Earthquake swarms in continental rifts — A comparison of selected cases in America, Africa and Europe

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Tectonophysics 452 (2008) 66 – 77 www.elsevier.com/locate/tecto

Earthquake swarms in continental rifts — A comparison of selected cases in America, Africa and Europe M. Ibs-von Seht a,⁎, T. Plenefisch b , K. Klinge b a

b

Federal Institute for Geosciences and Natural Resources (BGR), Germany Central Seismological Observatory, Federal Institute for Geosciences and Natural Resources (SZGRF/BGR), Germany Received 26 July 2007; received in revised form 1 February 2008; accepted 17 February 2008 Available online 26 February 2008

Abstract The occurrence of earthquake swarms is typically related to magmatic activity in volcanoes, yet swarms are also common in other intracontinental regions such as continental rifts. We present here a summary of geophysical observations that have been made in earthquake swarm areas of the Rio Grande, Kenya, and Eger rifts, focusing on characteristic parameters for the origin and generation of the swarm earthquakes. Our compilation of seismological parameters such as spatial distribution and focal parameters of hypocenters, magnitude statistics, and the location of the swarm centres in the rift environments reveals major similarities. The earthquake swarms take place at shallow depth between 0 and 10 km. The maximum magnitudes are mostly less than 4.5. The b-values, indicating the magnitude frequency relation of the seismicity, are about 0.8. They are hence not deviating from a normal non-volcanic intraplate environment, but are considerably lower than those of volcanic earthquake swarms. Focal mechanism studies give uniform pictures of stress field orientation and faulting style for the swarm areas. In all three rifts, the centres of swarm activity seem to be restricted to rift valley sections that may be influenced by large-scale fracture or shear zones that intersect the rifts. We conclude that these deep-reaching zones of weakness allow intrusions of upper mantle material into crustal layers, where magma-related fluids or fluctuations of the magma bodies themselves cause the generation of earthquake swarms. © 2008 Elsevier B.V. All rights reserved. Keywords: Earthquake swarms; Rifts; Seismicity

1. Introduction Earthquake swarms are seismic activity that is concentrated in time and space without an outstanding principal event. Thus, earthquake swarms differ from the common form of seismicity where a mainshock is followed by a number of aftershocks of significant lower magnitude. Mogi (1963) introduced a classification of earthquake sequences into three types, depending on their time–frequency relation. Types I and II comprise mainshock–aftershock and foreshock–mainshock– aftershock patterns, respectively, while swarms are assigned to type III which describes a gradual increase and decay of seismicity in time, without a distinct main shock. ⁎ Corresponding author. Federal Institute for Geosciences and Natural Resources (BGR), Stilleweg 2, 30655 Hannover, Germany. Tel.: +49 511 6432911. E-mail address: [email protected] (M. Ibs-von Seht). 0040-1951/$ - see front matter © 2008 Elsevier B.V. All rights reserved. doi:10.1016/j.tecto.2008.02.008

Earthquake swarms are most common in areas of active volcanism. A database compiled by Benoit and McNutt (1996) lists 136 volcanoes worldwide where swarms have been observed (Fig. 1). The occurrence of swarms in regions of active volcanism is commonly related to eruptions and magma movements (e.g. Karpin and Thurber, 1987). Volcanic swarms have an earthquake size distribution characterized by an unusual large b-value between 1 and 2.5. The observation of earthquake swarms can sometimes help predict forthcoming eruptions of active volcanoes. Mogi (1963) attempted to explain the nature of earthquake swarms on the basis of laboratory experiments and seismicity patterns in Japan. He proposed that earthquake swarms occur in regions where the crust is highly fractured and the stress system is heterogeneous. In such an environment, which is typical for volcanic regions, only low shear stresses are needed to produce seismogenic dislocations, resulting in a great number of small events while stronger earthquakes are sparse or absent. This explains the

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Fig. 1. Worldwide distribution of volcanoes (solid triangles) and swarm earthquake regions (open circles). Diamonds indicate swarm regions compared in this study.

