Degassing processes during lava dome growth: Insights from Santiaguito lava dome, Guatemala

Journal of Volcanology and Geothermal Research 202 (2011) 153–166

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

Degassing processes during lava dome growth: Insights from Santiaguito lava dome, Guatemala A.S. Peter Holland a,⁎, I. Matthew Watson a, Jeremy C. Phillips a, Luca Caricchi a, Marika P. Dalton b a b

School of Earth Sciences, University of Bristol, Bristol, BS8 1RJ, UK Department of Geological and Mining Engineering and Sciences, Michigan Technological University, Houghton, MI 49931–1295, United States

a r t i c l e

i n f o

Article history: Received 21 February 2010 Accepted 5 February 2011 Available online 22 February 2011 Keywords: SO2 UV camera Santiaguito degassing viscosity

a b s t r a c t Eruptions of intermediate magma may be explosive or effusive. The development of open system degassing has been proposed as a pre-requisite for effusion of intermediate magma, however processes leading to open system degassing are poorly understood. To better understand degassing processes during lava dome extrusion we report high temporal-resolution SO2 emission rate measurements collected with an ultra violet imaging camera at Santiaguito, Guatemala. Santiaguito is an ideal case study as the dome lava is compositionally very similar to products of the 1902 Plinian eruption of the parental Santa María volcano. We find that degassing is weak (0.4–1 kg s− 1) but continuous, and explosions are associated with small increases in emission rates (up to 2–3 kg s− 1). Continuous repose degassing occurs through a shallow cap rock which likely represents a proto-crust on the block lava flow which is extruded from the same vent. The continual permeability of the upper conduit argues against a mechanism of explosion triggering in which gas pressure builds beneath a viscous cap rock or plug. Rather, we consider degassing data better consistent with a model of shear-fracturing at the conduit margins. Using field constraints, we model the viscosity of Santiaguito magma as a function of depth and show that conditions for shear-fracturing are met from 150–600 m to the surface. This is in line with independent estimates of explosion initiation depth. We show that repose timescales are orders of magnitude longer than the timescale for shear fracture, and suggest that explosions are triggered when a continuous network of smaller-scale fractures develops, at which point decompression occurs and an explosion is triggered. Fracture healing occurs by viscous relaxation however near to the surface where viscosity is highest, an unconsolidated gouge layer may develop. Our model implies that the observed explosions are a by-product of extrusion. Shear-fracturing can drive open system degassing of crystal rich intermediate magma at shallow levels in the conduit, as high magma viscosity is able to overcome the low strain rates associated with slow ascent of magma. © 2011 Elsevier B.V. All rights reserved.

1. Introduction Volcanoes of intermediate composition (andesite–dacite) are responsible for many of the most explosive and hazardous eruptions recorded, for example Tambora 1815 (Self et al., 1984; Sigurdsson and Carey, 1989), Mount St. Helens 1980 (Lipman and Donal, 1981) and Pinatubo 1991 (Newhall and Punongbayan, 1996). However, intermediate magmas may also erupt effusively as lava domes and flows, for example Santiaguito 1922–present (Rose, 1972; Harris et al., 2003) and Mount St. Helens 2004–2008 (Sherrod et al., 2008). Understanding the factors which control whether intermediate magma will erupt explosively or effusively has significant implications for hazard assessment and remains a major goal of volcanology. Gas dissolved in magma at depth is the driving force of most explosive eruptions, but it has been shown that eruptions of varying

⁎ Corresponding author. E-mail address: [email protected] (A.S.P. Holland). 0377-0273/$ – see front matter © 2011 Elsevier B.V. All rights reserved. doi:10.1016/j.jvolgeores.2011.02.004

explosivity, as well as explosive and effusive phases of single eruptions, can be sourced from magma with similar initial volatile contents (e.g. Hervig et al., 1989; Barclay et al., 1996; Martel et al., 1998). Explosivity is instead likely to be determined by degassing processes occurring during rise through the crust from magma chamber to surface (e.g. Eichelberger et al., 1986; Newman et al., 1988; Jaupart and Allègre, 1991; Jaupart, 1998; Gonnermann and Manga, 2007; Castro and Gardner, 2008). During magma ascent the overburden pressure reduces and volatile solubility drops, causing volatiles to exsolve into a discreet gas phase. In intermediate magmas, viscous forces dominate the buoyant forces acting on bubbles and they are trapped in place, leading to coupled rise and closed system degassing (Sparks, 2003). During closed system degassing, gas bubbles grow as the pressure decreases, increasing the flow volume and hence upwards velocity, eventually leading to fragmentation and explosive eruption (e.g. Cashman et al., 2000). The emplacement of lava domes and flows thus requires the development of open system degassing, in which gas is able to escape the magma reducing gas overpressure (e.g. Gonnermann and Manga, 2007). However, processes which lead

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to open system degassing in intermediate magmas are not well understood, and may include the development of permeable foams (Eichelberger et al., 1986; Jaupart and Allègre, 1991; Eichelberger, 1995; Jaupart, 1998) or fracture networks through strain-induced shear fracturing (Gonnermann and Manga, 2003; Tuffen et al., 2003; Collier and Neuberg, 2006; Neuberg et al., 2006). The Santa María–Santiaguito volcanic complex (Fig. 1), Guatemala, provides an ideal location to investigate processes of degassing during extrusion of an intermediate lava dome. In 1902, Santa María was the source of one of the largest Plinian eruptions on historical record, ejecting at least 8.3 km3 of dacitic magma into the atmosphere (Williams and Self, 1983; Rose, 1987a). Twenty years later, extrusion of chemically similar dacitic magma began inside the eruption crater, representing the first growth of the Santiaguito dome complex which continues to grow today and currently has a total volume of N1.1 km3 (Rose, 1973, 1987a; Harris et al., 2003). From 1977 to the present, activity at Santiaguito lava dome has been characterised by continued extrusion of dacite block lava flows accompanied by frequent (0.5–2 h− 1) low-intensity explosions, producing ash and gas plumes to heights of 0.5–2 km above the dome (Rose, 1987b; Harris et al., 2003; Johnson et al., 2004; Sahetapy-Engel et al., 2008). The mechanism of these explosions is in itself a topic of considerable debate (e.g. Bluth and Rose, 2004; Johnson et al., 2008; SahetapyEngel et al., 2008; Johnson et al., 2009; Sahetapy-Engel and Harris, 2009), however understanding the provenance of these explosions has the broader context of understanding open system degassing processes during a period of low-explosivity. In order to make direct measurements of degassing processes we collect spectroscopic SO2 emission rate data. Studies of SO2 emissions have proven useful in outlining processes occurring at shallow levels of magmatic systems (e.g. Watson et al., 2000; Fischer et al., 2002; Burton et al., 2009; Mori and Burton, 2009; Boichu et al., 2010), however previous efforts to study SO2 emissions at Santiaguito (Andres et al., 1993; Rodríguez et al., 2004) have been limited by the low temporal resolution of standard spectroscopic techniques, precluding an in-depth look at gas emissions across the short eruptive cycle. In this study we use ultra violet imaging camera (UV camera) technology (Mori and Burton, 2006, 2009; Bluth et al., 2007; Dalton et al., 2009; Kantzas et al., 2010; Kern et al., 2010b) to collect a high temporal resolution SO2 emission dataset. This allows us to study the degassing processes occurring across the eruptive cycle in unprecedented detail, facilitating our dual aims of exploring the mechanics of low intensity explosive activity and understanding degassing processes during dome growth. In analysing the results, we combine

