Structure of the Grímsvötn central volcano under the Vatnajökull icecap, Iceland

Transcription

1 Geophys. J. Int. (2007) 168, doi: /j X x Structure of the Grímsvötn central volcano under the Vatnajökull icecap, Iceland Raimon Alfaro, 1 Bryndís Brandsdóttir, 2 Daniel P. Rowlands, 1 Robert S. White 1 and Magnús T. Gudmundsson 2 1 Bullard Laboratories, University of Cambridge, Cambridge, UK 2 Institute of Earth Sciences, University of Iceland, Reykjavík, Iceland. bryndis@raunvis.hi.is Accepted 2006 September 25. Received 2006 September 25; in original form 2005 August 2 1 INTRODUCTION Volcanism along the divergent plate boundary across Iceland is confined to close to 30 discrete volcanic systems arranged en-echelon within four major zones, the Reykjanes Peninsula (RP) and the Western (WVZ), Eastern (EVZ) and Northern Volcanic Zones (NVZ), Fig. 1 (Saemundsson 1978). Each volcanic system consists of a central volcano transected by a rift zone which can extend 20 km in width and over 100 km in length (Saemundsson 1979). Being rooted in crust two to six times thicker than normal oceanic crust (e.g. Menke et al. 1996; Darbyshire et al. 2000; Weir et al. 2001; Allen et al. 2002) these plume-enhanced volcanic systems are structurally and geochemically more complicated than their oceanic counterparts. Rhyolitic rocks, formed by remelting of hydrated crust in the vicinity of shallow crustal magma chambers (e.g. Sigmarsson et al. 1991), are common within the more evolved systems, some of which have developed calderas. Gravity data (White et al. 1995; Darbyshire et al. 2000; Gudmundsson 2003) and seismic measurements of crustal thickness within Iceland (e.g. Darbyshire et al. 1998; Weir et al. SUMMARY The subglacial Grímsvötn central volcano, lying within a volcanic zone directly above the core of the Iceland mantle plume, is one of the most active in Iceland. Local, regional and teleseismic earthquake data recorded on a temporary seismometer array across western Vatnajökull icecap during the summer of 1998 have provided a three-dimensional image of the shallow crustal structure of the volcano. Microearthquake activity at depths of 1 4 km along the Grímsvötn caldera rim coincided with inflation of a shallow magma chamber beneath the caldera, which culminated in a 0.1 km 3 eruption in December Tomographic inversion of these earthquakes define the extent of a low-velocity body beneath Grímsvötn with a volume of 20 km 3 extending to 3kmbelow the surface. This low-velocity body is flanked by high velocities under the caldera rim. Delays in the P-wave arrival times through the Grímsvötn caldera from regional and teleseismic earthquakes and from two detonations 150 km east of Grímsvötn are s greater than the delays through the uppermost 3 4 km of crust shown by local earthquake arrivals. This suggests the presence of a further low-velocity body at depths greater than 3 4 km beneath Grímsvötn, presumed to be due to the presence of melt. Using the distribution of local seismicity and shear wave attenuation we estimate the maximum lateral extent of the region containing partial melt to be 7 8 km E W and 4 5 km N S. P-wave delays require a thickness of less than 1 km of pure/high percentage partial melt, assuming a sill-like magma chamber. Key words: earthquake location, Iceland, tomography, volcanic systems. 2001; Allen et al. 2002) indicate increasing melt production towards the centre of the Iceland plume. A cluster of subglacial volcanoes marks the presumed centre of the Iceland hotspot, at the junction of the EVZ and NVZ (Björnsson & Einarsson 1990). This region has been seismically active since monitoring began (Tryggvason 1973; Brandsdóttir 1984). Local seismic networks, in operation during the last three decades, show the seismicity to be clustered within the Bár arbunga and Grímsvötn central volcanoes and beneath two subglacial geothermal areas at the Loki Ridge, northwest of Grímsvötn (Einarsson 1991; Jakobsdóttir et al. 2002). The Grímsvötn and Bár arbunga volcanic systems have been among the most active in Iceland during historical time (Larsen et al. 1998, and references therein). Their central volcanoes have produced some of the largest volcanic episodes documented along the plate boundary, such as the Vei ivötn eruption (from Bár arbunga) in ca (Larsen 1984) and the km 3 Laki eruption in (from Grímsvötn) (Sigurdsson & Sparks 1978). Grímsvötn has the highest eruption frequency of any volcanic system in Iceland GJI Volcanology, geothermics, fluids and rocks C 2006 The Authors 863

2 864 R. Alfaro et al N 64.6 N V V10 B6 W V9 V12 V Lok i Ridge V eruption site V13 E V6 High geothermal activity V16 V29 B9 Station used in LET of Alfaro (2001) Station on refraction Line 1 Other station active for shots V15 B5 B8 B4 B3 V5B2 V4 B1 V3 B0 B14 V2 V14 Gjálp N V24 V17 M B47 V1 V21 V23 V18 V19 V20 V22 E B46 V N 64 N Grímsvötn caldera V41 Line 1 RR RP 1400 WVZ 1500 He 0 10 KR EVZ Ka 1600 TFZ NVZ As SH1 SH2 25 W 20 W 15 W V42 B43 V48 V49 20 km 17.5 W 17.0 W Figure 1. Network layout of the passive and active survey used in the V98 experiment. Travel times along Line 1 were modelled using the Xrayinvr ray tracing software (Zelt & Smith 1992). Two major geothermal areas, the eastern (E) and western (W) Skaftár cauldrons, exist on Lokahryggur (Loki Ridge). The inferred margin of Grímsvötn caldera is denoted by black inward-ticked lines, (based on work by Gudmundsson & Milsom 1997) M: main caldera, N: north caldera, E: east caldera. The 1996 Gjálp fissure eruption and the 1998 eruption on the southern rim of the Grímsvötn caldera are shown by thick black lines. Inset shows tectonic map of Iceland, based on a map from Einarsson & Saemundsson (1987). Icecaps are represented in white; central volcanoes and caldera are outlined in black; rift zones are shown in light grey. The plate boundary is divided into the Reykjanes Peninsula (RP), the Western Volcanic Zone (WVZ), the Eastern Volcanic Zone (EVZ), Northern Volcanic Zone (NVZ), Reykjanes Ridge (RR), Kolbeinsey Ridge (KR), and the Tjörnes Fracture Zone (TFZ). Also shown are the volcanic systems Hengill (He), Katla (Ka) and Askja (As). SH1 and SH2 denote the location of two offshore detonations. since A.D. 1200, including five eruptions in (Gudmundsson & Björnsson 1991; Larsen et al. 1998; Vogfjörd et al. 2005). In October 1996 an eruption occurred at Gjálp, between Grímsvötn & Bár arbunga (Einarsson et al. 1997; Gudmundsson et al. 1997) and a smaller eruption took place within the Grímsvötn