Magma storage and migration associated with the 2011-2012 ...digital.csic.es/bitstream/10261/81685/1/IP_2013_JGR_old.pdf · Gonzalez et al., Magma dynamics El Hierro eruption Page

Gonzalez et al., Magma dynamics El Hierro eruption Page \1

Magma storage and migration associated with the 2011-2012 El 1

Hierro eruption: implications for crustal magmatic systems at 2

oceanic island volcanoes 3

4

Pablo J. González1,*

, Sergey V. Samsonov2, Susi Pepe

3, Kristy F. Tiampo

1, Pietro Tizzani

3, 5

Francesco Casu3, José Fernández

4, Antonio G. Camacho

4 and Eugenio Sansosti

3 6

7

1Department of Earth Sciences, Western University, Biological and Geological Sciences 8

Building, 1151 Richmond Street, London, ON, N6A5B7, Canada. 9

2Canada Centre for Remote Sensing, Natural Resources Canada, 588 Booth Street, Ottawa, ON 10

K1A0Y7, Canada. 11

3 Istituto per il Rilevamento Elettromagnetico dell’Ambiente, National Research Council (CNR) 12

of Italy, Via Diocleziano 328, Naples, Italy. 13

4Instituto de Geociencias (CSIC, UCM), Fac. Ciencias Matemáticas, Plaza de Ciencias, 3, 14

28040-Madrid, Spain. 15

* corresponding author 16

17

RUNNING TITLE: MAGMA DYNAMICS AT 2011-2012 EL HIERRO ERUPTION 18

Keywords: Magma migration, Volcano deformation, Radar interferometry, Time 19

inversion, Oceanic volcanic islands 20

Submitted: Journal of Geophysical Research 21

jft

Texto escrito a máquina

(JGR, in press, 2013, 2013JB010184)


ABSTRACT 22

Starting in July 2011, anomalous seismicity was observed at El Hierro Island, a young oceanic 23

island volcano. On the 12 October 2011 the process led to the beginning of a submarine NW-SE 24

fissural eruption at ~15 km from the initial earthquake loci, indicative of significant lateral 25

magma migration. Here, we conduct a multi-frequency, multi-sensor interferometric analysis of 26

spaceborne radar images acquired using three different satellite systems (RADARSAT-2, 27

ENVISAT and COSMO-SkyMed). The data fully captures both the pre-eruptive and co-eruptive 28

phases. Elastic modeling of the ground deformation is employed to constrain the dynamics 29

associated with the magmatic activity. This study represents the first geodetically-constrained 30

active magmatic plumbing system model for any of the Canary Islands volcanoes, and one of the 31

few examples of submarine volcanic activity to date. Geodetic results reveal two spatially 32

distinct shallow (crustal) magma reservoirs, a deeper central source (9.5 ± 4.0 km) and a 33

shallower magma reservoir at the flank of the southern rift (4.5 ± 2.0 km). The deeper source was 34

recharged, explaining the relatively long basaltic eruption, contributing to the observed island-35

wide uplift processes and validating proposed active magma underplating. The shallowest source 36

may be an incipient reservoir that facilitates fractional crystallization as observed at other Canary 37

Islands. Data from this eruption supports a relationship between the depth of the shallow crustal 38

magmatic systems and the long-term magma supply rate and oceanic lithospheric age. Such a 39

relationship implies that a factor controlling the existence/depth of shallow (crustal) magmatic 40

systems in oceanic island volcanoes is the lithosphere thermomechanical behavior. 41

42

43

44


1. Introduction 45

A deeper understanding of the magma plumbing system is critical to our ability to forecast the 46

eruptive behavior of a volcano and to provide a wider framework to constrain dynamic volcanic 47

processes, such as magma storage, supply rate, and migration [Takada, 1989; Poland et al., 48

2012; Clague and Dixon, 2000, Fialko and Rubin, 1998]. Attempts to understand the structure 49

and dynamics of magma plumbing systems at basaltic intraplate oceanic island volcanoes rely 50

heavily on a handful of well-studied volcanoes, e.g., Kilauea and Piton de la Fournaise [Tilling 51

and Dvorak, 1993]. However, the lack of data at other volcanoes makes it difficult to extrapolate 52

a well-constrained magma plumbing systems and dynamics that may be used as prototypes for 53

all oceanic island volcanoes. In particular, less active volcanoes such as the Canary or Cape 54

Verde Islands might not be described accurately by plumbing systems inferred from highly 55

active volcanoes. 56

57

An image of the active magmatic plumbing system of a volcano can be obtained via the 58

interpretation of various geological, geophysical and geochemical data, including seismic, 59

experimental petrology, gas emissions chemistry and geodetic measurements [Dzurisin, 2006]. 60

To study the long-term evolution of magmatic systems, experimental petrology provides a very 61

valuable tool through the analysis of erupted materials [Marsh, 1996; Hansteen et al., 1998; 62

Longpré et al., 2008]. However, at active volcanic areas, geophysical observations are an 63

important complement, providing constraints on magma locations and volumes. Geodetic 64

imaging is useful to infer complex deformation processes controlled by crustal rheology. In 65

volcanic environments, the temperature and density contrasts affect the stress distribution which 66

controls the dynamic magmatic processes. Nonetheless, under certain assumptions (usually 67


brittle/elastic deformation), it is possible to evaluate magma volume storage and migration 68

beneath a volcano [Dzurisin, 2006; Poland et al., 2012]. Therefore, the ground deformation 69

associated with an eruption allows us to constrain the magma plumbing systems, in particular at 70

basaltic intraplate oceanic island volcanoes with low eruption rates, and compare its structure 71

and dynamics against the prototypes. 72

73

Studies of the magma plumbing systems at basaltic intraplate ocean island volcanoes (e.g., 74

Reunion, Galápagos, Hawaiian, Canarian and Cape Verde Islands) have revealed fundamental 75

differences in their shallower (crustal) levels. It has been postulated that shallow (crustal) magma 76

chambers are not always present and differ in magmatic processes. At Hawaii, magma chambers 77

stagnate and distribute magma along elongated rift zones [Tilling and Dvorak, 1993]. However, 78

at Tenerife and Gran Canaria Islands, crustal magma chambers develop as long-lived features 79

that can fractionate parental basanitic composition melts into end-member (highly explosive) 80

phonolitic magmas [Ablay and Martí, 2000]. Nevertheless, multiple processes such as flank 81

instability also can play a role in the absence/presence of magma reservoirs by occasionally 82

disrupting the volcanic plumbing system [e.g. Amelung and Day, 2002; Ruch et al., 2012]. 83

Therefore, a deeper understanding of the crustal portion of the plumbing systems of oceanic 84

island volcanoes is essential for improving volcanic risk analysis. 85

86

A submarine eruption off El Hierro Island started in October 2011 and ended in early March 87

2012 [López et al., 2012]. The 2011-2012 El Hierro eruption represents the first eruptive episode 88

in 40 years, and the first systematically well-monitored eruption in the Canary Islands. 89

Consequently, the analysis of data acquired during the 2011-2012 El Hierro eruption is a unique 90


opportunity to image, for the first time, the active magmatic system of one of the Canary Islands. 91

In addition, it provides the first opportunity to study the volcanic phenomena associated with an 92

on-shore to off-shore migration of magma that culminated in to an eruption. This study also 93

contributes to our knowledge of the magmatic systems of oceanic island volcanoes. 94

95

96

2. El Hierro Island and the 2011-2012 eruption 97

El Hierro Island is small basaltic oceanic island volcano with a subaerial edifice ~20 km in 98

diameter and steep relief, with maximum altitudes of 1500 m (Pico de Malpaso) above the 99

surrounding seafloor and bathymetry of approximately 3600-3800 m. Located in the Canary 100

Islands volcanic archipelago, off the African continent (Figure 1), El Hierro is the south-101

westernmost and the youngest island, at 1.1 Ma old [Carracedo et al., 2001]. Geologically, El 102

Hierro is composed of three overlapping units (Tiñor, El Golfo, and rift volcanism) which have 103

suffered four flank collapses [Masson et al., 2002; Muenn et al., 2006]. The oldest formation is 104

the Tiñor shield volcano, with activity spanning 1.2 – 0.88 Ma. It has been inferred that this 105

volcano collapsed northwards. Later, another shield volcano, El Golfo, completely covered the 106

Tiñor collapse [Carracedo et al., 2001]. Eventually, El Golfo volcano again collapsed towards 107

the north, forming the El Golfo flank collapse (87-39 ka), a prominent embayment in the north 108

(Figure 1) [Longpré et al., 2011]. Over the last 158 ka, volcanism has been presented as effusive 109

along three major rift zones (northeastern, west-northwestern, and southern), with the 110

development of strombolian cinder cones and associated lava flows (Figure 1). At El Hierro, 111

Holocene volcanism is limited to a few dated eruptions; Tanganasoga volcano, a composite 112

volcano that erupted crystal-rich ankaramite lavas, and the Montaña Chamuscada volcano, 4000 113


and 2500 years before present, respectively [Guillou et al., 1996; Manconi et al., 2009]. For the 114

last 500 years, which corresponds with the historical record, no volcanic eruptions have been 115

confirmed. In 1793, a felt seismic crisis was dubiously associated with the eruption of the Lomo 116