degassing observations with a simple viscosity model and suggest that the explosions are caused by shear-fracturing at the conduit margins, and that this is an important process for promoting open system degassing during slow extrusion of crystal-rich intermediate magma. 2. Current activity During field campaigns in January/February of 2008 and 2009 we observed short duration (5–10 min) explosive events, which produced plumes with a small gas thrust region. Deceleration to buoyant rise occurs over the first 50 m above the edifice. Plumes typically attain heights of 300–1500 m above the dome. Classification of these explosions is difficult, as reflected by the variable terminology in the literature (Bluth and Rose, 2004; Johnson et al., 2004; Sahetapy-Engel et al., 2008; Marchetti et al., 2009). To avoid mechanistic implications associated with previous classifications, we use the term “explosion” to refer explicitly and exclusively to this form of activity throughout this paper. During short (30–120 min) repose periods, the gas plume is rarely visible to the naked eye, suggesting low levels of water vapour condensation in the plume. A block lava flow extrudes from the west side of the dome surface and regular rock avalanches occur from this flow. In 2008 we observed small rockfalls originating from the summit of the dome, as well as rockfall from an unstable section of the 1902 eruption scar on Santa María. In 2009 rockfall events were more frequent and originated primarily from the south-west rim of the dome surface. During the two week deployment in 2009, 3 small pyroclastic flows were observed on the east flank of the active Caliente Dome, coincident with vertical explosion plumes. 3. Data acquisition and processing The UV camera (Mori and Burton, 2006; Bluth et al., 2007) consists of an Apogee Instruments E6 Alta digital camera and a 105 mm focal length Coastal Optics lens (field of view = 13°). The camera sensor is a 1024 × 1024 pixel 16-bit Kodak KAF-1001E-2 CCD with a photo active area of 24.5 × 24.5 mm and a quantum efficiency of approximately 0.1 at 307 nm. This sensor was chosen as the best compromise between cost and quality, and as it has a known quantum efficiency at the wavelength of interest. The camera was fitted with a single Andover Optics UV band-pass filter centred at 307 nm with a 12 nm full-width half-max bandwidth. At typical exposure settings the instrument collects images at around 0.33 Hz.

Fig. 1. View of Santa María (right) and the active Caliente dome of Santiaguito, seen looking NNE from the imaging position at OVSAN, with observed activity marked schematically. Altitude difference between the summit of Caliente and that of Santa María is 1220 m.

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SO2 measurements were collected from the Santiaguito Volcano Observatory (OVSAN), 6.3 km SSW of the active Caliente dome. From this position the field of view at the distance of the plume is 1495 × 1495 m, with a single pixel resolution of 1.46 × 1.46 m. We collected visible light video footage from the same vantage point using a Sony Handycam to constrain visible activity. Data collection was only possible during daylight hours and when Santiaguito was not obscured by thick orographic cloud. This typically limits data collection to continuous periods of only a couple of hours per day. Raw UV camera images are processed following the protocols outlined in Holland et al. (submitted for publication). Raw images are converted from images of incident light intensity to absorbance images by solving the Beer–Lambert law for each pixel (Mori and Burton, 2006; Bluth et al., 2007). Vignetting is removed by use of a model background which is constructed from a series of second-order polynomials, fitted to clear (plume free) sky on either side of the subvertical plumes. A distance correction is applied to correct for light scattered into the instrumental optical path (Mori et al., 2006; Kern et al., 2010a). This is achieved by measuring the radiant signal of a number of ridge lines at different distances from the instrument, allowing us to constrain the rate of signal dilution with distance. Our field viewing geometry did not allow us to collect an image containing the necessary ridge lines for this analysis; however, such images have been collected previously at Santiaguito at a similar time of the year, providing us with a distance correction (βA = 0.182; Bouquet, 2007; Holland et al., submitted for publication). Corrected absorbance is scaled to path-length SO2 concentration (in parts per million metre, ppm m) using calibration curves derived from images of COSPEC-style fused silica calibration cells (Resonance Ltd.) of known SO2 concentration (90 and 270 ppm m in this study). In-plume ash and aerosols interfere with the single filter UV camera retrieval (Mori and Burton, 2006; Kantzas et al., 2010; Kern et al., 2010b). We correct explosive measurements using coincident UV-spectrometer measurements. UV-spectrometer data is processed by standard differential optical absorption spectroscopy (DOAS) methods (e.g. Platt, 1994; Galle et al., 2002), which shows no ash interference, and by a simulated UV camera retrieval. The difference between results from each retrieval reflects the effect of in-plume ash and aerosols (Fig. 2). Due to equipment failure, coincident DOAS data was only collected on a single day in 2009. Horizontal scanning DOAS data from two explosions observed on 9th February are used to develop a correction factor for explosive UV Camera measurements which we apply to the entire dataset. We apply a correction factor equal to the mean result. The correction factor applied to explosive measurements in this paper is 0.38.

SO2 column concentration (ppm m)

500 UV Camera DOAS

400

300

200

Corrected SO2 concentrations are integrated along a horizontal transect line across the UV camera image. The transect is chosen to traverse the entire plume width and is located close to the dome surface. As SO2 values are path-length concentrations this transect represents a plane through which the plume rises. The integrated SO2 concentration is multiplied by the vertical component of the plume velocity to provide an SO2 flux through this plane. Plume rise rates are manually picked by tracking notable features between successive UV images (Bluth et al., 2007; Holland et al., submitted for publication). Features which may be tracked include the plume front and packets of SO2 concentration within the plume. 4. Error analysis A number of errors are introduced during the capture and processing of UV images (Table 1). Model backgrounds were generated for images of clear sky to assess the error involved in predicting the background by the polynomial method, leading to an estimated error of ±3% (see Holland et al., submitted for publication). The distance correction removes a systematic error present in SO2 measurements collected under non-ideal conditions, which may have led to underestimates of N50% in uncorrected measurements for our field geometry (Kern et al., 2010a). However, in this study the temporal variation of this parameter is not considered. Assuming land-use (and hence anthropogenic contribution to atmospheric thickness) is the same for the duration of field work, the major cause of error in the application of this parameter will be humidity (Shettle and Fenn, 1979). The observed extinction coefficient (β) at Santiaguito (0.1468) suggests that the atmosphere is best characterised as approximately “tropospheric” in the classification of Shettle and Fenn (1979). Relative humidity records collected at OVSAN over the period 22nd January– 19th February 2007 and 2008 show that on 95% of days relative humidity is between 66 and 81%. The range of β values implied for tropospheric atmospheres of this humidity suggests that these variations in humidity of will lead to variations in the correction factor of ±14%, and we adopt this as our error estimate. Calibration curves in this study are derived from just two calibration cells (as per standard COSPEC techniques), and multiple calibration runs on the same images provide a spread of gradients. From this spread, the calibration error is estimated at ±4%. The SO2 burden across a transect plane is calculated by integrating across the plume as a function of distance. Pixel dimensions are calculated from trigonometric relationships. GPS measurements provide instrumental position to a high degree of accuracy; however a topographic map is used to locate the volcano. Errors in the derived distance are estimated to be 2% (120 m) in this study incorporating GPS errors, errors in map measurement and differences in the geometry of the dome between the map survey (1985) and the present day. Error is introduced during rise rate determination by manually tracking features due to pixel size calculation and difficulty picking rise rates of turbulent features. An error of ±5% is estimated for this picking difficulty. Additionally, it is not possible to pick rise rates for

Table 1 UV image error analysis. Total error is calculated as the root-mean-square of the contributing errors.

100

0 0

155

20

40

60

Spectrum # Fig. 2. Correcting for in-plume ash and aerosols by comparing a simulated UV camera retrieval with a standard DOAS retrieval. For this horizontal scanning measurement of an eruption plume on 9th February 2009 just 31.9% of the UV camera retrieval signal is due to SO2.