caldera in December 1998, four months after our seismic survey (Fig. 1) and again in November 2004 (Vogfjörd et al. 2005). Grímsvötn is a km wide and m deep depression within the central-western Vatnajökull ice cap. The bedrock topography of the Grímsvötn central volcano has been mapped by radioecho sounding (Björnsson & Einarsson 1990) revealing a composite caldera which Gudmundsson & Milsom (1997) divided into three regions, the east, north and main (or south) calderas (Fig. 1). A combined magnetic and seismic reflection survey of the main caldera indicated that the caldera floor is made of volcaniclastic sediments, lava flows and sills with lavas and sills being more prominent in its southern and southwestern parts (Gudmundsson 1992). The Grímsvötn caldera lake is largely confined to the main caldera and is overlain by a m thick floating ice shelf. Prior to the 1996 Gjálp eruption the ice shelf rose m per year as meltwater accumulated from numerous subglacial geothermal outlets within the caldera. The lake was partially drained in jökulhlaups (bursts of water) when its water level had risen to a critical height of m a.s.l. However, since the 3.6 km 3 jökulhlaup following the Gjálp eruption in 1996, during which the ice dam was damaged, the pattern of ice shelf rise and subsidence became more irregular and more frequent (Björnsson et al. 2001). The volume of the caldera lake was at a minimum during the summer of Evidence for shallow magma chambers in Iceland has been provided by shear-wave attenuation studies (Einarsson 1978), detailed seismic refraction studies (e.g. Gudmundsson et al. 1994; Brandsdóttir et al. 1997) and geodetic measurements during inflation and deflation episodes at central volcanoes (e.g. Tryggvason 1986, 1989; Sturkell et al. 2003; 2006). However, no undershooting refraction surveys had been conducted across any of the subglacial volcanic systems above the centre of the Iceland hotspot prior to this study. The 1995 ICEMELT profile (Darbyshire et al. 1998) lay

3 Structure of the Grímsvötn central volcano km to the northeast of Grímsvötn, straddling the eastern flank of Bár arbunga and did not cross the caldera. Lacking knowledge of the crustal structure of these hotspot volcanoes we conducted a seismic survey across W-Vatnajökull during the summer of Although focused on the Grímsvötn central volcano our array encompassed the Loki Ridge and extended north to Bár arbunga. Our main goal was to use wide-angle refraction data as well as local, regional and teleseismic events to map the outlines of a shallow magma chamber suspected to lie beneath the Grímsvötn caldera and to have fed a relatively small eruption along the southwestern caldera rim in 1934 and 1983 (Einarsson & Brandsdóttir 1984). The experiment was scheduled during a period of increased seismicity which turned out to stem from inflation of a shallow magma chamber beneath the Grímsvötn caldera (Sturkell et al. 2006), culminating with a 10 days eruption at the southwestern caldera rim in 1998 December. 2 THE V98 EXPERIMENT The V98 network consisted of an array of 40 three-component 1 and 2 Hz seismometers with an average spacing of 3 km deployed during 1998 May August within three connected networks (Fig. 1). Additional 16 stations were deployed during a brief refraction survey at the end of the field season. The station network was designed to cover regions which have experienced high seismicity during the last three decades: Grímsvötn, Lokahryggur (Loki Ridge) and Bár arbunga. Station failures within the Bár arbunga array resulted in limited data from this region, hence this study concentrates on the region around Grímsvötn. The ice thickness in the study region, derived from radio echo-sounding measurements, varies from afew metres near the southern caldera rim to m within the caldera, with an average thickness of 400 m outside the Grímsvötn region (Björnsson & Einarsson 1990). Initally, care was taken to deploy the stations outside the Grímsvötn caldera lake. A rise in the lake level during the experiment may have reached station V23 in August. Seismic arrivals were recorded from local, regional and teleseismic earthquakes and from two 200 kg explosions in the sea km to the east. Timing was from GPS clocks, and station locations from averages of 3-D GPS measurements. A total of 398 local events was picked, producing 5201 P- observations and 488 S-observations. The mean P- and S-wave picking uncertainty is 0.03 and 0.3 s, respectively. 3 3-D VELOCITY IMAGES OF GRÍMSVÖTN All seismic events were located initially using the computer program Hypoinverse (Klein 1978). For each of the locally recorded events a minimum of six P-wave arrivals was used, with the average being ten. S-wave arrivals were added where possible to further constrain focal depth. However, shear wave attenuation within the Grímsvötn caldera severly limited the number of well-constrained shear arrivals. In the following sections we discuss how the tomographic velocity structure was derived, starting from the Hypoinverse locations calculated with a best guess 1-D velocity model, through a minimum 1-D V p velocity model calculated by iterating both the hypocentral locations and velocities, to a final fully developed 3-D V p velocity model. 3.1 Minimum 1-D V p model The minimum 1-D V p velocity model was computed using the software Velest (Kissling et al. 1994), which inverts simultaneously for hypocentral locations, 1-D velocity model and station corrections. To ensure stability in the inversion we used only well-located events (Thurber 1993), with an azimuthal gap of 180 and at least 7 P- and S-wave observations for each event. This reduced the data set to 325 events with 4272 P and 421 S-wave traveltimes. Data from the nearby ICEMELT refraction profile (Darbyshire et al. 1998) which passes to the north of Grímsvötn was used as a starting model for the velocity structure. To add additional constraint to the shallow velocities beneath the caldera, data from a 1987 shot and a 1988 refraction profile (Gudmundsson 1992) were incorporated in the tomographic inversion (Fig. 2). The Velest ray tracer requires stations to be located within the top layer and only one layer is allowed above sea level. Including V98 station elevations, ranging from 1380 to 1950 m, an initial model with a 2 km thick top layer of 3.6 km s 1 (e.g. Gudmundsson et al. 1994) and increasing velocity with depth introduced instabilities in the inversion and poor convergence even at depths where good