Negro volcano, northwest of the Island [Carracedo et al., 2001; Romero and Guillén, 2012]. 117

Over the past decade, seismicity has remained low, with an average seismicity rate of <10 118

earthquakes/year. In 2004, a minor increase in seismic activity was correlated with a 119

simultaneous increment in CO2 soil diffuse efflux [Padrón et al., 2008]. 120

121

On 17 July 2011, the two permanent seismic stations of the Spanish Geographical Survey, IGN 122

[www.ign.es] detected unusual activity (Figure 2). From 17 July 2011 to early September small 123

magnitude seismicity, mbLg<3 (Figure 3a and 3b), occurred in small clusters at a depth range 124

from 9 to 15 km and varying slightly in position beneath the Tanganasoga volcano area (see 125

Figures 1, 2 and 3a-c). At the initial stage in late July, seismicity migrated generally towards the 126

west (Figure 3a). This trend shifted slightly towards the northwest in early-to-mid August 127

(Figure 3b). By mid-August, activity returned to the initial area (dark circles in Figure 3b), and 128

remained stationary for next the 15 days or so. Only one publicly-available GPS station (FRON) 129

was operating prior to the onset of the seismic swarm, located in an area northeast of the located 130

epicenters in El Golfo collapse valley (Figure 1). By the end of August, the GPS precise 131

positioning indicated a cumulative uplift of ~4 cm and ~2 cm motion towards the northeast. In 132

mid-September 2011, the seismicity began to migrate clearly southwards. Simultaneously, the 133

magnitude of events increased and earthquakes began to be strongly felt by the population, mbLg 134

~ 3-4 (Figure 3c). In late September to early October, the seismicity clearly stopped its 135

http://www.ign.es/


southwards migration and spread around a wide area, ~5 km in diameter, beneath the Mar de las 136

Calmas (Figure 1, 3c and 3d). 137

138

An energetic seismic swarm began late on 5 October 2011, peaking at a maximum mbLg=4.4 139

(Mw=4.0) on 8 October 2011 at 20:35 UTC, with a moment tensor and centroid at 6 km depth 140

[www.ign.es accessed 27 March 2012]. It was followed by a relatively quiet period of ~32 hours, 141

during which a few tens of low magnitude shallow events were located at depths of one to six 142

km. This period ended with the emergence of a growing amplitude tremor signal (10 October 143

2012, ~4:15 UTC). The tremor amplitude increased for ~24 hr, including clipping of some 144

seismometers [www.ign.es]. The tremor signal abruptly ceased after a peak in high frequency at 145

14:30 UTC on 12 October 2011, followed by several high frequency peaks during the next ~24 146

hr. On 15 October 2011 the first floating material began to appear off the coast near La Restinga 147

town (Figure 1). The material varied from a mixture of mafic and white pumice to large basanite 148

floating balloons [Troll et al., 2012; Carracedo et al., 2012]. Between 9 and 26 October 2011, 149

the submarine eruption continued to be associated with very minor seismicity (a few 150

earthquakes/day) and the magnitudes remained low, mbLg<3 (Figure 2c, and 3d). Subsequently, 151

very deep seismicity (20-30 km) appeared along the northern coast (El Golfo), accompanied by 152

visible changes at the eruption site such as volcanic bombs and bubbling associated with 153

vigorous degassing (Figure 3d and 2e). For the next three months, eruptive seismicity was sparse 154

and without clear patterns in either space or time (Figure 3f). While visible sea surface 155

discoloration disappeared in late December, by early January basaltic floating bombs re-appeared 156

and continued throughout February in a discontinuous fashion. In early March, the absence of 157

visible effects on the sea surface and the lower levels of seismicity triggered official declaration 158

http://www.ign.es/

http://www.ign.es/


of the end of the eruption on 5 March 2012 159

[www.gobiernodecanarias.org/dgse/noticias_sismo_hierro.html]. A more complete 160

comprehensive qualitative description of the chronology of the crisis can be found in Carracedo 161

et al., [2012] and López et al., [2012]. 162

163

Prior to July 2011, the geodetic monitoring of El Hierro Island was composed of two GPS 164

stations managed by different institutions. GRAFCAN (Cartográfica de Canarias) managed one 165

located in El Golfo valley with open data access (Frontera, FRON station) that was installed in 166

2010 for surveying purposes. A second station was installed in 2004 by ITER-INVOLCAN 167

(Instituto Tecnológico y Energías Renovables-Instituto Volcanológico de Canarias) in 168

cooperation with Nagoya University [González et al., 2006]. During August 2011, this minimal 169

network was augmented with five well-distributed GPS stations by INVOLCAN [Sagiya et al., 170

2012]. Meanwhile the Spanish IGN installed four GPS stations along a single transect in El 171

Golfo area [López et al., 2012]. Although this GPS profile distribution provided some 172

information with which to constrain depth and source geometry, it is not optimized to track the 173

location (and migration) of deformation sources due to its small aperture and the narrow back-174

azimuths of the displacement GPS vectors. The addition of five continuous GPS and three survey 175

GPS benchmarks in late September 2011 provided improved coverage of the entire island. In 176

addition to the GPS networks, tiltmeter and microgravity networks were installed to monitor 177

ground deformation and gravity changes [Arnoso et al., 2012]. Here, we complement the ground-178

based geodetic network with comprehensive processing, analysis and modeling of multiple 179

space-based radar interferometric data sets (RADARSAT-2, ASAR-ENVISAT and COSMO-180

SkyMed) in order to understand the dynamics of the magmatic system. 181

http://www.gobiernodecanarias.org/dgse/noticias_sismo_hierro.html


182

183

3. Data analysis: InSAR observations 184

Interferometric Synthetic Aperture Radar (InSAR) is a geodetic technique that measures surface 185

ground deformation between two different satellite passes over wide areas [e.g., Dzurisin, 2006; 186

Sansosti et al., 2010]. InSAR measurements are sensitive to a projection of the full 3D 187

displacement vector onto the satellite line-of-sight (LOS). Ideally, it is desirable to image the 188

area of interest with more than one satellite track (ascending and descending passes, with 189

variable incidence angles) in order to measure more than one projection of the full displacement 190

vector and, therefore, to better constrain models representing geophysical processes. 191

192

We used 32 SAR images from six different satellite tracks to image the ground deformation at El 193

Hierro Island. SAR images were acquired between December 2009 and March 2012 (Table S1). 194

Images were acquired from two ascending tracks of the Canadian Space Agency (CSA) 195

RADARSAT-2 satellite, beams S6 (look angle, 44º) and S7 (47º); and two descending tracks of 196

the European Space Agency (ESA) ENVISAT satellite, tracks 109 in I6 mode (36º) and 339 in I2 197

mode (20º). RADARSAT-2 and ENVISAT satellites are equipped with right-looking C-band 198

SAR antennas operating at wavelength of ~5.6 cm. In addition, images were acquired from 199

ascending and descending tracks of the Italian Space Agency (ASI) COSMO-SkyMed satellite, 200

both corresponding to the beam H4-8 (34º). The latter system is an advanced X-band SAR 201

constellation with 4 satellites operating at a wavelength of ~3.1 cm. 202

203


Interferograms were processed in two-pass differential mode, using a 30-m-resolution digital 204

elevation model (DEM) derived from the SRTM Topography mission, and a 5-m-resolution 205

DEM from IGN to remove topographic effects from C-band and X-band SAR data, respectively. 206

ENVISAT-ASAR data were processed using Doris software [Kampes et al., 2003], 207

RADARSAT-2 data were processed using GAMMA software [Wegmuller and Werner, 1997], 208

COSMO-SkyMed data were processed at CNR-IREA using their internally developed 209

interferometric SAR processing chain [Berardino et al., 2002]. Overall, we obtained 63 short 210

baseline differential interferograms (Table S2). Computed interferograms have temporal 211

separations ranging from three days to approximately two years. 212

213

El Hierro Island presented some challenges for InSAR analysis, including steep slopes, densely 214

vegetated areas and significant topographic elevation changes. The latter introduces strong 215

topography dependent phase variations due to temporal changes in atmospheric temperature 216

and/or pressure profiles (Figure S1), as observed in other islands of the Canary Islands 217

archipelago [e.g., González et al., 2010]. To compensate for this effect, we applied a commonly 218

used strategy based on an empirical polynomial relationship between atmospheric phase noise 219

and topography. We noted that a simple linear model is insufficient to remove most of the 220

atmospheric signal, a result that is in agreement with the non-linear relationships necessary to 221

model stratified troposphere signals [Doin et al., 2009]. Accordingly, we iteratively increased the 222

polynomial degree of the model up to 4. We selected an area assumed to be less affected by 223

deformation, in the northeastern portion of the island, to fit the data. To accept the optimal 224

model, at each iteration we computed an F-statistic ratio to test the significance of the decrease in 225

misfit, 2, between the two successive models [Stein and Gordon, 1984]. Another significant 226


disturbance is related to the imperfect knowledge of the satellite tracks. In this case, first order 227

artifacts can be compensated for by removing a linear or a quadratic ramp in both azimuth and 228

range image directions [Pepe et al., 2011]. Therefore, before interpretation, each interferogram 229

was corrected for orbital effects and atmospheric artifacts. This operation is critical because these 230

disturbances can bias the geophysical interpretation if attributed to deformation signal. Examples 231

of the original, modeled and residual interferograms are shown in Figure S1. 232

233

After the correction, most of interferograms did not show large signals for most of the analyzed 234

period (e.g., see the RADARSAT-2 interferogram in Figure 4a). However, clear and intense 235

deformation is detected in all interferograms spanning May 2011 to November 2011. Figure 4b-e 236

illustrates some of the interferograms with clear deformation patterns. Between 5 May 2011 and 237