high b-values that are commonly observed for swarm episodes connected to volcanic activity. Apart from active volcanoes, swarms also occur in intraplate environments. Also at these sites, swarms are mostly connected to recent volcanic, magmatic, or hydrothermal activity. Špičák (2000) gives an overview of intraplate earthquake swarm regions and highlights the general importance of fluids in the generation of earthquake swarms. In this paper, we concentrate on earthquake swarms in continental rifts because we consider them to play a key role in the understanding of the coupling between subsurface magmatic processes and seismogenic stress release. We provide an overview of what is known about earthquake swarms in continental rifts and discuss possible connections between seismological parameters and the geological and structural settings. In most of the prominent continental rifts worldwide, the Rio Grande rift, the European Cenozoic rift system, the East African rift system, and the Baikal rift, earthquake swarms do occur (e.g. Ibs-von Seht et al., 2001; Déverchère et al., 2001; Sanford et al., 2002; Klinge et al., 2003). Rifts are particularly suitable to study the occurrence of earthquake swarms because the crustal structure, tectonics and stress field parameters are often clearer compared to the situation for active volcanoes. We will focus on three areas where swarms have been studied in most detail. These areas are the Socorro area in the Rio Grande rift, the Lake Magadi area in the Kenya rift that is an element of the East African rift system, and the NW-Bohemia/Vogtland area in the Eger rift that is an element of the European Cenozoic rift system. A comprehensive compilation of geological and geophysical information on continental rifts can be found in Olsen (1995). 2. Rio Grande rift The Rio Grande rift is a major tectonic structure running north–south through Colorado and New Mexico (Fig. 2A). The rift basically consists of a series of asymmetrical basins that had formed in response to regional extensional forces and uplift.

Two main stages can be identified in the development of the Rio Grande rift (Morgan et al., 1986; Aldrich et al., 1986; Olsen et al., 1987): An initial stage (30 to 20 Mya) with low-angle faulting in response to NE–SW extensional stress formed broad shallow basins and significant regional extension, and a second stage (10 Mya) with high-angle normal faulting in response to E–W extensional stress formed graben and half-graben structures. The rifting was accompanied by extensive volcanic activity that was relatively modest in volume. A great part of the younger volcanism occurred along the Jemez lineament. There are several indications that rifting is still continuing at the present time. Geological evidence, such as fault scarps cutting Pleistocene surfaces and alluvial fans is supported by various geophysical observations, all indicating the presence of modern mid-crustal magma bodies in the rift centre and a mantle upwarp at the base of the crust (Bachmann and Mehnert, 1978; Wilson et al., 2005a,b). Based on statistical analysis of seismicity data and the observation of a clear lineation of topographic features, Sanford and Lin (1998) proposed the existence of a 1400 km long and ~85 km wide, ENE trending transfer zone, the “Socorro fracture zone”, that intersects the Rio Grande rift at about 34°N (Fig. 2A). Some similarities to another linear oblique structure crossing the rift further north, the “Jemez lineament”, were highlighted. Whereas the Jemez lineament is defined by an array of late Cenozoic volcanic fields (Baldridge et al., 1991), a mid-crustal magma chamber is assumed for the Socorro region (e.g. Sanford et al., 1973). It was further observed that the alignment of the rift axis, the width of the basins, and the faulting style change at these intersections. The geodynamic processes in the Rio Grande rift are accompanied by an increased level of seismicity. Most of the seismic activity is concentrated in several spots along the rift and occurs in the form of earthquake swarms. Most prominent for its anomalous seismic activity is a section in the central part of the rift near the town Socorro (Fig. 2B), known as the “Socorro seismic anomaly” (SSA). The Socorro seismic anomaly is the

68 M. Ibs-von Seht et al. / Tectonophysics 452 (2008) 66–77 Fig. 2. A (Left) Location map and seismicity of the Rio Grande rift. Solid lines: main tectonic lines of the rift (taken from Keller et al., 1991), circles: seismicity from 1962 to 2004 (M N 2), triangles: volcanoes, star: location of the Socorro seismic anomaly, dashed lines: approximate course of the basement structures Jemez lineament (JL) and Socorro Fracture zone (SFZ) (Sanford and Lin, 1998), dotted box: region of Fig. 3b, thin lines: rivers. Seismicity is based on catalogue data published by Sanford et al. (2002, 2006). B (Right) Shaded relief map and seismicity of the Socorro area of the Rio Grande rift. Circles and labels indicate clusters related to swarm periods. S = Socorro, S.A. = Santa Acacia.