resolution is expected. A more stable inversion was obtained by omitting station elevations and incorporating systematic errors in calculated traveltimes into individual station corrections (Kissling et al. 1994). Station corrections, including both variations in station elevation and near-surface structure, relative to a reference station (V15, Fig. 1) with good ray distribution (Haslinger et al. 1999) are fairly consistent across the V98 network with slightly earlier P-arrivals at stations around the caldera rim and delayed P-arrivals at stations west and north of the caldera. To test the performance of the final minimum 1-D V p model, hypocentres were systematically shifted to greater depth by 3 km and the inversion re-run using both damped and fixed velocity models. In both cases the events relocated back to their original locations, with only a slightly (<±0.2 km) larger spread in the damped velocity solution, which allowed trade-off with minor velocity changes. Within the array region, there is no significant trade-off between longitude or latitude and depth. However, there is an obvious tradeoff between velocities and focal depth. Available refraction data was therefore used to further constrain shallow structure within the Grímsvötn caldera (Fig. 2). The final minimum 1-D V p model produced an rms traveltime residual of 0.07 s for 325 events, a reduction of 0.03 s from the initial model. Events originally located at the surface were relocated to greater depths, producing an improved Gaussian depth distribution curve with a clear depth-frequency maximum between 2.3 and 2.5 km depth D upper crustal structure We performed an iterative, 3-D simultaneous damped least squares inversion of hypocentral parameters and velocity using the Simul2000 algorithm developed by Thurber (1993), Eberhart- Phillips (1993) and Thurber & Eberhart-Phillips (1999). A suite of ray paths connecting each source and receiver are calculated using an approximate 3-D ray tracing (ART) algorithm. The velocity model in Simul2000 is parametrized by a 3-D grid of nodes defined by the intersection points of three sets of orthogonal planes. Trial runs were carried out in order to determine optimum model parametrization. As discussed in detail by Kissling et al. (2001), model parametrization is a trade-off between parameter resolution and image fidelity. The node spacing for the V98 study region was optimized by taking

4 866 R. Alfaro et al. b) a) data variance (s ) damping = c) latitude solution variance (km/s2) d) damping = damping = longitude % Vp change, rel. to 1D initial mode l Figure 3. Trade-off curve to determine the damping value for the 3-D inversion. The selected damping value of five is denoted by a green cross. (b), (c) and (d) are plane views (at sea level, z = 0 km) of velocity perturbations (after one iteration) relative to those of the minimum 1-D model with varying damping value. Horizontal nodes are marked by crosses. Figure 2. (a) Parametrization of the model space excluding peripheral nodes and nodes at the edge of the model in the E W and N S direction. Note the larger grid node spacing (crosses) north of N. Stations used in the passive survey are marked by black triangles, epicentral locations by white circles and ray paths are grey. Black rays represent ray paths from the 1987 and 1988 refraction profiles. The Grı msvo tn caldera is denoted by black inward-ticked lines. (b) Projection of ray paths and seismic events across the caldera. into account station distribution and ray coverage. A horizontal node spacing of 2 km was chosen for the Grı msvo tn region (Fig. 2a). As ray paths are most sensitive to velocity variations near their turning points the vertical node spacing was smaller, 1 km in the uppermost 4 km and 2 km at greater depths (Fig. 2b). The initial velocity model for the 3-D inversion was derived from the minimum 1-D V p model by linear interpolation. Suitable velocity damping parameters for the inversion were found using a trade-off curve between data variance and solution variance (model roughness), (Eberhart-Phillips 1986). The optimal damping parameter of five was selected by running a series of one-iteration inversions with a large spectrum of damping values and selecting the value that resulted in a significant reduction in data variance without causing a marked increase in solution variance (Fig. 3). Model perturbations for underdetermined parameters are strongly dependent on the chosen damping value which explains why large velocity perturbations are observed with small damping values. The most suitable damping value after four iterations based on our grid spacing and ray coverage was also found to be five. The ICEMELT profile has relatively high bedrock velocities (4.06 km s 1 ) in comparison with other velocity models within the Neovolcanic Zones of Iceland (2 3 km s 1, e.g. Flo venz & Gunnarsson 1991; Brandsdo ttir et al. 1997). As ice has higher velocities (3.6 km s 1 ) than the underlying bedrock according to a shallow refraction profile across the Grı msvo tn caldera (Fig. 2, Gudmundsson 1992) which has near-surface velocities increasing from 2.6 km s 1 at 0.75 km elevation to 3.3 km s 1 at sea level, a thin low velocity layer is present beneath the glacier. In order to test the influence of the inversion on the uniqueness of the final model the 3-D inversion was also derived using the ICEMELT 1-D model and a constant velocity gradient model. All three models produced similar velocity models and rms residual values of s, although vertical smearing is minimized by the minimum 1-D model within the deeper, less well-resolved layers. C 2006 The Authors, GJI, 168, C 2006 RAS Journal compilation