8 August 2011 deformation observed with a RADARSAT-2 ascending track indicates a circular 238

pattern of deformation in the center of the island, with a maximum of ~10 cm of motion towards 239

the satellite, south of Tanganasoga volcano (Figure 4b). In a subsequent period, Figure 4c 240

displays the deformation (maximum ~10 cm towards the satellite) from a descending ENVISAT 241

interferogram spanning the entire month of September and a few weeks after the eruption started 242

(31 August 2011-30 October 2011). The intense deformation pattern is indicative of an 243

expansion source located offshore at Mar de Las Calmas. A slightly noisy descending CSK 244

interferogram spanning 12 days before the beginning of the eruption shows that most of the 245

observed deformation in Figure 4c occurred during September (Figure 4d). In Figure 4e, an 246

independent ascending interferogram (RADARSAT-2 S7) confirms the two areas of 247

deformation, a wide deformation centered in the island and a more intense and shorter 248

wavelength deformation located at northwest of the Mar de las Calmas. Finally, slight inflation 249


(motion towards the satellite) can be distinguished in a few interferograms spanning the 250

November 2011 period (e.g., the ENVISAT interferogram in Figure 4f). 251

252

253

4. Modeling: magma plumbing system and its temporal evolution 254

To analyze the large wealth of computed interferograms, we propose a two-step approach with 255

the aim of retrieving the temporal evolution of the ground deformation that is consistent between 256

the different look angles and satellite sensors. First, we determine the active sources during the 257

two seismically distinctive stages (pseudo-stationary in July-August and southeast migration in 258

September and early October). We use those interferograms spanning the shortest time in order 259

to constrain the best-fitting static displacements using simple elastic dislocation models. We 260

perform a non-linear inversion of the interferograms spanning the period May 2011 to October 261

2011. Once the static model is determined, a quasi-dynamic time-dependent model using a 262

Green’s function linear inversion scheme was employed to infer the variation of the source 263

volume at each interferogram, which subsequently were used to solve for the sources evolution. 264

265

4.1. Static model: Parameter estimation 266

We searched for the optimal model parameters of a few elastic models with simple geometries: 267

spherical point source [Mogi, 1958], horizontal circular crack [Sun, 1969], ellipsoid source [Yang 268

et al., 1988], and rectangular crack [Okada, 1985]. We included the effect of the topography by 269

using the approximate method of Williams and Wadge, [1998]. We adopted a non-linear global 270

inversion technique, bounded simulated annealing, to calculate the best-fitting model that finds 271

the minima of the misfit function in a least squares sense [González et al., 2010]. All the 272


interferograms were downsampled using a quadtree scheme [Jonsson et al., 2002] to reduce the 273

computational burden while preserving the observed deformation patterns. To avoid solutions 274

trapped in local minima, we restarted the non-linear inversion 250 times with a homogeneously 275

distributed random set of initial model parameters, with the same bounding limits as the model 276

parameters. We improved the method presented in Gonzalez et al. [2010] by introducing realistic 277

noise at each iteration [Wright et al., 2003]. We simulated the random noise using an exponential 278

and cosine theoretical covariance function, 279

C (r) = 2

[ exp (-a r)

cos (b r) ], (1) 280

where r is the distance between data points, 2 is the noise variance (we assume

2=4 cm

2) and, a 281

and b are coefficients which control the correlation lag (here a=b=0.05). Due to the small area of 282

the studied region, we could not estimate correlation parameters for the interferograms which 283

show deformation. Therefore, we assumed values that produce spatially correlated signals 284

similar to the observed interferograms before the onset of seismicity. Finally, we determined the 285

best-fitting model parameters as the mean of the 250 best-fitting set of modelled parameters and 286

the associated errors as their standard deviations. 287

288

In accordance with the seismicity evolution (Figure 2 and 3), we selected two interferograms 289

spanning short times over the intense periods of seismicity and deformation. The selection of the 290

specific interferograms was based on the possible sensitivity to different and independent sources 291

corresponding to the observed different seismicity and deformation patterns (Fig. 3 and 4). In 292

particular, we chose one associated with the pseudo-stationary period during July and August 293

(Figure 5a, covering the period 4 May 2011 to 8 August 2011), and another from the period 294

during which seismicity migrated south-to-southeastwards (Figure 5d, for the period 31 August 295


2011 to 30 October 2011). Solutions using the rectangular crack and ellipsoidal models were 296

poorly constrained displaying a very large dispersion of the set (250) of optimal best-fitting 297

parameters. We discarded those models because the model parameters had large uncertainties, on 298

the same order of magnitude as the prescribed search bounding limits, and a clear indication of 299

poor resolution. A circular horizontal crack and the spherical (Mogi) models provide essentially 300

the same results (Figure S2). However, in the case of the circular crack the source was generally 301

slightly deeper than the spherical source (Table 1). Although the location is well-determined, the 302

volume change was poorly resolved. Therefore, we preferred the spherical source results (Table 303

1) because it is the simplest model that explains the signal up to the noise levels. 304

305

In Figure 5, we show observed (5a and 5d), modeled (5b and 5e) and residual (5c and 5f) LOS 306

displacement for the best-fitting spherical sources for the two interferograms spanning the 307

selected periods. It is noteworthy that even with a very simple model, the spherical source, the 308

data fit is acceptable for the ascending 4 May 2011-8 August 2011 interferogram (Figure 5c), 309

despite some small residuals near the coastlines. In addition, we confirm the spherical best-fitting 310

parameter results for the first period by repeating the inversion process using a stack of four 311

observed RADARSAT-2 ascending differential interferograms with the 8 August 2011 as the 312

common slave image (Table 1; and Figure S2). For the August-early October period, a spherical 313

model is also a good approximation, as indicated by the fit to the descending 31 August 2011-30 314

October 2011 interferogram. However, small magnitude residuals along the Mar de las Calmas 315

coastline (Figure 5f) may reflect a slightly asymmetric causative source, with an axis along the 316

south rift direction (NNW-SSE). Considering these results, we subsequently assume the 317

existence of two spherical magma reservoirs consistent with elastic model point sources which 318


change in volume over time. The results constrain a static image of the magmatic system 319

composed of two spherical point sources: a deeper reservoir at ~9.5 km and a shallower at ~4.5 320

km depth (Figure 5g). 321

322

4.2. Tracking transient source volume changes 323

4.2.1. Volume changes over temporal intervals 324

Once a magmatic system of two spherical point sources was fixed, we performed a linear 325

inversion of all 63 available interferograms in order to estimate for volume changes of the 326

sources over the corresponding time interval. Ascending and descending interferograms from 327

COSMO-SkyMed data acquired on the same day (Table S1 and S2 in Supplementary Material) 328

were jointly inverted. For each interferogram we independently solved a forward problem of the 329

type: d = G m, where d is the data vector (InSAR LOS observations), m is the model parameters 330

vector (volume changes), and G is the Green’s function matrix (source impulse response 331

projected into the satellite LOS). The problem is solved using a bounded least squares method. 332

We allowed for a variation of the estimated volume change between 0.10 and -0.05 km3, which 333

are reasonable bounds considering the results of the non-linear inversion presented in the 334

previous section. As a result, we resolved for the source volume changes on both sources. 335

4.2.2. Source volume changes: temporal inversion 336

The problem of retrieving time series of the source volume changes from non-simultaneous and 337

temporally overlapped multi-sensor observations is an ill-posed problem. This problem can be 338

solved using regularization methods, e.g., first-order Tikhonov regularization [Grandin et al., 339

2010]. We consider a system of n increments of source volume changes (qij) between two 340

epochs ti and tj, where ti < tj, based on the linear inversion estimation of the corresponding 341


interferograms. We want to solve for the temporal evolution of volume change for each observed 342

epoch t [t0,..., tm], where m is the number of SAR acquisitions. In matrix notation, the problem 343

can be formalized as the inversion of the linear system q= A Q, where A is the differential 344

operator (i.e., the incidence matrix, with -1 and 1 for master and slaves dates respectively), q is 345

the known vector of increments of volume change (1 x n), and Q is the unknown vector of 346

temporal volume change (1 x m). A trivial solution to this problem, even in the least square sense 347