Error Source

Repose

Eruptive

Calibration Modelling background Distance Correction Pixel dimensions Velocity Picks Variable velocity Ash Correction Total RMS Error

4% 3% 14% 2% 5% 10% N/A 18.7%

4% 3% 14% 2% 5% 10% 23% 29.6%

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every frame due to a lack of trackable features, and due to the volume of data. We assume that plume velocity is constant between successive rise rate picks, however this is unlikely to always be true. In addition we assume a uniform rise velocity across the plume cross section. The error associated with these assumptions is estimated at ±10%. An additional source of error is introduced to explosive measurements by the application of the ash correction factor. The correction factor is based on a limited number of DOAS scans from two explosions. Although analysis of raw UV camera time series and visual observations suggest that these explosions can be considered representative, we are unable to correct for variable ash content as a function of time. To account for variable ash content we calculate the standard error of the sample mean, using our limited scans as the sample, and attribute an error of ±23% to this correction (0.38 ± 0.09). Assuming that individual errors are random, independent, and following the root-mean-square method of compounding errors (Stoiber et al., 1983; Williams-Jones et al., 2008), the total methodological error is ±18.7% during repose and ±29.6% during explosions (Table 1). Note that this comprehensive analysis of uncertainty is not directly comparable with many previous SO2 emission rate error analyses, which do not assess radiative transfer effects. 5. Results SO2 measurements presented here were collected on clear, cloudfree days under good conditions for making SO2 emission rate measurements — a total of six days of observation across two field seasons. For clarity, we refer to our two periods of observation as “2008” and “2009” without any inference that the details seen during these short periods of observation are representative of the entire years. In total, 14 796 individual SO2 emission rate measurements were collected covering 13 h 5 min and 47 s of observation. During this observation period 11 complete explosive events are imaged. All times in this paper are in Guatemalan local time (GMT −6). 5.1. Broad patterns of SO2 emission at Santiaguito In data which has not been corrected for the presence of ash, SO2 emission on all days appears highly cyclic. The highest SO2 emission rates are observed during explosions, with significantly lower emissions seen during the intervening periods of repose. However the ash correction makes this cyclic signal significantly less obvious (Fig. 3). The highest SO2 emission rates are still seen during explosions, however at times the magnitude of SO2 degassing during explosions is of comparable magnitude to the strongest periods of repose degassing on the same day (e.g. 29th January 2008). Emission rates regularly peak at ~2–3 kg s− 1 during explosions, although a maximum emission rate of 6.25 kg s− 1 is observed on 12th February 2009. By integration of the time series across complete explosions we calculate that individual explosions release 370–1270 kg of SO2 (Table 2), although most events cluster further between 500– 800 kg total SO2. Repose degassing rates are consistently low: average repose emission rates are between 0.28 kg s− 1 and 1.01 kg s− 1. Although degassing rates can be very low during repose periods (minimum = 0.09 kg s− 1), emissions never drop to zero — repose degassing is continuous. No obvious “sealing” signal is seen over the course of repose periods (c.f. Fischer et al., 2002; Nadeau et al., 2011). 5.2. Observation of different patterns of degassing Overall degassing rates observed during data collection in 2008 are higher than those observed in 2009 (Table 3). As activity at Santiaguito consists of a series of similar explosion and repose cycles, the average SO2 emission rate for a single cycle provides the best

practical estimator for the daily average. The daily emission rate (which, unlike the instantaneous emission rate, is most logically presented in tonnes per day of SO2) varies between 43.07 t d− 1 and 95.49 t d− 1 during this study (Table 3). Cycles recorded during a single day provide similar estimates of the daily emission rate, suggesting that this approach is valid. The average daily SO2 emission rate during our 2008 field season was 84.56 t d− 1compared to 55.38 t d− 1 for observations in 2009. Visual observations suggest that individual explosions were more energetic in 2009 (Johnson et al., 2009). However, with the exception of one explosion on 12th February 2009, there is no notable difference in the range of explosive emission rates or in total explosive emissions between the two periods of data collection. In fact, the average explosive emission rate is nearly identical (±0.01 kg s− 1) for each of the two periods of observation. The different overall rates of degassing observed are best accounted for by changes in the nature, and duration, of repose period degassing. Repose degassing is generally higher in our 2008 dataset (average repose emission rate 0.72 kg s− 1) than in 2009 (0.58 kg s− 1). In addition, repose periods observed during 2009 are longer than in 2008 (3801 and 1784 s respectively; as seen in Fig. 3). Longer repose periods, coupled with explosions of similar magnitude, lead to lower average emission rates. In 2009, repose degassing is consistently low in an apparently stable state. By contrast, the patterns of background degassing during 2008 can be more variable, with pulses of higher degassing interspersed with periods of very low degassing. Thus we observe two subtly different patterns of degassing at Santiaguito during our two field campaigns. While the results presented in this paper cover far too short a period of time for us to understand if this is a long-term change in degassing pattern, the existence of these distinct patterns during similar explosive activity is an important observation (Section 9.1). 5.3. Spatial variability in SO2 emissions Although the UV camera provides a 2D image of a 3D plume, and so does not facilitate the precise location of different sources of emissions, we are able to resolve several distinct sources of SO2 emissions with the UV camera. Changes in SO2 emission sources occur both during periods of repose and during explosions. Fig. 4 shows how the spatial distribution of SO2 emissions develops over the course of a single repose period and explosion — a total time of less than 30 min. During this example cycle at least three different sources of emissions are visible (Fig. 4). New emission sources open and close again over time periods of 5–10 min (Fig. 4). Such spatial variability is seen during both years of observation. 6. Long-term emission-rate trends at Santiaguito The long-term SO2 emission rate, calculated by weighted average of COSPEC scans and traverses collected between 1976 and 1991, was found to be 80 t d− 1 (Andres et al., 1993). The average emission rate calculated for 2001 and 2002 was 120 t d− 1(Rodríguez et al., 2004), an increase interpreted as the result of either a decrease in SiO2 content in Santiaguito magma, or high extrusion rates during the period of study. We observed lower daily emission rates of 84.38 t d− 1 and 55.38 t d− 1 during our field campaigns in 2008 and 2009 respectively. Even considering large errors in COSPEC measurements (typically 40%), the daily emission rates in 2009 are statistically different from those recorded in 2001 and 2002. If the SO2 emission rate is primarily dependent upon melt chemistry, the observed decrease in SO2 emissions in 2008–2009 requires a (currently unobserved) recent inflection in the SiO2 trend with time (see Fig. 10 of Harris et al., 2003) toward higher SiO2.

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6

29 Jan 2008 5 4 3 2 1 0

09:30

10:00

10:30

SO2 Emission Rate (kg s-1)

SO2 Emission Rate (kg s-1)

6

02 Feb 2008 5 4 3 2 1 0

11:00

08:00

Local Time

4 3 2 1

07:30

08:00

08:30

09:00

SO2 Emission Rate (kg s-1)

SO2 Emission Rate (kg s-1)

5

09:30

10:00

07 Feb 2009 5 4 3 2 1

0 09:00

09:30

Local Time

10:00

10:30

11:00

Local Time 6

09 Feb 2009 5 4 3 2 1

SO2 Emission Rate (kg s-1)

6

SO2 Emission Rate (kg s-1)

09:00

6

03 Feb 2008

0

08:30

Local Time

6

0

157

12 Feb 2009 5 4 3 2 1 0

09:30

10:00

10:30

Local Time

09:00

09:30

10:00

10:30

Local Time

Fig. 3. Time series of SO2 emission rate. Fully imaged eruption periods are shaded for emphasis. Dashed line on 9th February 2009 shows the pre-ash correction time series for comparison.