5 Structure of the Grímsvötn central volcano 867 Four iterations were required to reach the final preferred model. The final variance ratio was 1.08, after which model adjustments were considered to be insignificant. Data variance and rms residuals were reduced by 85 and 58 per cent, respectively, relative to the initial model. The resulting velocity anomalies relative to the minimum 1-D model are discussed below. The ice-bedrock boundary low-velocity layer was not inverted for and was approximated by a series of horizontal planes separated in elevation by 100 m. A velocity of 3.6 km s 1 was used for the top layer and held fixed during the inversion. Ray paths were calculated using the approximate 3-D ray tracing (ART) algorithm. Pseudobending (Um & Thurber 1987) has not been used in this analysis as it introduced large instabilities in the inversion. However, Eberhart- Phillips (1990) has shown that the introduction of pseudo-bending to the 3-D ray tracing has only a small effect on velocity variations, an observation also supported by this study. Velocity anomalies relative to the minimum 1-D model and a low-velocity model produced similar final velocity perturbations and rms residual values (0.056 s), even though the reference models are quite different in the top 2 km. The distribution of velocity anomalies is similar in deeper well-resolved layers (1 2 km b.s.l.), although the magnitude of the low-velocity perturbations at 2 km b.s.l. is greater for the low-velocity layer model. The low-velocity layer model produces a narrower low-velocity perturbation under the Grímsvötn caldera between sea level and 1 km a.s.l. with an absolute velocity of 2.8 km s 1 at sea level whereas the southern caldera rim, the Grímsfjall mountain ridge, is characterized by velocities in the range km s 1 at the same depth. The presence of a thin, low-velocity region at the base of the glacier thus has no significant effect on the velocity structure of the deeper layers or the epicentral distribution. Earthquakes relocate to slightly ( 0.5 km) greater depth which is within the location error. The velocity perturbations are stable because there is no absolute time reference in the inversion, with longer traveltimes through a low-velocity region being counter balanced by shifts in the computed earthquake origin times. 3.3 Resolution assessment Tests with synthetic data provide a tool for addressing the relationship between the effects that parametrization, the data set and damping have on the quality of the solution (Kissling et al. 2001). The characteristic model test has been used in this study to generate synthetic data. It is defined as containing the size and the amplitude of anomalies seen in the inversion results, but with different strike, shape and sign of amplitude (Haslinger 1999). Structures in the characteristic model are only recovered where the resolution is reliable. The characteristic model (Fig. 4a) consists of a series of positive and negative anomalies placed at three different depths (0, 2 and 4kmbelow sea level, which correspond to layers 2, 4 and 6). Synthetic data have been created using finite-difference modelling (Vidale 1990; Podvin & Lecomte 1991) with the same source receiver distribution and inverted using the same control parameters and parametrization as used to generate the final model. Due to uneven ray coverage there is a large lateral variation in resolution across the V98 array (Fig. 4b). Most of the boundaries between regions of high- and low-velocity are recovered in layers 2 and 4. Amplitude recovery is depressed due to the damping inherent in the inversion. Layer 6 also recovers partially the shapes of the anomalies but the images are blurred and amplitude recovery is poor. Strong vertical smearing is observed into layers 1 (z = 1 km), and 3 (z = 1 km) where ray paths are mainly orientated in avertical direction but less in layer 5 (z = 3 km) where many ray paths are near horizontal. To make a quantitative assessment of the reliability and accuracy of the final model, results obtained from the synthetics were compared with other diagnostic tools including hit count, derivative weighted sum (DWS) and resolution diagonal elements (RDE). Fig. 5 shows the hit count, DWS and RDE for layers 1 6 with white and black contours representing regions which are well illuminated from the tests with synthetic data. The regions with the highest hit count, DWS and RDE are contained mainly within the contours. In shallow layers (0 to 1 km above sea level), rays are mainly travelling vertically up to the receiver, giving patchy, non-zero RDE and DWS distributions. With increasing depth (0 to 2 km b.s.l.) there are more crossing ray paths which results in a more homogeneous distribution of DWS and RDE. At depths greater than 3 km b.s.l. the DWS and RDE decrease significantly due to lower ray coverage. The disadvantage of using solely resolution diagonal elements (RDE) is that because Simul2000 employs a damped least squares inversion, the resolution is dependent on the damping value (e.g. Eberhart-Phillips 1986; Kissling et al. 2001). Other elements of the resolution matrix need to be evaluated as well. Resolution contours (Reyners et al. 1999) which describe the relative size and pattern of the off-diagonal elements can be used to test the amount of image blurring within the velocity model. Resolution contours across the Grímsvötn caldera, delineating the region in which the values of the resolution matrix decay to 70 per cent of the value of the diagonal element, indicate that most nodes are well resolved, i.e. of similar size as individual grid cells. The nodes coincide with small spread values (e.g. Toomey & Foulger 1989), calculated by compressing each row of the resolution matrix. Smearing is more significant for nodes below 2 km depth, with higher spread values and resolution contours extending beyond adjacent nodes. Focusing our interpretation on the Grímsvötn region, further tests were conducted in order to assess the degree of smearing within the Grímsvötn caldera. A low-velocity anomaly was introduced from 0 2 and 2 4 km b.s.l., to test the resolution of the upper and lower part of the model. Whereas the shallow-velocity anomaly is well resolved, amplitude recovery is only 50 per cent (Fig. 6a). Regions of high velocity encompassing the low velocity anomaly are significantly higher in amplitude than edge effects observed in smearing tests (e.g. Kissling et al. 2001). At greater depth (Fig. 6b), the degree of recovery decrease due to reduction in ray coverage. Amplitude recovery of an anomaly at 3 km depth is close to 30 per cent. At 4 km depth the velocity has only been slightly perturbed from the background minimum 1-D model, indicating very low resolution. The P-wave velocity model within the Grímsvötn caldera is thus resolved to a depth of 2 3 km b.s.l. 4 RESULTS 4.1 Local earthquake tomography and focal mechanisms The main features of the 3-D P-wave velocity model is a well resolved region of lower velocities within the main Grímsvötn caldera, which is surrounded by higher velocity anomalies (Figs 7 and 8). A velocity contrast of up to 14 per cent from the minimum 1-D model (V p decrease of 0.28 km s 1 ) occurs beneath the main caldera at shallow depth, extending toward the northern caldera at 1kmb.s.l.