(generalized inverse), is not possible because the interferograms are formed by data from 348

different tracks and satellites and, therefore, not all acquisitions can be connected to each other. 349

As a result, the linear system is formed by independent linear subsystems, making the problem 350

underdetermined. 351

352

One way to solve this problem is to search for the least squares solution imposing an additional 353

minimum-norm condition (minimum volume); this can be accomplished through the use of the 354

Singular Value Decomposition (SVD) method. However, direct application of this strategy 355

provides unrealistic results with large discontinuities in the time series. 356

357

Similar to the approach of the Small BAseline Subset (SBAS) technique, which is designed to 358

recover the time series of displacements from a subset of disconnected interferograms 359

[Berardino et al., 2002], we reformulate the problem by minimizing the rate of the source 360

volume changes, instead of the volume variation itself. This avoids the presence of large 361

discontinuities in the final solution, which is retrieved after a final integration step, because the 362

minimum norm constraint implies minimizing the overall rate of volume change rates instead of 363

the volume change itself, which is a more physically sound constraint. Furthermore, to prevent 364


possible oscillation in the solution, we consider a Truncated Singular Value Decomposition 365

(TSVD) approach [Aster et al., 2005]. 366

367

According to the outlined strategy, we transformed the differential increments of volume change 368

into rates considering the time spanned by each interferogram, q = q / tij. We solved for the 369

vector of rates of volume change between each available epoch, Q, a vector 1 x (m-1), where the 370

underline denotes rates. The new system of equations is organized as follows, 371

q = B * Q, (2) 372

where B is a matrix n x (m-1) with elements corresponding to the spanned time (tk), for each 373

epoch between master and slave images (k ∈ i ≤ m < j), and zeros elsewhere. 374

375

To solve this problem, we decompose B using the SVD methods [Aster et al., 2005], 376

B = USVT, (3) 377

where U is an orthogonal matrix with columns that are the basis vectors of the data space, n x n, 378

V is an orthogonal matrix with columns that are the basis vectors spanning the model space, (m-379

1) x (m-1), and S is a diagonal matrix of the singular values (n x m-1). A solution for this 380

problem can be obtained as follows, 381

Q = VS-1

UT q . (4) 382

383

If rank(B) < m, the solution obtained using the SVD technique may contain numerical 384

instabilities when small singular values exist. In this case, a more stable solution can be obtained 385

using the TSVD method, which rejects model space basis vectors associated with small singular 386

values, up to a certain threshold p. Then an optimal solution can be obtained as follows, 387


Qp

i = Σp

i(U1xiT q / si ) V1xi , (5) 388

where i = 1,…, p, the solution using the p largest singular values (s). 389

390

Finally, the vector of rates of volume change must be integrated in time to obtain the time series 391

of volume change, Q(t), 392

Q(t) = Σ tk * qk , (6) 393

where tk is the time elapsed between each temporally contiguous SAR acquisitions. This 394

approach is able to handle the associated problem's rank deficiency, and effectively reduce 395

unstable results due to data noise as well. 396

397

In Figure 6 we show temporal evolution of the volume change for the considered sources (two 398

spherical sources), obtained as described above, for the pre-eruptive and eruptive periods (the 399

solution for the entire period, December 2009 – March 2012 is shown in Figure S4 of the 400

Supplementary Material). The inversion result subsequently was used to compute synthetic 401

interferograms via forward modeling based on the retrieved source parameters as shown in 402

Figure 7. We compare a set of the original observed interferograms with those simulated on the 403

basis of the time evolution of source changes (Figure 6). Those areas with significant residuals 404

may reflect atmospheric noise in the corresponding interferograms, activity on other structures 405

(e.g., shallow faults), a possible unaccounted eruptive conduit (e.g., see Figure 7c or 7d) or 406

departures from the assumed source geometries that were not included in our simplified model. 407

408

409

5. Discussion 410


5.1. Active shallow magmatic system during the 2011-2012 El Hierro eruption 411

A simple model (Figure 5g) with two active sources accounts for the vigorous motions towards 412

the satellite (uplift) observed before the onset of the submarine eruption (Figure 4). Two major 413

difficulties arise from the use of such a simple elastic model: the elasticity assumption that is 414

used to interpret all the observed ground deformation InSAR data, and the limited extension of 415

the imaged area. Due to uncertainties in the rheological structure beneath El Hierro Island, we 416

limit our discussion to crustal depths, where the rocks elastic behavior is commonly assumed. In 417

the case of the Canary Islands, Moho depths have been estimated to be approximately 10-18 km 418

[Banda et al., 1981; Ranero et al., 1995, Lodge et al., 2012]. The latter limitation reduces the 419

resolvability to sources as a function of depth. In particular, sources deeper than the Island’s 420

width will be difficult to detect because InSAR results are sensitive only to horizontal gradients 421

of the projected LOS motion, which are stronger for shallow sources than deeper sources. 422

423

InSAR results confirm that there was no slow precursory deformation signal before May 2011. 424

Therefore, most, if not all, of the deformation should have started with the initial onset of the 425

seismicity on 17 July. Between 4 May and 8 August 2011 a batch of magma with an estimated 426

volume of 0.028 ± 0.003 km3 intruded at the base of the oceanic crust (crustal-mantle reservoir, 427

CMR), at a depth of 9.5 ± 4 km. This is a depth consistent with the upper bounds of the 428

hypocentral depths (8-15 km) of the reported seismicity for this period. This depth is also 429

consistent with an intrusion stalled at the base of the oceanic crust, and likely not below it. The 430

oceanic crust-mantle (moho) boundary at the Canary Islands indicates a gentle deepening 431

towards the western Islands [Banda et al., 1981]. In particular, around El Hierro, it has been 432

estimated at about 12-15 km using gravity and seismic reflection data [Figure 7, in Ranero et al., 433


1995]. This batch of magma was located south of the main seismic locus (Figure 5g), and 434

directly beneath the central part of the Island, not to the north [López et al., 2012; Martí et al., 435

2013]. This southern location is consistent with a nearly ~N45E azimuth GPS velocity vector of 436

the public FRON GPS station. GPS data are inconsistent with a more northern location, as 437

suggested by the located seismicity (Figure 3a-b, and density curves in Figure 5g). Magma 438

located at the main locus of seismicity would have caused larger eastwards, rather than 439

northward motions. However, this behavior is not observed in the published GPS data. 440

441

At this initial phase, the location of the magma pulse seems to correspond quite well with the 442

inferred feeding system of the Tanganasoga volcano (4000 yr B.P.). This volcano is the only 443

polygenetic volcano in El Hierro Island, and the only area where it is common to produce 444

crystal-rich (ankaramite) lavas [Carracedo et al., 2001; Manconi et al., 2009]. Those crystal-rich 445

lavas are significantly denser than other El Hierro lavas, which indicate more primitive (deeper) 446

extraction regions, and interpreted as the effect of upper mantle decompression due to the El 447

Golfo flank collapse [Manconi et al., 2009]. This observation, coupled with the presence of a 448

new batch of magma at the CMR in this location, would indicate that this central area may be the 449

main feeding system for the entire El Hierro volcanism. Magmas would ascend further, erupting 450

directly above or distributed into the three rift zones, depending on the physic-chemical 451

conditions of the melt or its overpressure (gas content). 452

453

Between 8 August and 31 August 2011, our modeling results indicate the emplacement of an 454

additional magma volume of about 0.015 km3 into the CMR. The influx volume magma rate 455

apparently doubled (Figure 6). Although the increase in rate is not well constrained due to the 456


low temporal SAR data resolution during this period, the apparent increase in the magma volume 457

change time series is in accordance with an observed increase in low-magnitude seismicity rates, 458

which by middle August peaked >400 events/day (Figure 2c). That increase in seismicity rate 459

was not associated with a significant increment in the magnitude of the seismic events. However, 460

the estimated magma influx and observed daily seismicity rates were the largest for the entire 461

analyzed period. Simultaneously, between 10 and 17 August a station located in the central part 462

of the Island detected a sharp increase in the diffuse degassing efflux of CO2 [Pérez et al., 2012]. 463

Basaltic magma rising through a plumbing system saturates in CO2 at pressures corresponding to 464

about 30 km depth, then buoyant CO2-rich gas exsolves and can rise rapidly to the surface 465

[Gerlach et al., 2002]. It is worth noting that CO2 levels were at low levels in late July and early 466

August. This might be explained by the fact that CO2 is highly soluble in the waters of shallow 467

aquifers. Therefore, it is possible that the sudden increase marked the full saturation in CO2 due 468

to an increase in deep influx. After this, the levels of CO2 efflux and 222

Rn activity rates steadily 469

increased during September until the eruption in early October [Pérez et al., 2012; Padilla et al., 470

2013]. 471

472

In September, the CMR did not change significantly in magma volume. This correlates with the 473

gradual decrease in daily seismicity rates occurred until middle September. Since middle 474

September 2011, GPS stations located in the north of the Island started to indicate a more north-475

south deformation, with a slightly increase in daily deformation rates [López et al., 2012]. 476