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most seismically active area of the Rio Grande rift (Sanford et al., 2002). Swarms in the Socorro area have been reported since the beginning of the last century. From 1980 to 2004, at least seven individual swarm episodes can be identified (Sanford et al., 2002, 2006). The swarms occur in the middle of the rift in regions with an area measuring approximately 70 × 70 km (Fig. 2B). The duration of the swarms can be a few days up to many months. Up to ~500 events have been monitored during individual swarms. The hypocentral depths of the swarm events range from 3 km to 14 km. Depths greater than 10 km are restricted to a fault zone east of Socorro, although most of the seismic activity in the Socorro area cannot be correlated with surface features such as volcanics or faults. The maximum magnitudes occurring in the SSA are mostly smaller than 4.7, a swarm in 1906 had a maximum magnitude of 6.0 (Sanford et al., 1991). Magnitude– frequency distributions reveal b-values to be ~0.8 mostly, with some swarms showing b-values as low as 0.6. On the basis of an earthquake catalogue published by Sanford et al. (2002), we calculated a maximum daily rate of about 20 M N 1.3 events. Detailed information on individual swarms in the Socorro area can be found in numerous Geophysical Open File Reports of the New Mexico Institute of Mining and Technology (e.g. Ake, 1984; Jarpe, 1984; Jarmon, 1989). Besides the analysis of spatial and temporal distribution of the activity, focal mechanisms of events in the SSA have been studied (Wieder, 1981; Jarpe, 1984; Petrilla, 1987; Balch et al., 1994). Most of the solutions show N–S striking normal faults with a vertical P axis and the T axis trending WNW, which is consistent with the regional stress field of the Rio Grande rift. A minor group of events, mainly located near the fault zone east of

Fig. 3. Examples of focal mechanisms determined for swarm activities near Socorro in 1983 (modified after Balch et al., 1994).

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Socorro, exhibit strike slip mechanisms. There is some evidence that the mechanisms of a number of events are not pure doublecouple but have some volumetric components (Balch et al., 1995). Fig. 3 shows some examples of focal mechanisms observed during swarm activity in 1983. The swarm activity in the SSA is attributed to a mid-crustal magma chamber that was detected by the observation of reflected phases in microearthquake recordings (e.g. Sanford et al., 1973; Ake and Sanford, 1988). Interpretation of the seismic data suggest that the magma body forms a thin (~150 m) sill-shaped intrusion at about 19 km depth with a lateral extent of at least 3400 km2 (Balch et al., 1997, see Fig. 4), and possibly with a lower crustal extension acting as a conduit system (Schlue et al., 1996). The injection of new magma into the magma chamber is believed to cause extension in the overlying crust leading to the generation of earthquake swarms. The model of an inflating magma chamber has been supported by the detection of a longterm surface uplift of 2–3 mm/year in the Socorro area measured by levelling surveys (Larsen et al., 1986) and by interferometric radar imaging (Fialko and Simons, 2001). Fialko et al. (2001) calculated an inflation rate of about 0.006 km3/year for the Socorro magma body and emphasize that the volume increase may be due to a steady influx of magma from a deep source, as well as to the melting of the magma chamber roof. 3. Kenya rift The Kenya rift (Fig. 5A) represents a major part of the eastern branch of the East African rift system that extends from the Afar Triangle in the north to Mozambique in the south. Its central and southern parts form a classical 50 to 80 km wide rift valley with high escarpments on one side and en echelon fault steps on the opposite. Baker (1987) identified three major stages in the tectonic development of the Kenya rift: (1) the pre-rift stage (30– 12 Mya) with the forming of a depression and minor faulting (2) the half-graben stage(12–4 Mya) with the forming of the main boundary faults, and (3) the graben stage (b4 Mya) with an increase and an inward migration of faulting. All stages were accompanied by intense, mostly alkaline volcanism. Except for the Tanzanian Ol Doinyo Lengai, all volcanoes of the rift are extinct. However, some post volcanic hydrothermal activity can be observed in many places of the rift (Schlüter, 1997). There are a number of indications that the development of the Cenozoic Kenya rift and the present-day framework of fault lines were influenced by pre-existing basement structures (e.g. Braile et al., 1995). In particular, the central and southern parts of the generally NNE trending Kenya rift developed inside a broad, NW trending zone that is related to the suture between the Archean Tanzania Craton in the south-west and the Proterozoic Mozambique belt in the north-east. The craton margin zone is accompanied by several first-order NW trending and secondorder NNW–SSE trending lineaments, characterized by ductile and brittle shears and well documented in the Proterozoic rocks east and west of the rift. The shear zones are interpreted as major zones of weakness that were controlling rift propagation and accommodation of tectonic stress within several rift sectors (Smith and Mosley, 1993). At the intersections with the shear