6 868 R. Alfaro et al. Figure 4. (a) A model used to create synthetic traveltime data. Velocity perturbations are relative to the minimum 1-D model. (b) Results of the 3-D inversion using the synthetic data set. Thin green lines show the outlines of velocity anomalies in the original model and dashed lines areas of reliable resolution, i.e. good recovery of shape and amplitude. The minimum volume of this low-velocity body is 20 km3, as defined by the boundary marked by zero velocity change relative to the minimum 1-D model. The low-velocity body within the outlet of the caldera lake, east of Grı msfjall (Fig. 7) is predominantly an artefact of the inversion at the edge of the model as there is little resolution in this region. The higher velocity flanks along the caldera rims have a maximum velocity increase of +14 per cent from the minimum 1-D model. At 2 km b.s.l., they seem to be confined to the southern and western caldera rims. This P-wave velocity structure correlates well with the density structure observed by Gudmundsson & Milsom (1997), with the low-velocity body corresponding to their low-density body within the uppermost km beneath the caldera, and higher velocity/density flanks. Seismicity beneath Grı msvo tn is confined mostly to swarms along the inferred caldera margin forming a circular region surrounding the main and north Grı msvo tn calderas. The first seismicity map based on a local array, although a snapshot, reveals clustered activity which is mainly confined to the inner rims and western flank of the main caldera with a few more scattered events in the north and east calderas. We also recorded events north of the Grı msvo tn caldera complex, within the eastward extension of the Loki Ridge (Fig. 2). Earthquake activity within northwest Vatnajo kull has been persistent within this zone in recent years (Einarsson 1991; Jakobsdo ttir et al. 2002). Small earthquakes were also recorded in the vicinity of the 1996 Gja lp eruption site. Some correlation can be seen between the location of microseismicity and geothermal activity at the southern rim of the Grı msvo tn caldera (Fig. 7). A distinctive feature of the Grı msvo tn clusters is the sharp cut-off of seismicity at km b.s.l. (Fig. 8). Clusters along the western flank extend deeper than those within the caldera. Focal mechanism solutions were obtained from impulsive P-wave arrivals of the most accurately located microearthquakes. The angles of emergence and source receiver azimuths were determined using the Simul2000 program and plotted on an equal-area projection of the lower focal hemisphere using the Fpfit program by Reasenberg & Oppenheimer (1985). Non-unique solutions, as well as poorly constrained nodal planes within regions of limited azimuthal coverage, such as northeast of Gja lp, were discarded. The focal mechanisms are characterized by both pure and oblique normal faulting (Fig. 9), reflecting an extensional regime, most likely caused by the inflation of the shallow crustal magma chamber beneath Grı msvo tn (Sturkell et al. 2003). Both nodal planes tend to C 2006 The Authors, GJI, 168, C 2006 RAS Journal compilation

7 Structure of the Grı msvo tn central volcano 869 Figure 5. Hit count, DWS and RDE for different depth layers. White and black contours mark regions which are well illuminated as derived from synthetic testing. Green circles represent earthquakes and stations are represented by triangles. be oriented subparallel to the local trend of the caldera rim, suggesting that a common mechanism (inflation) is operating for a large proportion of the events. The outward dipping nodal planes most likely represent the fault planes. Although non-double-couple radiation patterns have been observed within high-temperature geothermal areas of central volcanoes in Iceland (e.g. Miller et al. 1998), we did not observe volume related sources. Bearing in mind that the Grı msvo tn microseismicity is generally low except during periods of magmatic unrest (a station has been operated on the southern caldera rim since 1985), it is logical to infer that cooling and contraction at depth due to circulating groundwater fluids from above is not responsible for the elevated seismicity. Instead, the distinct seismic clusters, 1 2 km in width and up to 2 km in depth along the main caldera faults, and the similar waveforms of earthquakes within each cluster are associated with localized weak zones responding to the inflation. 4.2 V98 regional, teleseismic and active source residuals In order to supplement the 3-D model of shallow velocity structure from local seismicity we analysed both regional and teleseismic events recorded during the V98 experiment. Having insufficient data for a tomographic inversion of these events we could only compare their traveltime residuals with results of the LET model (Fig. 10). Regional variations in source receiver distance were accounted for by correcting arrival times at each station to a common elevation and a uniform velocity of 7 km s 1. Theoretically, observed C 2006 The Authors, GJI, 168, C 2006 RAS Journal compilation travel time delays could be introduced anywhere along the ray path between source and receiver. However, since the receiver array is relatively small, we assume that variations in travel time across the V98 array are mostly due to the local structure beneath individual stations. Teleseismic arrivals were measured by picking a consistent waveform across the array, using the first peak or trough of the Pwaveform. Absolute residuals were obtained by subtracting predicted traveltimes, calculated using the IASP91 global model, from the observed traveltimes. Relative residuals at a common level across the array were then calculated by subtracting the mean residual for each event from the absolute residuals of every station for that event and by correcting arrival times at each station to a common elevation using the method of Achauer et al. (1986). P-arrivals of nine teleseisms exhibiting good signal-to-noise ratios were analysed. All events had magnitudes 6.0. Source backazimuths are from all directions apart from the south. The angles of incidence of most rays beneath the Grı msvo tn caldera were within of vertical. Having a dominant frequency of 1 Hz, the teleseismic P-waves illuminated a zone close to 2 km in diameter at 4 km depth beneath the caldera. The most noticeable feature from the regional and teleseismic travel time analysis is a region of positive residuals (low-velocity) beneath Grı msvo tn. Although broadly comparable in dimensions with the LVZ defined by the LET results, the magnitudes of the travel time anomalies from regional and teleseismic arrivals are much larger than those from the LET (Fig. 10a c). A delay of s is observed for the caldera stations using regional events from