Seismicity began to spread southwards, and hypocenters were located between 12 and 17 km 477

depth (deeper than in the previous period). The N-S clustering of the seismicity may suggest the 478

propagation of dikes, however the surface deformation does not confirm it. In fact, the deepening 479


of the events could be an artifact, reflecting that the southern part of the Island is significantly 480

more heterogeneous than the assumed 1D crustal-mantle seismic layered model [López et al., 481

2012], as indicated by the large horizontal and depth reported errors in the seismic catalog 482

(http://www.ign.es/ign/layoutIn/sismoFormularioCatalogo.do, note mean horizontal xy error 483

ellipse in Figure 2a, and mean depth error bar in Figure 2b). Therefore, hypocenter location 484

information should be interpreted carefully, at least for those events that occurred off-shore, 485

where the network configuration provides a poorer resolution. Nevertheless, these observations 486

suggest that magma migrated laterally away from the central CMR for a distance of about 10 km. 487

A possible reason for the migration could be the increase of magma overpressure due to 488

expansion/exsolution of CO2, as it might indicated by the increments in diffuse gases detected in 489

the central part of the Island [Pérez et al., 2012]. 490

491

During September the lack of dense temporal InSAR observations limits our ability to track 492

accurately the magma migration, particularly with depth. Modeling of the interferograms 493

spanning September 2011 did not support the existence of near vertical dikes. For example, the 494

descending interferogram 31 August to 30 October 2011 (Figure 4b) could be modeled by a dike 495

with a strike angle parallel to the south-southeast rift zone, but the location was either too far 496

from the coast or deepening strongly inwards towards the shore (not vertical). In addition, a 497

vertical dike poorly explains the observed descending interferograms. In Fig. 4d a small LOS 498

shortening signal south of the Island might corresponds to the opening of the submarine fissure 499

eruption feeding the dike between 30 September and 12 October. However, the data do not allow 500

for a reliable estimate of the causative source. As described previously, spherical or sill geometry 501

better explains the InSAR observations (Table 1, Figure 5d-f). In particular, interferograms 502

http://www.ign.es/ign/layoutIn/sismoFormularioCatalogo.do


modeling supports a shallower and smaller magmatic reservoir located slightly off-shore at the 503

Mar de las Calmas coastline and at 4.5 ± 2 km depth. The depth of this reservoir, hereafter SFR, 504

corresponds to the interface between the top of the oceanic crust, pre-volcanic sediments, and the 505

volcanic edifice. The thickness of the pre-volcanic oceanic sediments has been estimated to be at 506

least 1 km thick [Ranero et al., 1995]. This interface would generate a sufficient 507

density/rheology contrast to trap the ascending magmas. Geodetic modeling results points to an 508

increase in magma volume of about 0.01 km3 (Figure 6 and 7b). The SFR intruded magma 509

volume emplaced at some time between 30 August and 31 September 2011. 510

511

Cumulate geometric moment exhibited a steady-state increase, primarily due to events occurring 512

in the crust (Figure 2b and blue curve in Figure 6). During 23-27 September, the cumulative 513

geometric moment exceeded ~5 m3 (Figure 2b), a value which seems to be a critical threshold 514

value during this crisis (gray and blue curves in Figure 6). After 23-27 September (the last week 515

of September and early October), seismicity began to be located around the Mar de las Calmas 516

(Figure 3c and 3d). Cumulative geometric moment increased dramatically together with the 517

event magnitudes, with up to 90 being felt by the population [Martí et al., 2013]. We suggest that 518

this could mark the moment when the magma started to migrate and accumulate at the shallower 519

SFR. Furthermore, this upward migration provides a reasonable explanation for a positive 520

anomaly in differential geomagnetic signals that lasted for 3 days at this period. This was the 521

only anomaly registered between September and end of October 2011 [Figure 4 in López et al., 522

2012]. In addition, a mechanical loading model proposed to explain the magma migration 523

predicts that far from the load (the center of the Island) magma should emplace at shallower 524


depths [Martí et al., 2013]. In fact, the numerical model parameterization indicates that magma 525

should be trapped at about 3-5 km or deeper, at distances of 10 to 20 km from the load axis. 526

527

Between 30 September and 12 October, seismicity remained high, but we did not detect 528

significant changes in magma volumes at CMR or SFR. The largest event in that period occurred 529

on 8 October at 20:35 UTC. The relatively large magnitude of this event allowed for calculating 530

its focal mechanism parameters. Moment tensor waveforms inversion initially indicated 6 km 531

depth [www.ign.es accessed the 28 March 2012]. This event has been interpreted as the start of 532

the upward migration of the magma towards the eruptive vents [Carracedo et al., 2012; López et 533

al., 2012; Martí et al., 2013]. The shallow moment centroid depth (~6 km) seems to be 534

consistent with our estimate for the SFR and realistically could have opened a path for the 535

migration of magma towards the surface. The surface rupture involved in a Mw4.0 corresponds 536

to an area of approximately ~500x500 m2 with a few centimeters of slip, probably too small to 537

open a significant path at ~12 km depth for magma upward migration. In addition, a shallower 538

reservoir could explain the observed high levels and gradual increase in H2S efflux recorded at a 539

station located on the Mar de las Calmas coastline prior to the submarine eruption [Pérez et al., 540

2012]. On 7 October (installation date), this station recorded H2S efflux values about twice as 541

high as the levels after the eruption began, when H2S efflux stabilized to assumed background 542

levels. Hydrogen sulfide separates only from the parental magmas at hydrothermal conditions 543

(low lithostatic pressures). Conversely, Pérez et al., [2012] reported absence of anomalous H2S 544

efflux at the center of the island. It supports the conclusion that magma at a depth of 10-15 km 545

could not exsolve that sulfur species in measurable quantities at the surface. Therefore, this 546

http://www.ign.es/


indicates that the magma was already at shallow depths on the 7 October, at least a day before the 547

8 October earthquake. 548

549

We favor the hypothesis that magma started to migrate upwards and accumulate at the SFR on 550

26-27 September and for the next two weeks the reservoir pressurized. The pressurization 551

culminated with the 8 October 2011 event. Then, for approximately 32 hours the magma 552

migrated laterally in the topographically down-rift direction for ~5 km. During this period, very 553

mild and shallow (one to six km depth) seismicity occurred [www.ign.es]. This period ended 554

with the emergence of a growing amplitude tremor signal (10 October 2012, ~4:15 UTC). 555

Tremor amplitude increased for about 58 hr when the signal abruptly ceased after a peak in high 556

frequencies at 14:30 UTC on 12 October 2011. During the next 24 hours several additional high 557

frequency peaks were recorded. We interpret this as the magma started to ascend vertically on 558

the 10 October 2012, ~4:15 UTC. Tremor increased in amplitude as the magma was ascending 559

and decompression of vesicularized magmatic gasses or interaction with groundwater was 560

amplified, as commonly observed in other volcanoes [McNutt et al., 1996; Patanè et al., 2008]. 561

The sudden decrease in tremor amplitude provides a good indication of the initial/explosive 562

opening of the submarine vent. Therefore, the start of the submarine eruption occurred on 12 563

October 2012, when the first clear pale-color water plume was observed. Subsequent, high 564

frequency peaks in the tremor could be associated with submarine explosions. The temporal 565

evolution suggests a final magma ascend rate of ~0.02-0.03 m/s. This much simpler 566

interpretation accounts for the observations and contrasts with a previously proposed migration 567

model based on a 3 km long rift back-propagation towards higher topography, from 1000 m to 568

300 b.s.l. depth [Martí et al., 2013], which, to date, is not supported by the published bathymetric 569


surveys [http://www.ieo.es/hierro.htm]. It also explains the absence or minor effects observed on 570

the sea surface between the 10 and 12 October 2011 [López et al., 2012]. Some residuals in the 571

modeled interferograms, limited to the very south edge of the island, could be related to the 572

shallow conduit feeding the eruptive fissure (Figure 7c and 7d). 573

574

On 15 October 2011, floating material appeared off La Restinga. The material varied from a 575

mixture of mafic and white pumice to large basanite floating balloons [Troll et al., 2012; 576

Carracedo et al., 2012]. Pérez-Torrado et al., [2012] attributed the latter to material commonly 577

formed during shallow submarine eruptions, as observed during the La Serreta eruption in Faial, 578

Azores Islands [Gaspar et al., 2003]. Basanite lava ballons were observed during the entire 579

eruptive period. However, the first products called “restingolitas” or xenopumices, appeared just 580

during the first few days of the eruption. The xenopumices have been attributed as formed by 581

three main processes: a) partial melting and vesiculation of pre-island quarz-rich sedimentary 582

rocks carried by the ascending magma [Troll et al., 2012] ; b) remobilization of a stagnant 583

phonolitic melt present as late differentiate in the crust and successive variable interaction with 584

old oceanic crust and the volcanic edifice [Sigmarsson et al., 2012]; and c) the result of the 585

thermally-induced interaction between an ascending basanitic magma, a stagnant trachytic 586

magma pocket/s and an associated hydrothermally altered halo with rhyolitic composition 587