Rather, we propose that the observed variations in SO2 emission rate are due to extrusion rate variations. This is supported by comparison of the SO2 emission rate with published estimates of extrusion rate at Santiaguito (Harris et al., 2003; Jeff Johnson, pers. comm. 2010), which suggests that the SO2 emission rate varies with the extrusion rate on a decadal timescale (Fig. 5). Similar behaviour has also been observed during extrusion at Mt. Unzen (Hirabayashi et al., 1995). 7. Mechanism of degassing and eruption at Santiaguito We investigate the trigger mechanism of explosions at Santiaguito by consideration of degassing pathways and processes as outlined in SO2 emission rate data. Johnson et al. (2008) propose a model of Santiaguito explosions in which steady accumulation of gas beneath a thin (20–80 m) viscoelastic lava dome leads to building gas pressure in the subsurface and eventual dome failure through the detachment and uplift of the

viscous lava pad. However several features of the SO2 dataset disagree with aspects of this model. We observe that the viscous “cap-rock” overlying the conduit at Santiaguito is pervasively fractured and is permeable to gas flow throughout the duration of degassing cycles at Santiaguito, as reflected by continuous emissions during repose (Fig. 3). Emissions are observed from a number of different degassing sources on the dome surface over the course of a repose period, with up to three distinct sources observed at one time in our two-dimensional imagery (Fig. 4). Degassing pathways open, degas, and close again over periods of less than 10 min. This pattern may be controlled by the connectivity of bubble chains as degassing pathways through the shallow magma, allowing passive degassing to occur without requiring cracks to seal and unseal repeatedly. This continuous, crater-wide, permeability causes problems for any model based on generating gas pressure beneath this viscous cap. If we consider a model in which explosive degassing overprints continuous passive degassing, we can calculate the level of excess gas

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Table 2 Summary of SO2 measurements. “R” = repose period, “E” = eruption. Emissions which are not part of a fully imaged repose or eruption period at the start and end of the day are labelled “PRE” and “POST” respectively. Date

Period

SO2 (kg)

Time (s)

Average Rate (kg s− 1)

29 Jan 2008

PRE R1 E1 R2 E2 POST PRE E1 R1 E2 R2 E3 POST PRE E1 POST PRE R1 E1 R2 E2 POST PRE E1 R1 E2 PRE E1 R1 POST

157.47 1067.58 813.18 2009.46 823.81 2549.71 1883.45 528.82 1240.58 1269.60 1328.47 368.05 996.07 2127.99 723.96 2389.49 332.10 2569.63 646.85 496.76 558.51 39.57 115.10 965.97 2719.03 802.95 1926.82 1269.21 2668.73 581.86

137 1267 557 1997 652 2786 2300 449 1743 822 2128 325 1703 3234 623 4590 295 5049 562 1430 687 140 302 619 4506 643 2359 714 4219 284

1.15 0.84 1.46 1.01 1.26 0.92 0.82 1.18 0.71 1.54 0.62 1.13 0.58 0.66 1.16 0.52 1.13 0.51 1.15 0.35 0.81 0.28 0.38 1.56 0.60 1.25 0.82 1.78 0.63 2.05

02 Feb 2008

03 Feb 2008

07 Feb 2009

09 Feb 2009

12 Feb 2009

released during an explosion (Fig. 6). An average explosion emits 395 kg of SO2 in excess of that released during passive degassing processes (Table 4). If the mass of gas emitted during an explosion is stored progressively in the subsurface over the course of the preceding repose period (Fig. 6) we can calculate hypothetical “gas storage rates” (Table 4). Gas storage rates are extremely low (0.08– 0.40 kg s− 1) and are as little as 10% of the amount of gas which is free to degas through the dome. This suggests that shallow pressurisation of the dome would be highly inefficient, and argues against gas storage in the shallow subsurface as an explosion mechanism. Furthermore, no relationship between length of repose period and the amount of SO2 emitted during the subsequent explosion is seen in this dataset (Fig. 7). Although based on a limited number of data points, this observation is consistent with thermal, seismic and

Table 3 Daily emission rate at Santiaguito, estimated as the average of emissions across complete eruption–repose cycles. Date

Cycle

Total SO2 (kg)

Time (s)

Average Rate (t d− 1)

29 Jan 2008 29 Jan 2008 29 Jan 2008 02 Feb 2008 02 Feb 2008 02 Feb 2008 2008 AVE. 07 Feb 2009 07 Feb 2009 07 Feb 2009 09 Feb 2009 09 Feb 2009 12 Feb 2009 2009 AVE.

R1 + E1 E1 + R2 R2 + E2 E1 + R1 R1 + E2 E2 + R2

1880.76 2822.63 2833.27 1769.40 2510.18 2598.07

1824 2554 2649 2192 2565 2950

R1 + E1 E1 + R2 R2 + E2 E1 + R1 R1 + E2 E1 + R2

3216.48 1143.61 1055.27 3685.02 3521.98 3937.94

5611 1992 2117 5125 5149 4933

89.09 95.49 92.41 69.74 84.55 76.09 84.56 49.53 49.49 43.07 62.12 59.10 68.97 55.38

Fig. 4. Variability of emission source location during a repose/explosion cycle on 29th January 2008. Two sources of SO2 emission are resolvable in the UV camera image at 09:22:13 (top panel). The source of emissions changes during the cycle, as outlined in the transect SO2 profiles taken at different times throughout the cycle (bottom panel). Shading emphasises inferred emission sources. Transect line is ~ 220 m long.

infrasonic energy proxies (Sahetapy-Engel et al., 2008) and tilt (Sanderson et al., 2010), providing confidence in our conclusions. The lack of correlation between the repose period length and the characteristics of the following explosions suggests that a model in which steady accumulation of gas builds pressure in the subsurface is not viable. Such a model implies that longer time periods spent building pressure should lead to larger explosions (Sahetapy-Engel et al., 2008). We suggest that the observed continuous repose degassing is consistent with degassing of magma in the very shallow conduit beneath a fractured and permeable cap-rock which, in line with observations that the dome surface shows lateral motion toward the active lava flow (Johnson et al., 2008), likely represents a proto-crust of the block lava flow. However, the high permeability of the cap rock is not consistent with eruption model in which gas pressure builds in

A.S.P. Holland et al. / Journal of Volcanology and Geothermal Research 202 (2011) 153–166

1.0 SO2 (this study) SO2 (1976-1991) SO2 (2001-2002) Extrusion Rate

0.8

150 0.6 100 0.4 50

Table 4 Rates of SO2 storage in the subsurface (see also Fig. 6). “Corrected SO2” refers to the total SO2 released in an eruption, minus the background degassing contribution.

Extrusion Rate (m3 s-1)

Average Emission Rate (t d-1)

200

0.2

0 1970

1980

1990

2000

159

0.0 2010

Date

Eruption

Corrected SO2 (kg s−1)

Repose Time (s)

Repose Rate (kg s− 1)

“Storage rate” (kg s−1)

29 29 02 02 02 03 07 07 09 09 12

E1 E2 E1 E2 E3 E1 E1 E2 E1 E2 E1

295.14 217.42 218.13 700.80 143.16 364.25 383.13 236.14 600.11 423.92 770.39

1267 1997 1743 2128 5049 1430 4506 -

0.84 1.01 0.71 0.62 0.51 0.35 0.60 -

0.23 0.11 0.40 0.07 0.08 0.17 0.09 -

Jan 2008 Jan 2008 Feb 2008 Feb 2008 Feb 2008 Feb 2008 Feb 2009 Feb 2009 Feb 2008 Feb 2009 Feb 2008

Date Fig. 5. Variation in SO2 emission rate and magma extrusion rate from 1970 to the present. The SO2 emission rate is broadly correlated with the extrusion rate on a decadal time scale. 1976–1991 SO2 average from Andres et al. (1993). Triangles represent measurement dates, showing that this study did not sample much of the high extrusion rate period. 2001–2002 SO2 average from Rodríguez et al. (2004). 1970–2002 extrusion rates from Harris et al. (2003). 2007–2009 extrusion rates from field observations in January 2007 and 2009 (Jeff Johnson, pers. comm. 2010).