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Fig. 4. Seismicity of the Socorro area, contours of estimated surface uplift between 1912 and 1980 (Larsen et al., 1986) and areal extent of the Socorro magma body estimated by Balch et al. (1997) (shaded region) and Hartse and Sanford (1992) (dotted line) (modified after Fig. 1b in Schlue et al., 1996).

zones the main course and faulting style of the rift change. On a regional scale, the rift sections influenced by the reworked craton margin are located roughly between Lake Baringo in the north and Lake Magadi in the south. On a local scale, NW–SE and NNW–SSE trending fractures that are assumed to be under basement control can be identified within the rift (Fig. 5A). These zones, described in detail by Smith and Mosley (1993), are the Kerio–Bogoria–Marmanet zone in the Lake Bogoria/Lake Baringo area and the Engorika–Magadi–Lembolos zone in the Lake Magadi area. Further discussions on oblique shear zones crossing the Kenya rift can be found e.g. in Baker et al. (1972) and Atmaoui and Hollnack (2003). Compared to the whole of the East African rift system, the Kenya rift shows a relatively low seismic activity. Except for a M=6.8 earthquake in 1928 in the central part of the rift, no MN 5 events connected to the rift have been reported since 1928. Nonetheless, numerous local earthquake studies since the 1960s have established a high microearthquake activity of the rift (e.g. Tongue et al., 1992). Almost the entire N–S trending rift system, including the E–W trending Kavirondo rift, shows a more or less high micro-seismic activity. Ibs-von Seht et al. (2001) reported a constant rate of ~10 Mb 3 events per day for the southernmost part of the Kenya rift.

Some of the microearthquake activity of the Kenya rift occurs in the form of earthquake swarms. During a three month local earthquake study in the Lake Bogoria region of the Kenya rift, Young et al. (1991) recorded more than 500 Mb 2.7 events. Some of the activity occurred during intense, short-lasting (~1 h) swarms within a small epicentral area. Furthermore, a number of multiplets were identified, and it was observed that the deeper events below 20 km were of unusual low frequency. Young et al. (1991) attributed the low frequency events to magma movements. A b-value of 0.87 was determined for the activity by Tongue et al. (1992). Tongue et al. (1994) observed a short-lasting (less than one day), low-magnitude (Mb 2) earthquake swarm in the Lake Baringo region. The swarm earthquakes were found to form a narrow, elongated cluster in the centre of the rift at ~5 km depth. Focal mechanism investigations suggested the presence of sub-vertical, N–S-trending fault planes in the hypocentral region and a WNW– ESE directed extensional stress field. The rift centre is characterized by recent volcanism and geothermal activity. An upper crustal low velocity zone, as had been established by a local earthquake tomography, was explained by elevated temperatures and the presence of fluids and the observed swarm activity was attributed to the emplacement of crustal dikes (Tongue et al., 1994).