[Meletlidis et al., 2012]. Indeed, the xenopumices interpretations could provide some clues on 588

the migration of the magma. Petrological experiments (mineral assemblage, composition and 589

crystal content) carried out with xenopumice samples are compatible with an isobaric (P=100 590

MPa, ~ 4 km) crystallization of trachytic melt cooling from 1035 to 900 ºC [Meletlidis et al., 591

2012]. However, other analytical results of xenopumice incompatible trace elements, designed to 592


distinguish different magmatic differentiation processes are still controversial [Meletlidis et al., 593

2012; Troll et al., 2012]. It is interesting to note that all hypotheses for the generation of the 594

xenopumices require magma stalled with different degrees and/or time residences, at shallow (~4 595

km) depth [Meletlidis et al., 2012; Troll et al., 2012; Sigmarsson et al., 2012], in agreement with 596

a pre-island seafloor and volcanics interface. These observations further support a previously 597

undetected shallow SFR reservoir. 598

599

After the start of the eruption, the volume change at the deeper CMR started to decline, while the 600

shallow reservoir, SFR, did not change significantly for the remainder of the study period, as 601

confirmed by the absence of shallow depth cumulative geometric moment (blue curve in Figure 602

6). This decrease lasted until beginning of November 2011, when it ceased. The increment in 603

magma volume seems to correlate with a period of seismic quiescence and the beginning (26 604

October 2011) of very deep (20-30 km) seismicity located in the north of the Island (Figure 3d-605

e). During this phase, earthquakes magnitude increased until to the occurrence of a mbLg 4.6 606

event on 8 November 2011. Just after the mbLg 4.6 earthquake, an increment in volume of 607

approximately 0.01 km3 occurred at the CMR increase. The earthquake sequence could be the 608

result of a collapse of the deeper segments of the magma plumbing system due to magma 609

withdrawal and/or chamber walls relaxation [McNutt et al., 1996]. A geodetically undetected 610

deeper reservoir or plexus of intertwined intrusions and sills at this depth range corresponds with 611

petrologically constrained results [Stroncik et al., 2009]. The collapse of an ~25 km deep magma 612

reservoir could favor a transient increase in magma ascent rate. Evidence for a new, less-613

degassed (more primitive) magma batch is supported by the significant increment of CO2 soil 614

diffuse efflux surveys in November-December 2011 [Melián et al., 2012]. 615


616

In late December, visible sea surface discoloration disappeared. This might be related to the 617

observed steady volume change decrease at the CMR in December (Figure 6). That clear trend 618

stopped in early January when basaltic floating bombs re-appeared and occurred discontinuously 619

throughout February. Again, this observation agrees with a mild and not completely significant 620

increase in volume change at the CMR reservoir. This also correlates with a modest increase in 621

the detected daily seismicity rate (Figure 2c). During February and March, CMR volume change 622

further decreased. By early March, the absence of visible effects on the sea surface and the lower 623

levels of seismicity triggered the official declaration of the end of the eruption on 5 March 2012 624

[www.gobiernodecanarias.org/dgse/noticias_sismo_hierro.html]. The invariant (overpressure) 625

volume change at the shallower rift flank reservoir, SFR, the behavior of the volume change in 626

the CMR, and its relation to the visible changes at the eruptive site is evidence for an active 627

(magmatic) hydraulic connection between the deep plumbing system and the eruptive vent. We 628

illustrate a conceptual model of the entire eruption in Figure 8. 629

630

These results, derived from the geodetic data of the 2011-2012 eruption, support a mechanical 631

scenario beneath the El Hierro Island which is quite similar to that proposed for the Island of 632

Hawaii [Ryan, 1993]. A central magmatic plumbing system focus magma (dikes) ascent, which 633

propagates from the mantle upwards due to buoyancy forces until it reaches a horizon of neutral 634

buoyancy [Tilling and Dvorak, 1993], the oceanic crust. Then, depending of the relative magma 635

overpressure (e.g., gas content) new and/or evolved magmas might be distributed into the rift 636

zones or continued along a vertical ascent path. Therefore, at least two horizons of neutral 637

buoyancy would be present at El Hierro. A deep horizon associated with olivine-rich cumulates 638

http://www.gobiernodecanarias.org/dgse/noticias_sismo_hierro.html


(main magma underplating region) exists at a depth of 6-10 km for Hawaii, or ~9.5 km at El 639

Hierro (lower oceanic crust) and a shallower one at 2-4 km in Hawaii, or ~4.5 km at El Hierro. 640

Magma accumulation occurs preferentially at or near those horizons, which facilitates lateral 641

migration [Fialko and Rubin, 1998; Grandin et al., 2012] and possible eruption along rift zones. 642

643

5.2. Implications for the active magmatic systems of oceanic island volcanoes 644

We obtain a detailed and complex image of the main magma storage locations. We track the 645

magma migration before, during and after the El Hierro 2011-2012 submarine eruption by 646

combining data from different SAR sensors. Geodetic data and modeling results, coupled with 647

independent observations (e.g., gas emissions, loading mechanical models, and geochemistry and 648

petrology of erupted material) from this eruption revealed its crustal magmatic plumbing system. 649

Our results for the El Hierro eruption indicate two levels of magma stagnation that persist for a 650

minimum of weeks to months at the base of the oceanic crust (~9.5 km), and at least days to 651

weeks at the paleo-oceanic seafloor (~4 km). 652

653

Evidence presented here is in contrast with previous magma plumbing systems models (depths) 654

based solely on petrological analysis of Canary Islands lavas. Petrology evidence shows very 655

short-term magma stagnations of, at most, few days at shallow crustal levels [Kluegel et al., 656

2005]. Most studies systematically indicate magmatic systems with main storage volumes at 657

mantle pressures of, 600-1400 MPa or ~15-45 km depth [Kluegel et al., 2005; Longpré et al., 658

2008; Stroncik et al., 2009]. However, Kluegel et al., [2005] also identified a bi-modal 659

distribution of pressures based on the analysis of phenocrystals and fluid inclusions at La Palma, 660

with fluid inclusions indicative of shallower depths (7-14 km). They concluded that fluid 661


inclusions re-equilibrate faster than the growth of phenocrystal rims at equilibrium conditions. 662

The same bimodality has been observed at Fogo Island [Hildner et al., 2012]. Based on 663

considerations of typical re-equilibration times of phenocrystals and fluid inclusions, it excluded 664

shallower magma stagnation before eruptions with residence times longer than hours to days 665

[Kluegel et al., 2005; Longpré et al., 2008]. Therefore, our deeper inferred depth range (~9.5 666

km) for the active magmatic reservoirs during the El Hierro eruptions is consistent with depths 667

estimated from fluid inclusions analysis for oceanic island volcanoes [Kluegel et al., 2005; 668

Hildner et al., 2012]. However, our results indicate residence times at this depth range that are 669

significantly longer than the previously suggested hours to days. 670

671

Regarding the shallowest level of magma stagnation at El Hierro (~4.5 km), at least two of the 672

proposed hypotheses for the origin of the xenopumices imply the existence of shallow reservoirs 673

with trachitic magmas evolving to phonolitic compositions [Sigmardsson et al., 2012], or long 674

lasting rhyolitic alteration halo surrounding a hot shallow trachitic magma pocket [Meletlidis et 675

al., 2012]. Additionally, evidence for the relevance of this depth range for the Canary Islands can 676

be found from the analysis of other volcanoes. In Gran Canaria, estimated pressure in fluid 677

inclusion samples indicate reservoirs at 2-4 km depth, which are interpreted as a relevant horizon 678

of neutral buoyancy [Hansteen et al., 1998]. The analysis of fluid inclusion pressures of Gran 679

Canaria samples is relevant because those are from eruptive products in its shield basalt stage, 680

the current oceanic island volcano evolution stage for El Hierro. A similar magma depth has 681

been estimated for Tenerife Island (Canary Islands) during a more evolved stage (central sub-682

plinian volcanism). Recently, a magma storage depth was inferred at about 5 ± 1 km beneath the 683

summit of Teide volcano using phase equilibrium experiments [Andújar et al., 2010]. In 684


addition, P-wave seismic velocity contrast horizons at ~4 km depth have been found for La 685

Palma, Tenerife and Lanzarote [Lodge et al., 2012]. This suggests that this depth range may 686

represent a regional horizon of neutral buoyancy for the Canary Islands magmas. It clearly 687

corresponds to the average pre-island seafloor with a variable effect due to lithospheric flexural 688

volcanic loading. Such magma stagnation levels can be interpreted as an initial stage in the 689

formation of magma chambers, similar to those actually present in other, more evolved, Canary 690

Islands oceanic island volcanoes such as at Tenerife and Gran Canaria [Ablay and Martí, 2000]. 691

When magma chambers establish, primitive melts can fractionate to generate evolved magmas 692

(e.g., phonolites) that could reach eruptable conditions similar to the central Canary Islands. 693

However at least for the current stage of volcanic evolution, we find that those conditions are not 694

met in the El Hierro magmatic plumbing system. 695

696

At the current shield-building stage in El Hierro, we also cannot rule out that the detected 697

reservoirs may be more ephemeral and solidify completely at depth, thus confirming part of the 698

proposed active magmatic underplating beneath Canary Islands volcanoes [Dañobeitia and 699