the shallow subsurface beneath this cap rock, and thus we must further consider the source of these explosions. Sanderson et al. (2010) find that the eruption cycle is accompanied by cyclic inflation and deflation of the dome, with a pressurisation source located at depths of 250–275 m beneath the dome surface. They suggest that this reflects the base of a viscous, impermeable plug beneath which gas pressure builds by exsolution into bubbles until internal pressure exceeds magmastatic pressures at which point the plug overcomes friction and slips, and an open fracture system is developed extending to the surface. However once more the observed continuous passive degassing is not consistent with such a model, as the viscous plug (which is viscous as it is itself already degassed) again must be pervasively fractured and permeable. Several other studies (Bluth and Rose, 2004; Sahetapy-Engel et al., 2008; Sahetapy-Engel and Harris, 2009) have proposed plug flow models in which the process controlling the cyclicity is shear fracturing at the conduit margins. In this rheological model, high strain rates at the conduit margins lead to brittle fracture of otherwise ductile magma. A gas and pyroclast mixture is released into these

fractures, which act as permeable degassing pathways for explosive degassing (e.g. Gonnermann and Manga, 2003; Tuffen et al., 2003; Collier and Neuberg, 2006). This model is attractive in the context of gas emission data as presented in this study for several reasons. The shear fracture model implies that gas is stored in deeper fractures and within magma during the course of the repose period, removing the requirement for building gas pressure in the shallow subsurface which, as discussed previously, is problematic in light of the continued permeability of the dome cap-rock. Additionally, during shear fracturing gas acts in a passive manner with respect to explosion triggering allowing the observed explosions to be triggered without the requirement of large gas overpressure to overcome the tensile strength of magma or to lift an overlying cap-rock. The energy for explosion is rather created by the sudden decompression of gas stored deeper in the system. This is compatible with the small amounts of SO2 emitted during single explosions (Table 4) and the small increases in gas emission rate between repose and explosion (Fig. 3). The shear-fracture model has a subtle advantage over that proposed by Sanderson et al. (2010) in that it accounts for the coexistence of pyroclastic explosions and extrusion of block lava flows from the same vent. Shear fracture is driven by the rise of magma and thus the explosions can be viewed as a “by-product” of extrusion, facilitating degassing of a highly viscous magma. Additionally whilst shear-fracturing generates plug flow, it does not necessarily imply that the plug is an anomalously viscous and impermeable body under which gas pressure can form. In fact, in Sections 8 and 9 we explore 2000 2008 2009

Excess SO2 424 kg

-1

SO2 Emission Rate (kg s )

3

Eruption SO2 (kg)

1500 2 4506 s Storage: 0.09 kg s -1

1

0 09:30

Baseline: 0.60 kg s-1

10:00

10:30

Local Time Fig. 6. Annotated time series from 9th February 2009, outlining the calculation of subsurface gas storage rates. SO2 emitted during an eruption is calculated by integrating the time series and subtracting the integral of baseline degassing over the same period (shaded region). Assuming gas is exsolved from magma at a constant rate into a pressurising space during repose, we calculate the rate of gas storage in the subsurface during the repose period.

1000

500

0

0

1000

2000

3000

4000

5000

6000

Repose Time (s) Fig. 7. The total mass of SO2 emitted during single explosions plotted against the length of the preceding repose period. No repose-intensity relationship is observed, consistent with thermal, infrasonic, seismic and tilt observations. Note that there is no notable difference in explosive gas release between the two years. Background-corrected data show similar patterns.

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shear fracture processes and show that a combination of high strain rates and cooling at the conduit margins may lead to the viscous onset of shear-fracturing, whilst the flow interior may be less viscous. This is an important distinction as the magma involved in plug flow eventually feeds block lava flows which have been observed to extend up to 3.75 km from the vent (Harris et al., 2002, 2004). We also note that the annular pattern of explosive emissions is consistent with the shear fracture model (Bluth and Rose, 2004; Sahetapy-Engel and Harris, 2009). Another attractive aspect of shear-fracture at Santiaguito is that conditions for shear fracture can be met continuously in the conduit as long as the conduit geometry and physical characteristics of the magma remain similar, and as long as block lava flow extrusion continues at the surface. As the dacitic lava dome at Santiaguito is near-homogeneous through time, and as extrusion has been continuous through the 30 year observation period over which this type of low intensity explosion has been observed, we suggest that this mechanism provides the stability required for such a continuous sequence of explosions to occur over three decades. Under the shear fracture model, degassing is a two-stage process, reflected in SO2 emissions during explosions and repose periods respectively. Shear fracturing generates degassing pathways at the conduit margins and leads to explosive degassing, and plug flow in the conduit centre. Magma, upon approaching the surface, is able to continuously degas through the pervasively fractured and highly permeable viscous cap rock. This is most likely possible through connected bubble chains. The thin, permeable upper carapace is considered as the beginning of the viscous boundary layer on the lava flow, and is not mechanically important. This model suggests that the observed explosive activity at Santiaguito is a by-product of an extrusive phase of eruption, and is important in the context of understanding processes occurring during extrusion of crystal rich silicic lava during lava dome growth. However, although we reason that shear fragmentation is an attractive driving force of the observed activity, it has been noted that it is not clear that shear fracturing will occur in a system characterised by such low extrusion rates (Massol and Jaupart, 2009). To address this we perform calculations to assess the viability of shear fracture at Santiaguito. 8. Viability of shear fracture at Santiaguito Silicate melts are visco-elastic, meaning that they can behave as viscous melts or brittle glass depending on temperature and deformation timescales. The transition from viscous to brittle behaviour is crossed by lowering temperature or by increasing the applied strain rate (Dingwell and Webb, 1990; Webb and Dingwell, 1990). The failure criterion for melt is given by the relationship (e.g. Goto, 1999; Gonnermann and Manga, 2003; Tuffen et al., 2003): γ˙ η N τs ;

ð1Þ

where γ˙ is the strain rate (s− 1), η the melt viscosity (Pa s) and τs the shear strength of the melt (Pa). We follow Caricchi et al. (2008) by applying Eq. (1) to the failure of magma by substituting the melt viscosity for the apparent viscosity of crystal-bearing magmas. The shear strength of silicic magma is not well constrained but lies in the range 106–108 Pa and may be lowered by the presence of crystals due to local stress amplification (Collier and Neuberg, 2006). The strain rate during Poiseuille flow in a cylindrical conduit is maximum at the conduit walls and is given by: 3

γ˙ max = 4Q =πR ;

ð2Þ

where Q is the volumetric magma flux (m3 s− 1), and R is the conduit radius (m). Thermal images of the Santiaguito summit crater show