M. Ibs-von Seht et al. / Tectonophysics 452 (2008) 66–77 Fig. 5. A (Left) Location map and seismicity of the Kenya rift. Solid lines: main tectonic lines of the rift (taken from Smith, 1994), circles: seismicity from 1998 to 2002 (M N 2.5), triangles: volcanoes, stars: locations of swarm earthquake areas, dashed lines: approximate course of the Precambrian basement structures Kerio–Bogoria–Marmanet fault zone (KBM) and Engorika–Magadi–Lembolos fault zone (EML), dotted box: region of Fig. 6B. Seismicity was recorded by the Seismological Network of the University of Nairobi, Kenya (Ibs-von Seht et al., 2002). B (Right) Shaded relief map and seismicity of the Lake Magadi area in the southern Kenya rift. Circles indicate precisely located epicentres (Ibs-von Seht et al., 2001). Note the linear shaped earthquake cluster North of the Lake which belongs to an intense swarm activity in May/June 1998. 71

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Fig. 6. Orientation of P-, B- and T-axes (left) and direction of extension (right) in the southern Kenya rift as inferred from fault plane solutions.

In the Lake Magadi area, an intense and long-lasting microearthquake swarm was recorded and analyzed by Ibs-von Seht et al. (2001). The swarm occurred in the centre of the rift, lasted for more than 2 months, and comprised more than 4000 events (Fig. 5B). In contrast to the background activity of about 10 events per day, a peak rate of more than 300 events per day was recorded during the swarm activity. The strongest swarm event had a magnitude of M = 4.2 and caused a surface crack of several kilometres length in the epicentral region (Ibs-von Seht et al., 2001; Atmaoui and Hollnack, 2003). The frequency–magnitude relation revealed a b-value of 0.75. Precise hypocenter locations using a 3-D velocity model showed that the clustered swarm earthquakes form a sub-vertical planar structure that follows the trend of the main rift valley axis. Compared to the background seismicity of this rift segment which was found to be restricted to depths greater than 10 km, the hypocentral depths of the swarm activity are generally shallow: Most of the hypocenters are between 0 and 6 km deep, a small fraction reaches 9 km depth. The volume occupied by the swarm activity is about 10 km long, 6 km deep and 2 km wide. The temporal development of the swarm activity was found to be clearly characterized by a migration of hypocenters in time from south to north. A progressive front of activity started at the southern tip of the cluster and moved at about 300 m per day in 14°NNE direction. At the same time, a slight increase in depth extent could be observed. The volume underneath the cluster was found to be virtually free of hypocenters. Fault plane solutions of the swarm events indicate predominantly normal faulting on fault planes aligned parallel to the rift valley axis, and a minor fraction of strike slip mechanisms. The faulting occurs in response to a WNW–ESE directed tensional stress field, which is in accordance with the general NNE–SSW alignment of the southern Kenya rift (Fig. 6). According to Ibs-von Seht et al. (2001) there are indications that the Magadi earthquake swarm can be attributed to a midcrustal magma body located in the centre of the rift. These are a locally upwarped brittle–ductile transition, a positive mid-crustal P-velocity anomaly overlain by a negative anomaly, a positive gravity anomaly, and the migration of activity which might be directly related to the movement of magma-related fluids (Fig. 7). 4. Eger rift The European Cenozoic rift system is a somewhat special case in regard to the occurrence of earthquake swarms, because