8 870 R. Alfaro et al. Figure 6. Testing the reliability of a low-velocity body within the Grímsvötn caldera using synthetic data from an anomaly at 0 2 and 2 4 km depth (top three plots). White circles denote local events, projected onto the two E W cross-sections. Good recovery of shape and amplitude are observed along both cross-sections (lower four plots). Hengill (southwest), Askja (northeast) and Katla (south-southwest). Teleseismic delays also range from 0.06 (V5) to 0.22 s (V23). These delays are more pronounced than the s local delays, indicating that ray paths under the caldera are traversing an anomalously low-velocity body below the region sampled by LET. Higher velocities which cause arrivals to be up to 0.1 s early on the southern caldera flank are also in agreement with LET results. Arrival times from two offshore detonations further constrain the extent of low-velocity zones beneath the V98 array. P-wave travel time delays of s were observed along Line 1 (Figs 1, 10d and 11). The traveltimes were modelled using the Xrayinvr ray tracing software (Zelt & Smith 1992). A simplified P-wave velocity model based upon the nearest section of the ICEMELT profile (Darbyshire et al. 1998) is used as regional structure and the LET model added in the uppermost 4 km beneath the Grímsvötn caldera. However, this model does not match the delays observed in other crustal P-wave arrivals. A further low-velocity body is required at depths greater than 3 4 km. 5 DISCUSSION 5.1 The shallow structure of Grímsvötn Seismic P-wave velocities can be altered by a wide variety of mechanisms including changes in temperature and ambient pressure, phase changes in the material and composition of the rocks, degree and style of fracturing and nature and degree of saturation of fluid content such as water and/or melt. A reduction in P-wave velocity alone, without information from S-waves and other geophysical data, cannot be used to identify the contribution made by each of these factors without some ambiguity. As in other parts of the Neovolcanic zones of Iceland, the uppermost 2 3 km beneath the Grímsvötn caldera is likely to be made up of alternating layers of basaltic lava flows (pillow basalts) and hyaloclastites with a progressive increase of intrusives with depth. Lithological cross-sections from boreholes within high-temperature geothermal areas contain numerous units of these rocks corresponding to glacial and interglacial sequences. Low near-surface velocities ( km s 1 ) characterize the Neovolcanic zones (Flóvenz & Gunnarsson 1991; Brandsdóttir et al. 1997; Weir et al. 2001). However, the region imaged by LET within the Grímsvötn caldera has markedly lower velocities in the uppermost 2 3 km than the Krafla caldera in northern Iceland (Brandsdóttir et al. 1997), a likely product of the higher eruptive activity and accumulation of volcanoclastic rocks within the Grímsvötn central volcano. Positive Bouguer anomalies with overall amplitudes of mgal and centralized lows are associated with calderas in Iceland (Gudmundsson & Högnadóttir 2002). The main caldera of Grímsvötn has a relative low of 5 mgal amplitude superimposed on a gravity high of up to 20 mgal. Gudmundsson & Milsom (1997)

9 Structure of the Grı msvo tn central volcano z = -1.0 km 871 z = 0.0 km N N 64.4 N GFL GFL z = 1.0 km z = 2.0 km N N 64.4 N GFL W GFL W % Vp change, rel. to minimum 1D mode l Figure 7. Percentage velocity perturbations relative to the minimum 1-D model for different depth layers. Regions in red have lower velocities than the initial model, regions in blue have higher velocities. Dashed lines delineate areas of reliable resolution, i.e. good recovery of shape and amplitude. Topographical ice contours in light grey have a spacing of 50 m. Node spacings are marked by black crosses (2 km spacing in the horizontal direction), seismometers are represented by triangles and green circles are microearthquakes between each depth grid node (0.5 km either side of the grid node). The 1998 eruption site is shown by the solid black line. N = north caldera, GFL = Grı msfjall ridge. Areas of surface geothermal activity are shown by black diamonds. put forward two possible explanations for the low-density material. It could be mainly volcaniclastic material that has accumulated within the caldera, or it could be directly related to the caldera subsidence as a 0.4 km downwards displacement could explain a negative density contrast of 0.2 Mg m 3. Comparison of the gravity map with our velocity structure shows good correspondence between low-velocity and low-density regions. The low-velocity region is considerably thicker within the caldera than on the flanks (Figs 7 and 8). Based on density contrasts the low-density material has a volume of km3, which is similar to the low-velocity body imaged by LET (20 km3 ). In addition to a marked low-velocity body within the main caldera lower velocities appear SE, E and N of Grı msvo tn (Fig. 7). These lower velocities are not well constrained but they may indicate that the uppermost crust flanking Grı msvo tn has slightly lower velocities than the minimum 1-D model. The top few hundred metres within the main caldera have velocities of km s 1, consistent with unconsolidated or loosely consolidated water-saturated sediments. Between 0.5 and 1 km depth (Fig. 8) consolidated hyaloclastites with gradually increasing percentage of intrusives could explain the gradual rise in velocity from 2.5 km to 4 km s 1. Batches of melt under the main caldera could explain the pronounced shear-wave attenuation observed within the Grı msvo tn caldera (Figs 12 and 13) and discussed below. Such magma batches could be located below 1 km b.s.l., the maximum depth of earthquakes within the main caldera. C 2006 The Authors, GJI, 168, C 2006 RAS Journal compilation Figure 8. (a) Plane views of velocity structure within the Grı msvo tn caldera. Microearthquakes within the main region of influence in each layer are marked by white circles. Contours represent ice topography in metres. (b) E W cross-section along the dashed line in (a). 5.2 The Grı msvo tn magma chamber Evidence to suggest that a magma chamber exists under Grı msvo tn comes from a variety of sources (Bjo rnsson et al. 1982; Einarsson & Brandsdo ttir 1984; Gudmundsson & Milsom 1997; Larsen et al. 1998). The Grı msvo tn caldera lake acts as a natural calorimeter of the geothermal (magmatic) activity within the caldera (Bjo rnsson et al. 1982; Bjo rnsson & Gudmundsson 1993). An 80 yr record of heat release through measurements of meltwater accumulation and drainage testifies that the geothermal areas are being frequently replenished by dyke intrusions into the shallow crust. Bjo rnsson & Gudmundsson (1993) estimated that the total heat release in was equivalent to the energy released by the solidification and