three concentric regions: a hot outer annulus through which preferential flow of heat and gas occurs; a cool inner annulus of blocky lava; and a warm central core of hot lava which is connected to the lava flow (Sahetapy-Engel and Harris, 2009). We interpret the warm central core as the surface expression of the conduit, providing a conduit radius of 18 m. The extrusion rate at Santiaguito in January 2009 was measured at 0.25 m3 s− 1 (Johnson pers. comm. 2010). At this extrusion rate the maximum strain rate is 5.5 × 10− 5 s− 1 (Eq. (2)), and brittle fracture of Santiaguito magma will occur for viscosities of 1010.26–1012.26 Pa s and above (Eq. (1)). This fracture criterion can be considered an upper estimate as non-Newtonian behaviour of crystal-bearing magmas may lead to additional strain localisation at the conduit margins. Magma viscosity varies as a function of temperature, melt composition, melt H2O content, and bubble and crystal content. During ascent the viscosity increases in response to degassing and crystallisation. Here we model the magma viscosity over the uppermost 1000 m of the conduit to investigate if the viscous criterion for shear fracture is met and, if so, at what depth. In common with other intermediate magmas (Reubi and Blundy, 2009), Santiaguito dacite is a mixture of rhyolitic melt (72.8 wt.% SiO2, Rose, 1987a) and crystals. Melt viscosity varies strongly as a function of temperature and H2O content (Fig. 8a). Here we make the simplifying assumption of isothermal ascent of magma over the depth range of interest (temperature = 850 °C; Harris et al., 2002). However, temperature is likely to vary radially within the conduit. For long lived magma pathways a thermal boundary layer develops at the margins due to the cooling effect of the surrounding wall rocks. At low extrusion rates (and thus rise rates) conduit marginal magma will spend significant time in this thermal boundary layer, lowering the temperature locally by up to 200 °C (Newman et al., 1988; Tuffen and Dingwell, 2005; Collier and Neuberg, 2006) and increasing melt viscosity by several orders of magnitude (Fig. 8a). As we are interested in the development of shear-driven fracture at the conduit margins we include this effect by modelling the viscosity at both 850 °C and 700 °C. We note that some crystallisation-driven heating may occur in the shallow conduit (Blundy et al., 2006) but that the effect of this heating on viscosity is likely to be minor. Melt is assumed to be H2O saturated across the depth range of interest. Saturation H2O content is modelled using the solubility model of Papale et al. (2006) and decreases from 1.89 wt.% at 1 km depth to 0.02 wt.% at atmospheric pressure. Assuming that the melt composition is otherwise invariant over this depth range, and equal to the matrix glass composition (Rose, 1987a), the melt viscosity is modelled using the viscosity model of Giordano et al. (2008). Magma viscosity further varies due to the presence of crystals and bubbles suspended in the melt. The presence of crystals increases magma viscosity (Caricchi et al., 2007). Total crystallinity at Santiaguito, calculated using K2O mass balance between melt, bulk rock, and plagioclase (the dominant phase at Santiaguito) compositions (Rose, 1987a), is 47%. Unlike previously quoted crystallinities (31%, Rose, 1972) this accounts for both phenocryst and ground mass crystallisation. We assume that phenocrysts grow at chamber depths, whilst the additional 17% crystallisation occurs during rise in the conduit in response to degassing (Blundy and Cashman, 2005; Blundy et al., 2006). Much of this crystallisation occurs outside of the boundaries of our modelling (during ascent from a modelled chamber depth of 5 km; Barmin et al., 2002), so we simplify the model by assuming constant crystallinity. Magma viscosity is calculated using the model of Caricchi et al. (2007), and the observed crystal cargo increases the magma viscosity by ~1.5 orders of magnitude above the equivalent melt viscosity (Fig. 8a). This increase may be considered a minimum estimate as the presence of elongated crystals (such as plagioclase) results in higher magma viscosity than for a magma with the equivalent crystal fraction of spherical crystals (Mueller et al., 2010).

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a

9. Shear fracture processes

20 Wet Melt + Xtls Degassed

Log Viscosity (Pa s)

18

In this section we use our gas emission data and further physical analysis to propose a more detailed model of processes occurring before, during and after shear fracture events, and hence explosions, at Santiaguito.

16

D 14

9.1. Fracture 12

C

10

B

8

A 6 600

650

700

750

800

850

900

Temperature(oC)

b

Log viscosity (Pa s) 6 0

Depth (m)

161

8

10

12

14

16

In Section 6 we present evidence that the extrusion rate of magma scales with the emission rate of SO2. Thus, by combining a known extrusion rate measured in 2009 (0.25 m3 s− 1, Jeff Johnson pers. comm. 2010) with the ratio of observed degassing rates in 2008 (84.56 t d− 1) and 2009 (55.38 t d− 1) we infer an extrusion rate of 0.38 m3 s− 1 in 2008. Corresponding values of γ˙ MAX for each year are 10− 4.08 in 2008 and 10− 4.26 in 2009 (Eq. (2)). Tuffen et al. (2003) consider the development of shear stress as a function of time during visco-elastic deformation, and find that the time taken for shear stress to accumulate to shear fracture is given by:

200

tFRAC = −ðη = G∞ Þlnð1−ðτs = γ˙ MAX ηÞÞ;

400

where G∞ is the shear modulus at infinite frequency (1010 Pa). Modelling the time to fracture as a function of depth (Fig. 9) we see that the fracture time is heavily dependent on the magma strength. Assuming a strength of 107 Pa the time to shear fracture over most of the 265 m depth susceptible to shear fracture is 12 s in 2008 and 19 s in 2009. This calculation leads to two important observations. For the lower extrusion rate in 2009 it takes longer for stress to build to shear fracture. This may partly explain the observation of two distinct styles of degassing during our two field campaigns (Section 5.2). At low extrusion rates, and thus low rise rates in the conduit, fresh magma is supplied to the shallow conduit at lower rates, leading to a reduction in the observed rate of repose degassing. At the same time, the lower rise rate means that stress takes longer to accumulate at the conduit margins, and thus it takes longer for the system to reach failure (i.e. longer repose periods). A second important observation is that the calculated time to fracture is 1–2 orders of magnitude shorter than the repose timescales, meaning that the repose time is not controlled by a single cycle of stress build-up and dissipation through shear fracture.

600 o

800

850 C + Xtls o 700 C + Xtls

1000 Fig. 8. Results of simple viscosity modelling. Fixed parameters throughout are conduit radius (18 m), extrusion rate (0.25 m3 s− 1) and strain rate (10− 4.26). The shaded region represents the range of possible critical viscosities for onset of shear fracturing (for τs = 106 − 108 Pa). a) The effect of different processes is shown in this figure by taking the viscosity of the melt phase at 1000 m (point A, with water content at saturation), adding 47% crystals (point B) and cooling the marginal region by 150 °C (point C). Shear fracture does not occur at this depth for any reasonable value of τs. The path from C to D represents decompression driven degassing as magma ascends in the conduit. This path crosses the critical viscosity for shear fracturing, indiciating that at some point in the conduit shear fracturing will occur. b) Modelling the variation of viscosity with depth to determine the depth of onset of shear fracturing. Modelling results show that at 850 °C shear fracturing is unlikely to occur, although at high crystallinities it is possible in the shallow subsurface if the magma strength is low. At 700 °C, however, the onset of shear fracturing occurs at between 140 and 600 m depth.

ð3Þ

Time to fracture (s) 100 0

102

103

100

τs = 108

200

Depth (m)

We follow Caricchi et al. (2008) and consider the effect of bubbles on magma viscosity secondary, as the presence of crystals results in larger viscosity increases and a more significant decrease in viscosity with increasing strain rate. We thus simplify the problem by omitting the effect of bubbles. Although we have employed a number of assumptions and simplifications during modelling, the predicted magma viscosity at the surface (at 850 °C, Fig. 8b) is in excellent agreement with calculated lava flow viscosities (109.6–1010.8 Pa s, Harris et al., 2004). This provides confidence that our model includes the most important controls on the magma viscosity. The results of our modelling suggest that although Santiaguito magma ascends slowly, the high magma viscosity means that sheardriven fracturing will occur at the conduit margins. For our best estimate of conduit geometry (r = 18 m) the onset of brittle fracture in the magma occurs between 140 and 600 m depth (Fig. 8b), consistent with independent geophysical estimates of the depth of explosion initiation (Sahetapy-Engel et al., 2008; Sanderson et al., 2010).

101

300

τs = 107

400 500 600

τs = 106

2008 2008 2009 2009

700 Fig. 9. Results of modelling the time to shear fracture, taking into account the interplay between shearing at the conduit margins and relaxation. Fracture times are modelled for magma strengths of 106, 107 and 108 Pa, although any intermediate magma strength is possible. Time to fracture is differs from 2008 to 2009 due to different extrusion, and hence shear strain, rates.