for the main branch of the rift, the Rhine Graben, swarms are virtually non-existent. However, the Eger rift is part of the European rift system (Ziegler, 1992) and is prominent for its earthquake swarms occurring in the NW-Bohemia/Vogtland area. The Eger rift is a major tectonomagmatic zone cutting through the Bohemian Massif. The formation of the rift took place during the Alpine orogenic phases between Late Eocene and recent times and has reactivated the variscan suture zone between Saxothuringicum and Moldanubicum. The rift system developed along this major crustal inhomogeneity with a ENE–WSW trend, and is approximately 50 km wide and 300 km long. The main rifting phase with the beginning of the graben formation and intraplate alkaline volcanism is attributed to the time span 42–9 Mya (Ulrych et al., 2003). Quaternary volcanism was active until Holocene, the youngest known volcanic activity occurred about 0.3–0.5 Mya (Wagner et al., 2002; Geissler, 2004). An oblique extension of the area led to the formation of the Eger Graben sedimentary basins at the western end of the Eger rift, while other large-scale structural features such as the NNW-trending Mariánské Lázně fault zone influenced the development of the rift. High emanations of CO2, manifested by mofettes and mineral springs (e.g. Kämpf, 1994; Weinlich et al., 1999; Weise et al., 2001; Bräuer et al., 2003), neotectonic movements (e.g. Bankwitz et al., 2003) as well as an unusual intraplate earthquake swarm activity in NW-Bohemia/ Vogtland (e.g. Fischer and Horálek, 2003; Klinge et al., 2003) indicate that the rifting process is still active. The seismicity of the NW-Bohemia/Vogtland area is largely dominated by recurring swarm activity (e.g. Fischer and Horálek, 2003; Klinge et al., 2003; Neunhöfer and Meier, 2004). Since 1552, earthquake swarms have been reported for the area (Karnik and Schenkova, 1988; Grünthal, 1989). The occurrence of the swarms is concentrated in several epicentral regions, among which the Klingenthal and Novy Kostel areas show the strongest activity (Fig. 8). The swarm areas are located several kilometres off the main rift axis. The rift centre shows virtually no seismic activity. While approximately 50 years ago, the Klingenthal area was more active, the strongest recent swarms in 1985/86, 1994, 1997 and 2000 took place in the Novy Kostel area. During the 2000 swarm, about 10,000 events were

Fig. 7. Profile sketch running perpendicular to the axis of the southern Kenya rift, illustrating a crustal structure model derived from local earthquake tomography and hypocentral depth studies. Circles: projections of hypocenters, dotted lines: crustal velocities derived from tomography, dashed lines: assumed depth ranges of active faults (after Ibs-von Seht et al., 2001).

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Fig. 8. Location map of the NW-Bohemia/Vogtland area in the western part of the Eger rift. Solid lines: main tectonic lines of the rift, triangles: volcanoes, circles: seismicity (ML N 0.5) from 1962 to 1999, stars: locations of main swarm areas, dashed lines: approximate course of the Mariánské Lázně fault zone (MLFZ). The seismicity map is based on the VOCATUS catalogue of Neunhöfer (2000).

recorded in a three month period. While previous swarms in the area had magnitudes of up to M = 4.6, the maximum magnitude of the 2000 swarm was ML = 3.7. We calculated b-values between 0.8 and 1.1 for the individual swarm periods. Fischer (2003) and Fischer and Horálek (2003) performed a precise relocation of the hypocenters that shows the activity is clustered in a volume of a few cubic kilometres and confined to a narrow, almost planar source volume describing a steeply dipping, N–S striking plane at 6 to 11 km depth. The N–S striking plane of hypocenters, as well as the lateral distribution of mofettes with evidence of CO2 degassing from the subcrustal mantle, match quite well with the newly found N–S trending Počatky–Plesná zone (Bankwitz et al., 2003). The morphologically more prominent Mariánské Lázně fault zone intersecting the Počatky–Plesná zone acutely, however, seems to be seismically inactive. Detailed source parameter and stress field investigations for the Novy Kostel swarm area have been carried out e.g. by Plenefisch and Klinge (2003). Fault plane solutions indicate mainly normal fault and strike slip mechanisms with nodal planes

