved solid line
represents the calorimetric Tg which increases as water is lost; below this line
conditions are equivalent to Regime 1. The line is calculated approximating time to a
decreasing H2O content (measured by FTIR spectroscopy, Fig 5b) and from the
calorimetric Tg calculated using Giordano et al. (2008; see inset and Supplementary
Table A2). b) T normalized to Tg and t normalized to a critical strain rate for structural
failure of the melt (τr). This plot shows that transitional Regime 2 plots at T/Tg ~ t/τ0.046
,
r
below this critical power law degassing occurs through pre-existing structures and
above power law viscous behaviour allows vesicle collapse.
Regime 2 (800 ≤ T ≤ 1100 °C, t ≤ 6 h; T/Tg ≈ C t/τ0.046
) is a transir
tional regime showing both brittle and viscous deformation (Fig. 6).
Degassing of magmatic water (b0.1 wt.%) continues (Fig. 5) and
vesicle and crack expansion cause 5–20% increases in overall porosity
(Fig. 4). Observations show the porosity changes are due to local
viscous vesicle growth and coalescence with concurrent brittle
cracking in areas of high strain (Fig. 3a–c). Our experimental samples
have initial porosities of 0.29–0.31, and connected porosity of 0.27;
the growth of isolated vesicles frequently impinge on other vesicles
(Fig. 3a–c), resulting in coalescence and increasing connected
porosity with relatively small increases in isolated porosity
(Fig. 4a). In addition, the opening of cracks (Fig. 3b, c) connects
isolated vesicles, increasing the connected porosity and reducing
isolated porosity (Fig. 4a).
In Regime 3 (800 ≤ T ≤ 1200 °C, t ≥ 30 min; T/Tg b Ct/τ0.046
) porosity
r
loss occurs due to vesicle deflation (Fig. 6). This regime is
characterized by crack sealing and vesicle collapse leading to a
transient but generally decreasing overall porosity (Fig. 4). Both
135
processes involve viscous relaxation and appear to occur at similar
temperatures and timescales. Vesicle collapse is driven as the internal
pressure of the vesicle can no longer balance the surface tension and
gas leaks out of the vesicles through the connected porosity. Crack
healing occurs as melt connects opposite sides of a crack (Fig. 3e,f).
Healing of cracks reduces connected porosity and increases isolated
porosity (Fig. 4). However, generally, in our experiments observations
of crack healing correlate with decreases in overall porosity which can
only be explained by vesicle collapse (Figs. 3 and 4). Occasional small
increases in porosity are seen in this regime and we interpret these to
be due to continued degassing and vesicle growth, as larger coalesced
vesicles collapse.
In Regime 4 (T ≥ 1200 °C, t N 30 min; T/Tg N N Ct/τ0.046
) textural
r
collapse occurs due to viscous flow of the melt in response to gravity
(Fig. 6). Although we could not measure the residual porosity of these
run-products, SEM image analysis shows a significantly reduced
porosity comprising isolated pores (Fig. 3k) implying a much reduced
permeabilty.
The temperature/time space of each of these regimes is ultimately
controlled by the viscosity of the melt/glass framework in the dome
lava which is itself strongly influenced by the dissolved water content
(inset Fig. 6a). The calculated dry value for Tg of this melt (see
Supplementary Table A2) is 770 °C. The equilibrium volatile content
of the glass will be controlled by pressure and therefore the
“openness” and porosity of the volcanic system. An open, degassed
system such as the system reproduced by our experiments has a high
Tg of 740 °C (~ 0.1 wt.% H2O) and is relatively difficult to remelt and
initiate Regimes 2, 3 and 4. However, a closed system with a higher
ambient pressure and equilibrium volatile content will have a
depressed Tg. For example, this melt with 1 wt.% H2O has a calculated
Tg of 606 °C (see Supplementary Table A2, Giordano et al., 2008).
Therefore, if such a system is reheated it may achieve Regimes 2, 3,
and 4 at lower temperatures.
The style of hydration, the original texture, and the amount of time
a reheated rock/magma spends in Regime 1 will strongly influence the
relationship between water content and its calorimetric Tg, and ability
to continue to other regimes. Similarly the amount of time a rock
spends in permeable Regimes 1 and 2 will influence the ability of any
magma beneath it to degas. For these reasons heating rates and
partially closed systems become important considerations and
avenues for future research.
6. Conclusions and implications for Tarawera
The frozen magma forming the conduit walls will be reheated
during the ascent and eruption of fresh magma. The time and
temperature dependent textural changes in the plug or walls have
implications for the monitoring of degassing and deformation of
active volcanoes, and on the resultant style and magnitude of an
eruption. In particular, a key issue is whether reheating leads to the
production or destruction of permeability in the conduit walls. For
example, enhanced permeability facilitates outgassing, which can
reduce the overpressure in the intruding magma, and favour effusive
volcanism (e.g., Quane et al., 2009). In contrast, reduced permeability
can inhibit outgassing, leading to greater overpressure in the new
magma and an increased likelihood for explosive volcanism.
From our experiments, the evolution of wall rock texture during an
eruption depends on the temperature and duration of the event (Fig. 6).