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100

9.2. Fracture connectivity as a trigger

Critical Volume Fraction

Oblate We envisage that the critical stress for fracture is crossed repeatedly during repose, each time opening small fractures and dissipating stress which immediately begins to build once more. The critical timescale for an explosion (i.e. the repose time) is controlled by the generation of a fully connected fracture network extending to the surface. For an average background-corrected explosion, 395 kg of SO2 is emitted (Table 4). Assuming an SO2 mass fraction of 0.0135–0.0297 (Gerlach, 2004; Ohba et al., 2008) the total gas mass emitted during an explosion is 13 000–29 000 kg. The volume of this mass of ideal gas composed of pure H2O (by far the dominant species) is 1.1– 2.3 × 105 m3 at atmospheric pressure. Assuming that Santiaguito explosions can be approximated as instantaneous finite-volume releases (Yamamoto et al., 2008), this gas volume is equivalent to an initial explosion cloud of radius 33–43 m (assuming an oblate spheroid shape). This fits well with initial plume widths measured in UV images. For a modelled fracture initiation depth of 265 m (τs = 107 Pa) gas is stored in the subsurface at an average pressure of 3.1 MPa (assuming even distribution with depth and magmastatic conditions) and has a stored volume of 3480–7656 m3. Prior to an explosion this volume of gas is stored in fractures throughout the conduit marginal region, which are prevented from closing on relaxation timescales due to the presence of gas in the fractures (slow H2O diffusion in rhyolitic melt is the rate limiting process; Zhang and Behrens, 2000). The marginal region which is susceptible to shear fracture has a theoretical width of ~2.5 m (Gonnermann and Manga, 2003), in good agreement with field observations at dacitic–rhyolitic centres (Stasiuk et al., 1996; Cashman et al., 2008). This marginal region extends from 265 m to the surface (τs = 107) and thus, for a conduit of 18 m radius, the volume of the region susceptible to shear fracture at any given time is ~67 000 m3. Gas is stored throughout this region and makes up a volume fraction of 0.05–0.11 at the point of explosion. Continuum percolation simulations (Garboczi et al., 1995; Saar and Manga, 2002) suggest that a volume fraction of 0.05–0.11 of softcore (i.e. intersecting), randomly oriented ellipsoids will lead to the development of continuous networks for oblate spheroids with aspect ratio of 7.32–15.46 and greater (Fig. 10). However, as the gas erupted at the surface represents only that stored in the connected “backbone” of such a simulation (e.g. Walsh and Saar, 2008) the overall amount of gas stored in the marginal region will likely be higher, making this critical aspect ratio a maximum estimate. Tuffen and Dingwell (2005) observe no preferred fracture orientation of stage 1 shear fractures, justifying the use of an isotropic simulation here. Shear fractures at exposed rhyolitic conduits (Stasiuk et al., 1996; Tuffen and Dingwell, 2005) typically have apertures of 1 cm and lengths of tens of centimetres, suggesting that natural fractures (modelled as oblate spheroids) exceed the critical aspect ratio for connectivity. We thus suggest that the development of a fully connected network of shear fractures to the surface is a likely explosion trigger at Santiaguito. This model is consistent with field and experimental observations. Benson et al. (2008) showed that during shear deformation samples deform by micro-cracking, generating low-level, high-frequency, seismicity. However, under high pore fluid pressure (expected at Santiaguito due to water saturation) the initial high frequency seismicity is suppressed, and seismicity is confined to a supra-exponential increase in seismicity leading to failure (Benson et al., 2010). At Santiaguito, repose periods are seismically quiet until seismicity begins 2–4 min prior to an explosion, possibly reflecting this late increase in seismicity (Johnson et al., 2009). In the laboratory, the passage of a gas and pyroclast mixture through the tortuous fracture network is recorded as a long period event, equivalent to the explosion long-period event seen at Santiaguito. Whilst this laboratory

Prolate

10-1

10-2

10-3

10-4 10-4

10-3

10-2

10-1

100

101

102

103

104

Aspect Ratio Fig. 10. Modelling the critical volume fraction of soft-core ellipsoids required to develop full connectivity, as a function of ellipsoid aspect ratio. Calculated volume fractions of gas within the shear fracture susceptible region is marked along with the equivalent aspect ratios for oblate and prolate fractures. We consider fractures to be oblate ellipsoids, and find that connectivity is possible for fractures of aspect ratio of 7.3–15.5 and above. Field observations (Stasiuk et al., 1996; Tuffen and Dingwell, 2005) suggest that natural fractures exceed this aspect ratio. Thus modelling suggests connectivity as a possible eruption trigger. Figure modified after Garboczi et al. (1995) and Saar and Manga (2002).

work was carried out at room temperature, Tuffen et al. (2008) have shown that brittle failure in high temperature magma is seismogenic. In other intermediate systems, the short fracture timescales calculated in Section 9.1 are in line with the spacing between events in seismic swarms at Montserrat (as short at 8 s, Green and Neuberg, 2006), which are considered to be an expression of brittle failure of melt at the conduit margins (Neuberg et al., 2006). They are also in line with the periodicity of “drum-beat” earthquakes observed during the 2004–2008 period of extrusion at Mt. St. Helens (~60 s Iverson et al., 2006; Moran et al., 2008). Both types of seismicity are overlain by longer-period tilt cycles (Green and Neuberg, 2006; Anderson et al., 2010) which we speculate may be controlled by connectivity of successive shear fractures. Opening gas-filled shear fractures leads to an increase in volume and hence inflation, whilst the development of connectivity leads to expulsion of gas, closure of fractures and deflation. 9.3. Gas source Gas is likely to be sourced from the magma immediately surrounding shear fractures. Assuming a melt sulphur content of 200 ppm (Andres et al., 1993) and a bulk dacite melt fraction of 0.53 we calculate that 1 m3 of dacite (density 2500 kg m− 3) contains 26.5 kg of S. For an H2S/SO2 ratio of 1 (see data collection in Edmonds et al., 2010), 395 kg of sulphur is released during a single explosion, equivalent to complete degassing of 14.9 m3 of dacite per explosion. The total gas volume released implies the equivalent of 5.9 × 106 ellipsoidal cracks of dimension 1 × 15 × 15 cm. The observed sulphur degassing is equivalent to a uniform melt layer of thickness 1.74 × 10− 5 m (17 μm) around each fracture. 9.4. Gas decompression — the driving force When a fracture network reaches the surface the pressure inside the entire network drops to atmospheric pressure, facilitating the explosive expansion of gas trapped within the network. For a network extending from the surface to 250 m depth under magmastatic pressure, atmospheric connectivity leads to a pressure drop inside the fracture averaging 3.1 MPa. This leads to a gas expansion of 30 times by volume (ideal gas, 100% H2O), providing the driving force for the observed gas venting at the surface. This calculated pressure is

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consistent with experimental results of Chojnicki et al. (2006), who suggest that Santiaguito explosions are driven by 5–100 times gas volumetric expansion. As gas is passive with respect to fracture processes, this mechanism does not require large gas overpressure, in contrast to models of Vulcanian eruptions where gas pressure must overcome the tensile strength of the overlying magma. This is consistent with the relatively small volume of gas released during a single explosion at Santiaguito (Table 4). 9.5. Fracture healing — driving cyclicity After expulsion of the gas phase, the pressure differential between the fracture walls and interior closes the fracture network. Annealing of shear fractures occurs by diffusion-related processes which establish cohesion between adjacent surfaces and by viscous deformation which acts to remove pore space between imperfectly packed particles (Tuffen et al., 2003). Assuming that the fracture is closed to the atmosphere when cohesion is established, a shear fracture will heal over a duration equal to the relaxation time of the melt trelax = η = G∞

ð4Þ

The presence of crystals will act to disturb melt relaxation and we assume that we can substitute the magma viscosity for the melt viscosity to take this effect into consideration. Fig. 11 shows the evolution of the relaxation timescale with depth. At 250 m depth, relaxation of Santiaguito magma occurs over timescales of ~30 s, and this increases toward the surface. Thus the deepest material will heal most robustly, and material will tend toward an unconsolidated gouge as the surface is approached (Section 9.6). Tuffen et al. (2003) show that tuffisite veins form through brittle fracture, heal, are subject to ductile flow, and then fail once again through brittle fracture. Similarly Santiaguito explosions, triggered at the moment of connectivity of a system of shear fractures, are ended when gas is expelled from the system of fractures which subsequently relaxes and heals. This healing allows stress to begin to build and a new fracture network to develop during the next repose period, driving the observed periodicity at Santiaguito.