Canales, 2000; Lodge et al., 2012]. Thus, the magmatic underplating process could explain the 700

observed rates of island uplift, as seen e.g., at La Palma [Staudigel and Schmincke, 1984; 701

Hildenbrand et al., 2003] and Tenerife [Kroechert et al., 2008]. Since the end of the eruption, 702

several episodes of renewed seismicity occurred in June-July 2011, September 2011, and 703

December 2011-January 2012 (www.ign.es). Those seismic swarms, all deeper than 10 km, were 704

accompanied with clear uplift signals at the IGN and INVOLCAN GPS monitoring networks 705

[www.ign.es; Sagiya et al., 2012]. Each post-eruption seismic swarm distributed over areas that 706

were not affected by previous seismicity during the pre-eruption and co-eruption periods. In 707

http://www.ign.es/


general, the seismo-deformation swarm patterns indicate that magmatic underplating contributes 708

to a relatively homogeneous uplifting of the entire volcanic edifice (island). The observed 709

episode at El Hierro suggests that the magma plumbing system and underplating process could 710

have significant differences among Islands. For example, at La Palma a strong asymmetry affects 711

the uplifted formations (seamount), which are tilted towards the west [Staudigel and Schmincke, 712

1984], which may partially explain observed azimuth differences in P-to-S receiver functions 713

beneath this island [Li et al., 2003]. These results are unexpected and reveal substantial 714

differences in the magmatic plumbing systems among the Canary Islands chain. Here, we 715

speculate that differences in the distribution of magmatic underplating and the relative 716

importance of the shallowest horizon of neutral buoyancy may result in the development of 717

central phonolitic volcanic systems in some islands, e.g. Tenerife and Gran Canaria, and absence 718

in other islands, e.g., Lanzarote, Fuerteventura, La Gomera and perhaps La Palma. 719

720

Studies of the magma plumbing systems at (basaltic) oceanic island volcanoes have revealed 721

fundamental differences in their shallower (crustal) levels (e.g., Reunion, Galápagos, Hawaiian, 722

Canarian and Cape Verde Islands). Shallow (crustal) magma chambers have been postulated but 723

they are not always present or detected. At Hawaii, magma chambers stagnate and distribute 724

magma along elongated rift zones [Tilling and Dvorak, 1993]. At Tenerife and Gran Canaria 725

Islands, crustal magma chambers develop as long-lived features able to fractionate parental 726

basanitic composition melts into end-member (highly explosive) phonolitic magmas [Ablay and 727

Martí, 2000]. Our results indicate two active crustal magma reservoirs at ~9.5 and ~4.5 km for 728

the 2011-2012 El Hierro eruption. Interestingly, volcanic systems with two crustal magma 729

accumulation levels are relatively common for a number of oceanic island volcanoes. For 730


example, in the Island of Hawaii the central feeding system of Kilauea volcano is composed of 731

main two main levels of magma storage, in particular at about 1 km and about 3 km depth 732

[Cervelli and Miklius, 2003; Baker and Amelung, 2012]. Two active magmatic chambers have 733

been detected in the last 30 years at La Piton de la Fournaise, at ~2.5 and ~7 km [Peltier et al., 734

2009]. At the Galápagos Islands, active shallow crustal magma reservoirs have been detected 735

geodetically at Wolf, Darwin, Alcedo, Cerro Azul and Sierra Negra volcanoes. However, only at 736

Fernandina Island volcano has two levels of magma storage been reliably estimated at depths of 737

~1 and ~5 km [Chadwick et al., 2011; Bagnardi and Amelung, 2012]. As it seems a common 738

feature, we can designate this two-level magma crustal storage system in oceanic island 739

volcanoes as a doublet crustal magmatic system. 740

741

The existence of shallow crustal magmatic systems has been previously attributed to relative 742

importance of the deep mantle magma melt supply rates [Clague and Dixon, 2000]. Our results 743

indicate that the high melt supply is not a sufficient condition for the existence of the shallow 744

magma reservoirs, because the Canary Islands has one of the lowest magma supply rates for 745

oceanic island volcanoes, ~0.0002 km3/year [Amelung and Day, 2002]. Other conditions have 746

been proposed to explain the existence or not of crustal magmatic chambers. Hildner et al. 747

[2012] proposed that the age of the underlying lithosphere could control this behavior. In Figure 748

9, we show the relation between age of the underlying lithosphere and the depth of crustal 749

magma chambers (constrained using geodetic data, under relatively similar assumptions). The 750

figure illustrates a weak trend from shallower systems to deeper ones, as oceanic lithosphere 751

ages. However, an anomaly occurs in the case of the crustal magmatic system from Hawaiian 752

volcanoes, which are resting on a ~90 Ma old oceanic lithosphere. The active Hawaiian 753


magmatic systems are significantly shallower than the one expected for an age progression. If we 754

consider mid- to long-term magma supply rates, a clearer trend is observed (Figure 10). Both the 755

shallowest and deepest reservoirs in single and doublet systems are deeper when the magma 756

supply rate is lower. This result illustrates that the main controlling factor is the 757

thermomechanical structure of the crust, which depends strongly on the deep mantle magma 758

supply rate, and also on the age of the oceanic lithosphere. This occurs because older oceanic 759

lithospheres are significantly cooler than younger and buoyant oceanic crust (lithosphere), 760

causing the seafloor to deepen due to thermal contraction and increase in density. 761

762

The observation that there is a pervasive absence of reservoirs at the shallow crustal levels at 763

some oceanic islands needs to be revised [Amelung and Day, 2002; Kluegel et al, 2005]. The 764

lack of shallow magmatic systems may be attributed to fairly local processes that sometimes can 765

significantly disturb their structure, such as catastrophic flank collapses [Amelung and Day, 766

2002; Manconi et al., 2009]. On the other hand, it might a sampling artifact, because many 767

volcanoes lack sufficient or good quality data over the typical long-recurrence eruptive periods, 768

in particular for oceanic island volcanoes with modest magma supply rates (e.g., Azores, Canary 769

Islands, Cape Verde or Comoros). 770

771

772

6. Conclusions 773

We successfully combined multi-frequency and multi-satellite interferometric data to study the 774

volcanic deformation associated with the 2011-2012 El Hierro submarine eruption, with a 775

temporal sampling of ~10 days during the eruption. This study demonstrates the potential to 776


obtain high-resolution temporal deformation and mitigate poor geodetic volcano monitoring 777

systems using a combination of space-borne radar satellites. In addition, our results also 778

demonstrate the applicability of radar observations to studying off-shore eruptions, if occurring 779

close to the coast. In the near future, the next generation of radar satellites, such as the satellite 780

constellations COSMO-SkyMed and the planned RADARSAT Constellation Mission, and short 781

revisit time missions such as ESA Sentinel-1 and the Japanese ALOS-2, will allow the study of 782

sub-weekly to daily magma volume changes in land volcanic systems at a global scale. 783

784

This data processing and modeling approach allows us to analyze the full range of temporal 785

changes of the magmatic sources that produce the observed deformation (Figure 8). The 786

interpretation of the results provides evidence for two reservoirs, CMR, larger and ~9.5 km 787

depth, and SFR, a smaller off-shore rift-elongated at ~4.5 km depth. This dense temporal 788

sampling suggests magma started to accumulate in the centre of the Island for about two months, 789

migrated southwards, and ascended to about 4 km depth in late September. Finally, lateral and 790

vertical migration occurred between 9 and 12 October 2012. Starting on the 12 October 2012 and 791

continuing for the subsequent 5 months, magma erupted along a submarine NW-SE fissure 792

vent(s) without significant fissure migration. Volume changes in the CMR are consistent with the 793

ascent of one or two more primitive magma pulses from an upper mantle (deeper) source during 794

November 2011 and January-February 2012. The latter batches of magma contributed to the 795

relatively long eruption duration (~5 months), as compared to previous shorter Canary Islands 796

eruptions. 797

798


The detection/existence of a very shallow magmatic reservoir has important implications in the 799

understanding of the Canary Islands volcanoes crustal magma plumbing systems. We conclude 800

that the existence of shallow reservoirs, active for relatively long periods of time (at least weeks 801

to months) might suggest initial stages of magma chamber formation. Shallow magmatic 802

reservoirs could fractionate primitive basanitic magmas into evolved trachites or phonolite, as 803

erupted in La Palma and more voluminously in Tenerife and Gran Canaria. Such magmas could 804

reach eruptable conditions as the magmatic system that feed the oceanic island volcano evolves 805

to a more mature evolutionary stage at least for the western Canary Islands. However, at El 806