10 872 R. Alfaro et al. a) Local delays W W N 64.4 N b) Regional delays W W Figure 9. Lower hemisphere focal mechanism solutions computed using the Fpfit program (Reasenberg & Oppenheimer 1985) for microearthquakes around the Grímsvötn caldera rim following 3-D relocation. Black quadrants indicate compression, white quadrants dilatation N N cooling of 2.1 ± 0.4 km 3 of basaltic magma ((8.1 ± 1.6) J), with a contribution of 45 per cent (max.) from a magma chamber, 35 per cent (min) from shallow intrusions and 20 per cent from eruptions. High velocity/density bodies (basic intrusives) have been identified within both active and extinct/eroded central volcanoes in Iceland (Pálmason 1971; Flóvenz & Gunnarsson 1991; Gudmundsson et al. 1994; Brandsdóttir et al. 1997). A slight increase in velocities is also observed along the flanks of Grímsvötn (Figs 10 and 11). The 20 mgal Bouguer gravity high across Grímsvötn has been interpreted to be caused by 400 km 3 of basic intrusives (gabbros) at a depth between 1.5 and 4 km beneath the main caldera, extending toward the surface along its flanks (Gudmundsson & Milsom 1997). This body is essentially nonmagnetic, indicating that its temperature may be above the Curie point or that it has undergone severe hydrothermal alteration. Although a magma chamber beneath the main caldera of Grímsvötn is not required by the gravity and magnetic data, a 10 km 3 chamber (ca. 1 km thick and 4 km across) close to 2 km b.s.l. is consistent with these observations. The pattern of seismicity preceding and following the 1983 Grímsvötn eruption suggest that a magma chamber, located somewhere beneath the caldera, inflated, deflated and then re-inflated (Einarsson & Brandsdóttir 1984; Björnsson & Einarsson 1990). Single station GPS measurements by Sturkell et al. (2003) indicate that a magma chamber, located at least 1.6 km depth beneath the caldera complex, deflated briefly during the relatively small (0.1 km 3 ) December 1998 eruption. The GPS data indicates that the 1998 eruption was fed from a magma body located within the northern margin of the low-velocity region in the final LET model (Fig. 7). However, the location of this magma body should be treated with caution since data from only one GPS station was available and a single source of inflation and deflation was assumed. A more complicated magma storage region than can represented by a single Mogi source may be present and crustal heterogeneity can deflect the displacement vector. Based on the GPS data, increased microseismicity within the caldera during the months preceding the December 1998 eruption c) Teleseismic delays V W W d) Shot 2 delays V W W B0 B47 V1 V24 Gríms fjal l B1 V24 V3 V3 V23 V N 64.4 N V22 B46 V N 0.10 s residual N Figure 10. Velocity variations across the Grímsvötn caldera from local events (green circles) (a) relative to the minimum 1-D model between 1 km above sea level to 2 km b.s.l. A positive residual (delayed arrivals) marked in red, indicates a region underlain by anomalous low-velocities, while a negative residual (blue) implies higher than normal velocities. (b) Average traveltime residuals from regional events from Hengill, Askja and Katla. (c) Average delays from nine teleseismic events recorded during the V98 deployment. (d) Delays from Shot 2. Location of the 1998 December eruption site is marked by a bold black line in (c).

11 Structure of the Grı msvo tn central volcano 873 (a) Caldera Depth (km) Shot2 V5 6.7 Hengill Time - Distance / 6.5 (s) (b) 25 V km s 1 PmP/SmS Caldera V22 V4 P/S Teleseismic raypaths V N ~ 4km s V41 V23 V. E. 1:1 V24 Distance to Shot2 (km) V N V km V W W SmS 15 Figure 12. S-wave ray paths (grey lines) are mostly confined to the southwestern and western rims of the main caldera. V23 10 S 5 PmP P B04B03V05V03 Caldera Time - Distance / 6.5 (s) (c) Distance to Shot1 (km) SmS 15 S 10 5 PmP P Distance to Shot2 (km) Figure 11. (a) E W cross-section showing rays traced by Xrayinvr from Shot 2 (P and P m P) and from a Hengill earthquake located around 180 km WSW of Grı msvo tn. Teleseismic ray paths are also shown. The box around the Grı msvo tn caldera is shown at a larger scale in Fig. 14. (b) E-W component record section from Shot 1 along Line 1, bandpass filtered 1 5 Hz. Clear S-waves are visible in the vicinity of the shot until station V23 at around 110 km offset. The P waves are strongly attenuated at greater distances and no direct S-waves are visible. Traveltime curves for P and P m P are from Xrayinvr modelling. S and S m S traveltime curves are derived by multiplying corresponding P and P m P times by (c) E W component record section from Shot 2 along Line 1, bandpass filtered 1 10 Hz. No S-waves are observed west of the caldera but clear S m S arrivals are seen at B03 and B04. Possible lower amplitude S m S arrivals are observed at stations V03 to V05. was most likely generated by strain changes within the region above and adjacent to the inflating magma chamber. The relatively minor increase in seismicity prior to the December 1998 eruption may indicate that the crust within the caldera is only moderately seismogenic, i.e. that the current state of stress within the uppermost crust is low. The Icelandic regional SIL-network only recorded eight events within Grı msvo tn during the first 11 months of 1998, ranging in magnitude from M1.6 to M3.1, most being around M2. A total of 35 events, with magnitudes ranging from 1 to 3.5, accompanied the dyke intrusion along the southern caldera rim at the outbreak of the eruption. Based on the seismic distribution and increase in geothermal activity along the southern caldera rim following the eruption, the feeder dyke may have been 4 6 km long C 2006 The Authors, GJI, 168, C 2006 RAS Journal compilation Figure 13. Multiple north-component record section of events at the western rim of the caldera (reduced traveltime with a reduction velocity of 4.0 km s 1 ). No bandpass filter is applied. (a) Waveforms at stations located on the western rim of the caldera exhibit strongly attenuated S-wave arrivals. (b) Stations within or northeast of the main caldera have severely attenuated P-waves and are devoid of S-waves. and 2 3 km high. A marked decrease in microearthquake activity following the eruption can also be taken as an indicator of a relaxed state of stress. This activity bears close resemblance to the cycles observed during the Krafla rifting episode where seismicity associated with the inflating magma chamber decreased markedly during brief deflation events and only reappeared when the magmatic stress had again overcome the lithostatic/regional stress (e.g. Einarsson 1991).