In Regime 1 (T/Tg N Ct/τ0.046
), there is degassing of rehydrated meteoric
r
water from the rhyolite. For Tarawera dome rock, up to 1.4 wt.% water
could be released by conduit wall and old dome rocks into the erupting
magma during reheating. In Regime 2 (T/Tg ≈ Ct/τ0.046
) both ductile and
r
brittle deformation cause increases in connected porosity and inflation.
This increase in connected porosity should correlate with an increase in
permeability and hence aid the degassing of the fresh magma below.
Regime 3 (T/Tg b Ct/τ0.046
) is a result of vesicle collapse and porosity
r
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B.M. Kennedy et al. / Earth and Planetary Science Letters 299 (2010) 126–137
reduction. Although a connected pore structure still remains, permeability will be reduced progressively closing the degassing system
(Westrich and Eichelberger, 1994; Yoshimura and Nakamura, 2008;
Yoshimura and Nakamura, in press). Regime 4 produces a generally
isolated pore structure and permeability would be significantly
decreased causing pressure build up, that could give rise to explosive
eruptions.
The magnitude of thermally-induced permeability
change is
pffiffiffiffiffi
expected to depend on the length scale L = 2 κt , where the thermal
diffusivity k of potentially foamed rhyolitic dome rocks is order
107 m2 s− 1 (Bagdassarov and Dingwell, 1994). For Tarawera, over the
proposed 5–24 hour magma rise and eruption time (Houghton et al.,
2004; Keam, 1988) L is likely to be approximately 8–25 cm (Fig. 6a).
Indeed, on-going field studies of the margins of basalt dykes in the
lava domes reveal a 5 cm thick remelted layer of rhyolite with reduced
porosity. This 5 cm thick remelted layer is remarkably consistent with
the eruption timescale and thermal diffusion length scale (Fig. 6a).
However, detailed descriptions of these dyke margins are not yet
complete and beyond the scope of this paper.
For the 1886 eruption, we propose a reheating trajectory of 5 h
(Fig. 6a) which is consistent with (1) the eruption timescale
(Houghton et al., 2004; Keam, 1988), (2) the presence of all three
naturally heated sample types, and (3) the thicknesses of remelted
dyke margins.
The volume of AD 1314 Tarawera rhyolite affected is difficult to
estimate due to uncertainties in the subsurface geometry of these and
older rhyolite lavas and feeder dykes, and their relationship to the
basalt dyke (e.g. Fig. 4 in Carey et al., 2007). We take the basaltic
fissure to be 8 km long (Fig. 1) in map view and 0.2–8 km deep (the
lower bound assumes that the basalt interacts only with the shallow
dome; the upper bound implies interaction with a conduit over the
full depth to the 8 km rhyolitic magma source (Shane et al., 2008)).
Over the 5 hour eruption duration the intruding magma will heat and
degas the surrounding rhyolitic wall rocks to a distance of 8 cm
(Fig. 6a). From these dimensions and assuming that the wall rocks
contain 1.4 wt.% H2O, reheating would release around 6.4 × 106–
2.6 × 108 kg of H2O. From our experiments, the majority of this
volume would be released within the first 2 h of reheating and may
significantly contribute to precursory signs of eruption, although it
may not be significant in terms of the total volatile budget of the
eruption.
The similarity between the textures in the rhyolite erupted from
Tarawera volcano and the experiments defining Regime 4 imply that
conduit-wall permeability may have been catastrophically reduced
during the early stages of the Tarawera eruption. We propose that
such a sealing of the system may have produced a closed system and
aided the unusually explosive basaltic eruption of Tarawera in 1886.
This offers an additional and complimentary explanation to the
driving mechanisms suggested by previous authors e.g., phreatomagmatic interactions and bubble/microlite coupling (Carey et al.,
2007; Sable et al., 2009).
Finally, our work also has implications for eruptions related to reintrusion of similar temperature magma. A closed pressurized lava
dome and conduit with 1–3 wt.% H2O (e.g. Burgisser et al., 2010)
could potentially depress Tg (Fig. 6a inset) and hence allow melting
associated with the re-intrusion of magmas of similar composition
(Giordano et al., 2008), i.e. a 700 °C rhyolite could re-intrude and melt
a water rich rhyolite with a Tg depressed to 600 °C. Our work predicts
that as an old lava dome or plug heats up it will initially inflate,
partially degas, then deflate, and seal up leading to continued
pressurization and possibly eruption (Fig. 6). This sequence may be
followed by another reheating event and lead to the cyclic
deformation observed at many lava domes (Johnson et al., 2008;
Matthews et al., 1997; Voight et al., 1999).
Supplementary materials related to this article can be found online
at doi:10.1016/j.epsl.2010.08.028.
Acknowledgements
Funding for MJ was provided by NSERC, Canadian Institute for
advanced Research, and Marsden (UOC0508). Funding for BK was
provided by Marsden Fast start (09-UO-017C). Help with measuring
water contents was provided by Paul Wallace, and John Stix.
Additional help with field access was provided by Ken Ruaeti and
the Ruawahia 2b Trust and Judy Collins of Mt Tarawera tours, Paul
Ashwell, Felix Von Aulock, and Fabian Wadsworth. We would also like
to thank reviewers Dr. Heather Wright and Dr. Hugh Tuffen for their
detailed and insightful reviews that significantly improved the
manuscript.
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