Hierro, those conditions do not meet the current stage of volcanic evolution. In addition, magma 807

stagnated at or near the crustal-mantle reservoir(s) may contribute to the deep cumulitic 808

complexes. Cumulitic complexes are common at oceanic island volcanoes and have been 809

detected using geophysical surveys. The complete solidification of magma would correspond to 810

an active magmatic underplating beneath Canary Islands volcanoes. Spatial variability of 811

magmatic underplating could explain the observed rates of island uplift and patterns, as observed 812

at La Palma and Tenerife Islands. 813

814

Geodetic interpretation of the deformation associated with the 2011-2012 eruption revealed that 815

the crustal magma plumbing system of the Canary Islands oceanic island volcanoes are relatively 816

similar to other oceanic volcano groups (Hawaii, Piton de la Fournaise, Galápagos). In particular, 817

doublet magma storage levels seem to be a recurrent feature in many oceanic island volcanoes. 818

Lack of shallow crustal magma chambers (e.g. El Hierro) has been suggested for low-magma 819

supply rate volcanic areas. This could lead to artifacts resulting from the low activity and long-820

recurrence time of eruptions. However, local catastrophic events might sometimes disturb the 821


shallow magmatic system, e.g. Fogo Island. Data from El Hierro shallow magmatic system 822

suggest that the existence and depth of crustal magma chambers can be attributed to the 823

thermomechanical conditions of the oceanic lithosphere on which the volcanic islands rest. Such 824

thermomechanical conditions would depend strongly on the mantle magma supply rate, and on 825

the age of the oceanic lithosphere. 826


Acknowledgments 827

KFT and PJG are partially supported by an Ontario Early Researcher Award and an NSERC 828

Discovery Grant. PJG also acknowledges the Banting Postdoctoral Fellowship (Canadian 829

Government). SP has been supported by the Italian Space Agency (ASI). Research by JF and 830

AGC has been supported by the MICINN research project AYA2010-17448. It is a contribution 831

for the Moncloa Campus of International Excellence. RADARSAT-2 images were acquired 832

under SOAR-E project #28209 from the Canadian Space Agency. ENVISAT images from ESA 833

CAT1 project #6745. COSMO-SkyMed data were processed at IREA-CNR within the Italian 834

Space Agency (ASI) SAR4Volcanoes project (agreement I/034/11/0). The authors are grateful to 835

Mike Poland for insights gained during an early review of this paper, two anonymous reviewers 836

who improved the manuscript quality, and Mimmo Palano for his rapid response to the 837

computation of FRON GPS time series. We also thank Simona Zoffoli of ASI for her support in 838

planing and providing COSMO-SkyMed data. GMT Open source software was used for all 839

figures, except Figure 8 which was designed using Open source software Inkscape. 840

841

842

843

844


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Figure 1. Map of El Hierro Island and the three rift zones (Northeastern, Western-

Northwestern, and South-Southeastern) showing topography (shaded topography).

Simplified geological map of El Hierro. 1.2–0.88 Ma Tiñor volcano (green), 545–176 ka

El Golfo volcano (blue), and rift volcanism, pre- and post-El Golfo collapse in orangish

and red, respectively. Flank collapses are denoted with blue dashed lines. Populated areas

are shown with gray polygons. Inset: El Hierro (EH) location in the Canary Islands NW

off African continent. Red star indicates the approximate location of the eruption site

(same in all map figures).

Figure 2. (a) Seismicity located according to near-real time Spanish Instituto Geografico

Nacional seismic catalog (www.ign.es). Events are color-coded according to date. Gray

circles are lower and upper standard errors for events location reported in López et al.,

[2012] and black ellipse illustrates the mean standard error taking into account horizontal

correlation (http://www.ign.es/ign/layoutIn/sismoFormularioCatalogo.do). (b) Vertical

cross-section projecting all seismic events along longitude -18.02. Black bar shows the

standard errors in depth location. (c) Seismicity rate, earthquakes per day, and cumulative

seismic geometric moment, in cubic meter. Shaded area indicates the eruption period.

Top color marks indicate satellite SAR acquisition dates (blue Radarsat-2, red ENVISAT

and black CSK).

Figure 3. Monthly seismicity according to Spanish Instituto Geografico Nacional

(www.ign.es). (a) July 2011, (b) August 2011, (c) September 2011, (d) October 2011, (e)

November 2011, and (f) December 2011 and March 2012. In panels (a) to (e) events are

color coded according to date of month. For panel (f) events are color coded according to

date since 2011/12/01.

Figure 4. Observed interferograms spanning (a) period of without deformation

(2010/09/06-2011/05/04); (b) initial stages of the magmatic activity with seismicity

located in the center of the Island (2011/05/04-2011/08/08); (c) the period of southwards

migration of seismicity in September-early October (2011/08/31-2011/10/30); (d)

interferogram covering days before the onset of the submarine eruption (2011/09/30-

2011/10/12); (e) independent ascending Radarsat-2 interferogram spanning all the

deformation associated with the entire pre-eruptive phase (2009/12/09-2011/10/12); and

(f) an eruptive interferogram illustrative of the deep recharge of the system, a wide

deformation pattern in the central part of the Island (2011/10/30-2011/11/29). All

interferograms shown here indicate motion primarily towards satellite (negative is

shortening LOS).

Figure 5. (a) Observed ground-deformation (4 May 2011-8 August 2011) using ascending

Radarsat-2 images, with a maximum motion towards satellite of ~9 cm. (b) Simulated

ground deformation predicted by the best-fitting single spherical source model (see g) of

(a). (c) Residual of (b) and (a). (d) Observed ground deformation (31 August 2011-30

October 2011) using descending ENVISAT images from Track 109, with a maximum

motion towards the satellite of ~10-12 cm. (e) Simulated ground-deformation predicted

with the best-fitting single spherical source model (see g) of (d). (f) Residual of (b) and

(a). (g) Location of the best-fitting spherical point sources: orange, deep crustal source

(b); and dark red, the shallower crustal reservoir (e). Seismicity flux (events/km2), which

represents the 2D clustering of background seismicity, is shown the background. Inset

shows the vertical cross-section b-b'.

Figure 6. Time series of the volume change between May 2011 and March 2012 for the

two spherical point sources (orange circles, CMR deep source; and dark red diamonds,

the SFR shallow rift reservoir) from the TSVD inversion. Cumulative seismic geometric

moment, in cubic meter, is indicated as a shaded gray area for all events, and shaded blue

for earthquakes shallower than 15 km. In addition, we show in gray, the time series of the

NS component from FRON GPS station, applying a moderate temporal filter (boxcar

with width 0.05 years) the time series shows minor oscillations after September 2011,

which seems to be correlated with the behavior of the CMR reservoir (~9.5 km).

However, these oscillations are well below the time series scatter and will require a

deeper data analysis and additional GPS time series to extract robust conclusions.

Figure 7. Matrix of observed, simulated, residuals and profiles (model in red and data in

black) for various interferograms spanning the pre-eruptive phase (a) 2011/05/04-

2011/08/08; (b) 2011/08/31-2011/10/30; (c) 2011/09/30-2011/10/12; (d) 2009/12/09-

2011/10/12; and the eruptive phase (e) 2011/10/30-2011/11/29.

Figure 8. Cartoons showing North-South cross sections through El Hierro Island. Each

panel illustrates a schematic model of the active magma plumbing system at the different

phases discussed in the text.

Figure 9. Depth of inflating/deflating magma chambers at different oceanic island

volcanoes compared with the age of the underlying oceanic lithosphere. A weak trend

towards deeper magma chambers as the oceanic lithosphere aged can be observed in the

shallow (red circles) and deep (blue squares) crustal reservoirs. This trend is disrupted in

the case of the active volcanoes of the Hawaii Island (Kilauea and Mauna Loa). Data and

references in Table S3.

Figure 10. Same as Figure 9, but comparing depth of crustal reservoirs against the magma

supply rate for oceanic island volcanoes. A clear trend can be observed in the shallow

(red circles) and deep (blue squares) crustal reservoirs. Data and references in Table S3.


1

TABLE 1. Numerical results for inverted magmatic source parameters 2

3 Phase 1 (June-August 2011)

Interferogram Source Lon [deg. ± km] Lat [deg. ± km] Depth [km] Radius [km] Vol [km3] RMS [cm]

20110504-20110808 Mogi -18.0584 ± 3.7 27.7103 ± 4.3 8.4 ± 4.1 - 0.031 ± 0.010 0.93

Sill -18.0739 ± 2.2 27.7587 ± 5.4 9.7 ± 2.5 0.615 ± 0.546 0.012 ± 0.360 0.70

Stack R2S6 Mogi -18.0681 ± 1.8 27.7126 ± 4.2 10.2 ± 4.7 - 0.034 ± 0.015 0.66

Sill -18.0641 ± 1.2 27.7178 ± 2.2 11.2 ± 1.8 0.75 ± 0.396 0.022 ± 0.195 0.77

Phase 2 (August-October 2011)

Interferogram Source Lon [deg. ± km] Lat [deg. ± km] Depth [km] Radius [km] Vol [km3] RMS [cm]

20110831-20111030 Mogi -18.0370 ± 2.2 27.6653 ± 2.3 4.2 ± 2.4 - 0.012 ± 0.010 0.91

Sill -18.0471 ± 3.4 27.6701 ± 4.4 4.9 ± 3.8 3.604 ± 0.889 0.012 ± 0.809 1.81

4

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