12 874 R. Alfaro et al. Figure 14. Summary E W cross-section showing the ray paths under the Grímsvötn caldera from an earthquake in the Hengill region (H), a teleseismic event (T) and from Shot 2 (P and P m P). The source of the delays observed in Fig. 11 is dependent on the depth sampled by individual ray paths. The velocity model shown is a hybrid model, containing a low-velocity body representative of a sill of pure melt with an underlying 1.5 km thick zone within which velocities increase with depth from values typical of a high percentage of partial melt to values consistent with normal Icelandic lower crust. See text for further discussion. Grey-shaded areas indicate zones of negative velocity contrast relative to the background velocity of the surrounding crust at that depth. Dark grey indicates a high negative velocity contrast, lighter shades of grey indicate progressively smaller negative velocity contrasts. P-wave velocities for the background crustal structure are shown in km s 1. Results from the V98 experiment provide further evidence for the existence of magma at a shallow level beneath Grímsvötn. Delays in P-wave traveltimes from local, regional and teleseismic earthquakes, as well as from the two explosive shots, are consistent with the existence of a magma chamber beneath the Grímsvötn caldera. Delays from the two shots and from the regional and teleseismic earthquakes are s larger than can be explained by the LET model alone within the uppermost 3 km under the caldera (Fig. 10). The region where ray paths from these different events intersect (Fig. 11a) suggests that a low-velocity-body must lie within a depth of 3 7 km beneath the ice surface. Microearthquakes within the main caldera were mostly confined to 1 3 km depth during V98 (Fig. 8), i.e. within the region above the magma chamber. An additional observation is the region of attenuated shear waves from local earthquakes and passive sources within the caldera (Figs 11 13). Clear S-waves are observed from shot 1 at all stations east of and including station V23 (Fig. 11b). Within and to the west of the caldera, however, these S-waves are completely attenuated. Clear Moho reflections (S m S) are seen west of the caldera at stations B04 and B03 (Fig. 11c). Possible attenuated S m S phases are also observed between stations V05 and V03. These observations of clear shear energy on either side of the caldera, together with the high attenuation observed within and immediately west of the caldera, are best explained by a melt lens centred beneath the main caldera, and a zone of high percentage partial melt (Fig. 14). The clear shear energy observed at V23 and the attenuated S m S phase at V03 constrain the maximum lateral extent of any pure melt lens to be 7 8 km in an east west direction between these two stations. Two end-member cases may be considered for the shape of the magma chamber; a sill-like body containing pure/high percentage partial melt and a thicker gradational zone consisting of pure/high percentage partial melt at the roof and little or no melt at the base. If one assumes the first end-member, a magma chamber containing pure/high percentage partial melt with a velocity of km s 1 (Murase & McBirney 1973), then the thickness of the magma chamber must be 1kminorder to account for the observed s regional, teleseismic and shot delays. Using the distribution of local seismicity and shear wave attenuation we estimate the maximum lateral extent of the magma chamber to be 7 8 km E W and 4 5kmN S. If there is also a region of partial melt extending over a greater depth zone, then the magma chamber containing pure melt could be correspondingly smaller. In comparison, the Katla magma chamber has been inferred to be 1 kmthick and its volume 10 km 3 with at least 5 km 3 of melt (Gudmundsson et al. 1994). A somewhat larger (12 54 km 3 ) and thicker ( km) region of pure/partial melt was imaged under Krafla (Brandsdóttir et al. 1997). If there is a high-velocity region deeper beneath the inferred Grímsvötn magma chamber, as observed beneath Krafla and Katla, then the magma chamber itself could be somewhat thicker, since traveltime delays caused by the low velocities of the magma chamber would be partly compensated by the high velocities underneath. However, it should be noted that below Grímsvötn a high-velocity region is not resolvable by the seismic data available. Alternatively, we can assume that the magma chamber grades downwards from pure/high percentage partial melt at its roof to little or no melt at its base. This situation is represented by P-wave velocities of km s 1 at the top of the magma chamber, increasing to the background velocity of the adjacent crust at the base. In this case the region of lowered velocities presumed to contain partial melt may be 2kmthick in order to lie in the region containing delayed ray paths (Fig. 14). The seismic data currently available is not sufficient to allow us to discriminate between these two end-member models. The Grímsvötn magma chamber may be a hybrid of the two end-member cases of pure melt and distributed partial melt. A thin pure melt lens, perhaps m thick, could be underlain and flanked by a zone of high percentage partial melt. This region of high melt, of 0.5 km thickness, could in turn be underlain by a 1 km thick mush zone, grading downwards from a high degree of melt at its roof to no melt at its base. Such a structure would account for all observed P-wave delays as well as the attenuation of shear energy from the explosive shots and the gravity and magnetic data. It is also broadly consistent with models of magma plumbing and crustal accretion at both the Mid-Atlantic Ridge (Henstock et al. 1993; Magde et al. 2000) and in Iceland (Gudmundsson et al. 1994; Brandsdóttir et al. 1997; Menke et al. 1998), indicating the presence of hot material advecting downwards and outwards beneath the magma chamber during spreading. 6 CONCLUSIONS Using local earthquake tomography a 20 km 3 region of lowered velocities has been imaged below the main Grímsvötn caldera to a depth of 3 kmbelow the ice surface. The low velocities are inferred to be caused by higher amounts of unconsolidated material with recent intrusives and are flanked by high velocities along the caldera rim.