Relaxation Time (s) 10-1 0

100

101

102

103

104

105

Depth (m)

50

163

9.6. Viscous cessation of healing We assume once again that the annealing timescale can be approximated by the relaxation timescale, as this is the timescale over which cohesion is established. By modelling the relaxation, and hence annealing, timescale as a function of depth it is clear that in the shallowest regions of the conduit the relaxation time exceeds the repose interval between explosions (Fig. 11). If an explosion is accompanied by some amount of plug-flow magma ascent, then at these shallow levels the conduit marginal material will be repeatedly disrupted without ever managing to heal. We propose that this results in the formation of an unconsolidated gouge layer. Modelling suggests that this layer at Santiaguito will have a depth of 30–50 m. The presence of an unconsolidated fault gouge zone has been reported at Mount St. Helens (Cashman et al., 2008). This gouge forms a voluminous source of ash which is emitted during gas venting episodes (Rowe et al., 2008). Evidence for the fluxing of a gas and ash mixture through a pre-existing, unconsolidated gouge zone can also be found at Mule Creek Vent, New Mexico, where tuffisite veins are observed winding between autobrecciated clasts, implying that autobrecciation occurred prior to tuffisite veining (Stasiuk et al., 1996; Rust et al., 2004). At Santiaguito, evidence for such a gouge layer comes from field observations. The hot outer annulus in the Santiaguito crater, from which most explosive degassing occurs, is made up of fine material which is pitted following venting of gas during an explosion (Bluth and Rose, 2004; Sahetapy-Engel and Harris, 2009). Furthermore, thermal imaging of explosion plumes (Sahetapy-Engel and Harris, 2009) suggests that much of the ash erupted is cool and ‘accidental,’ consistent with a gouge body as a source of ash. 9.7. Model summary Stress builds up in the conduit marginal region during flow of viscous, crystal-rich magma leading to shear fracturing from depths of approximately 250 m to the surface (Fig. 12). Stress accumulates and causes brittle failure on timescales 1–2 orders of magnitude less than the repose timescale, and individual shear fracture events generate large numbers of small cracks. Gas and ash are released into fracture interiors, keeping fractures open beyond the relaxation timescale. This process repeats until fractures obtain connectivity to the surface at which point the fracture network pressure drops and gas explosively expands, causing the observed explosions. Following the flushing of the gas–pyroclast mixture from the crack network, fractures heal over relaxation timescales. New melt rises to meet the fracture criteria and the process repeats. At the shallowest sections of the conduit marginal region an unconsolidated gouge zone may form, providing a source of “accidental” ash during explosions. 10. Wider implications

100

150

200

o 850 850 oC C 700 700 C C

250 Fig. 11. Relaxation timescale calculated as a function of depth. Fracture healing occurs by development of cohesion between fracture walls and occurs over the relaxation timescale, which varies as a function of magma viscosity. We consider that an unconsolidated gouge may develop once the relaxation time exceeds the repose time between events, as healing does not occur between successive eruptive plug flow motions. The shaded critical region represents the average 2008 and 2009 repose period durations. Both model runs are for melt with 47% crystals. This calculation suggests that an unconsolidated gouge of between 33 and 48 m depth will develop.

Omitting the effects of bubbles on magma rheology (Llewellin and Manga, 2005) is a key approximation adopted during modelling in this study. However, it has previously been demonstrated that shear fracturing is possible over a range of conditions for bubbly flows, where crystallisation is omitted (Gonnermann and Manga, 2003). This contribution, displaying the importance of crystallisation in the development of high viscosity and shear-fracture, thus represents an alternate end-member study. An integrated approach is not possible at present as current understanding of three phase rheology, and the interactions between crystals and bubbles in a three phase system, is currently poor (e.g. Walsh and Saar, 2008). In this study we find that shear fracturing can occur, and promote open system degassing, during very slow extrusion of intermediate magma. This is in line with previous studies suggesting that low ascent rates of magma are associated with open system degassing

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Explosive Degassing

Repose Degassing

Magma

11. Conclusions

Carapace

High temporal resolution SO2 emission rate data collected with a UV camera shows that explosive activity at Santiaguito is inconsistent with gas accumulation in the upper conduit beneath a viscous cap rock. The cap rock is shown to be pervasively fractured and is permeable to gas flow at all times. Rather, gas emission data favour shear fracturing at the conduit margins as an explosion mechanism. Passive degassing occurs continuously and may be facilitated by the connectivity of bubble chains in the shallowest subsurface. By combining field observations with viscosity modelling, we show that shear fracturing is likely to occur at Santiaguito, even through the rise rate of magma in the conduit is low. Furthermore, we are able to recreate the depth of explosion onset determined by independent methods (150–600 m) using realistic parameters and simplifying assumptions. Results of viscosity modelling suggest that shear fracture-driven degassing occurs by the generation of large interconnected networks of smaller shear fractures, and that healing of these fracture networks following degassing drives the cyclic nature of activity. It is likely that healing of these fractures will cease at the shallowest depths, leading to the development of an unconsolidated fault gouge. Our results suggest that the explosive activity at Santiaguito is the by-product of an extrusive phase of eruption, and that shear fracturing is a viable mechanism of promoting open-system degassing during ascent of crystal-rich intermediate magmas, as the high viscosity of such magma overcomes the low strain environment provided by the typically slow extrusion rates of these magmas.

Fractures

Onset of Brittle Failure

Plug Flow Strain Localisation Poiseuille Flow

Fig. 12. Schematic diagram outlining degassing processes and pathways at Santiaguito. During magma ascent degassing, crystallisation and conduit-marginal cooling lead to increases in viscosity until the shear fracture threshold is crossed and brittle failure occurs (dashed line). Magma ascent is modelled as Poiseuille flow (bottom velocity profile) but strain likely localises toward the conduit margins (middle velocity profile) leading to shear fracture at slight deeper levels than ismodelled in this paper. The brittle shear fracture process opens fractures which develop into networks extending to the surface, allowing gas trapped within the network to explosively decompress (large wavy arrow). This results in plug flow in the upper conduit (upper velocity profile), and the upwelling magma plug flows into the block lava flow (thick arrow). The viscous carapace is pervasively fractured and allows continual release of gas derived from shallow degassing of magma in the shallow conduit (small wavy arrows).

(Eichelberger et al., 1986; Jaupart and Allègre, 1991; Eichelberger, 1995; Jaupart, 1998; Gonnermann and Manga, 2007). We speculate here that it is possible that low extrusion rates may be necessary for the development of shear fracture networks. Whilst high rise rates of magma ascent lead to higher strain rates, such rapid decompressions also lead to lower levels of crystallisation during ascent and thus lower magma viscosity (e.g. Castro and Gardner, 2008). As such, shear fracturing conditions may not be met for rapidly ascending magma, as calculated for the 2008 eruption of Chaitén where rhyolitic magma ascended at a rate of 1 m s− 1 without reaching the shear fracture criterion (Castro and Dingwell, 2009). A fuller understanding of the role of shear fracturing in aiding the transition from explosive to effusive eruption styles, and the role of ascent rate in determining the eruptive style, requires numerical modelling of magma flow which includes crystallisation kinetics. We suggest that the current low-level explosive activity at Santiaguito is a by-product of a dominantly extrusive eruption. Our study supports previous field and geochemical evidence (Gonnermann and Manga, 2003; Tuffen et al., 2003; Tuffen and Dingwell, 2005; Collier and Neuberg, 2006) which suggest that shear fracturing is an important process for driving open system degassing at shallow depths in the conduit during extrusion. However, whilst shear fracturing may act to promote extrusive activity, we are unable to show uniquely that without this degassing mechanism activity at Santiaguito would be more explosive and hazardous. Further understanding of this problem requires stronger constraints on the pre-eruptive volatile contents of Santiaguito magma.

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