striking between N–S and NNW–SSE, which is similar to the orientation of the Mariánské Lázně fault zone (Fig. 9). For the most recent swarm of 2000, we applied relative moment tensor inversions to derive the complete moment tensor. The solutions exhibit non-double-couple components of up to 30%. These results are in good agreement with those found by Dahm et al. (2000), Horálek et al. (2002) and Vavryčuk (2001, 2002) for the swarm of 1997. The non-double-couple components consisting of positive isotropic and compensated linear vector dipole components can be explained by tensile faulting indicating high pore fluid pressure. Several inversions for distinct subsets of the focal mechanisms of NW-Bohemia/Vogtland earthquakes revealed a local stress field similar to the general overall Central European stress field (Wirth et al., 2000; Vavryčuk, 2002; Plenefisch and Klinge, 2003). This means that dislocations caused by the swarm activities are controlled by the regional stress field of Central Europe which is a strike slip regime with a SE–NW directed σ1axis and a NE–SW directed σ3-axis. Similar results were obtained by Ibs-von Seht et al. (2006) for the Marktredwitz swarm area located in the western continuation of the Eger rift.

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anomalous upper mantle. Many authors emphasize the role of fluids in the generation of the swarms, which manifests in a high CO2 flux, He-anomalies, increased heat flow and frequent occurrence of mineral springs and mofettes in the area (e.g. Kämpf, 1994; Weise et al., 2001; Bräuer et al., 2003). Weinlich et al. (1999) assume a magmatic body at the base of the lower crust as the source of CO2 gas that ascends along a deepreaching fracture system at the intersection between the Eger rift and the Mariánské Lázně fault. Geissler et al. (2005) found a local crustal thinning underneath the area of interest using receiver function methods. On the basis of seismic and gas geochemical data Geissler et al. (2005) proposed a new tectonic model for the mantle–crust interaction in the NW-Bohemia/ Vogtland region. Therein, an active zone of mantle melting and presently active magmatic underplating beneath the western Eger rift are the sources for ascending CO2. According to Geissler et al. (2005) the triggering of earthquake swarms is caused by high pore fluid pressure in local captured upper crustal environment, high permeability in the crust enables the high permanent CO2 transport to the earth surface. A hypothesis proposed by Špičák et al. (1999) relates earthquake swarms to magmatic activity in the crustal layers. Injection of magma or related fluids is assumed to act as a triggering mechanism for hydraulic fracturing and the cause of the swarm earthquakes. Detailed seismo-statistical analyses of Hainzl and Fischer (2002) suggest that the activity of the 2000 swarm was initiated by a fluid intrusion, whereas the subsequent activity can be explained by stress transfer from previous events and associated pore pressure perturbations. 5. Discussion and conclusions

Fig. 9. Focal mechanisms of the 1994 and 1997 swarms and solutions of moment tensor inversions of the 2000 swarm near Novy Kostel.

The occurrence of seismic swarms in the NW-Bohemia/ Vogtland area is commonly attributed to a complex crustal structure extending to large depths in combination with an

A combination of geophysical observations for the three rifts discussed above reveals a number of similarities and agreements for the properties of earthquake swarms in these continental rifts. A compilation of relevant parameters is given in Table 1. In this section, we discuss the relation between the appearance of the earthquake swarms and the geological and structural settings. Perhaps the most striking similarity between the three rifts is the fact that swarms predominantly occur in those sections of the rifts that are influenced by large-scale fracture zones intersecting the main course of the rifts. The Rio Grande rift is crossed by

Table 1 Properties of swarm earthquake seismicity in continental rifts Rio Grande rift

Period regarded Max. rates (M N 1.3) Duration of swarms Max. magnitudes b-value Shape of active volume Area of scattered activity Depth Spatio-temporal migration of activity Faulting style Stress regime

Kenya rift

Eger rift

Socorro

Magadi

NW-Bohemia/Vogtland

1962–1995 ~20/day Days...months 4.7 ~0.8 Scattered spots N3000 km2 3–14 km Not observed Normal fault, strike slip Extensional

1998 (3 months) ~80/day Days...months 4.2 ~0.8 Planar sub-vertical None 0–9 km Clear Normal fault, strike slip Extensional

1985–2000 ~ 100/day Days...months 4.6 ~ 0.8–1.0 Planar sub-vertical, scattered spots N500 km2 6–11 km Diffuse Normal fault, strike slip Strike slip

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major fracture zones in the north near Los Alamos (Jemez Lineament) and further south near Socorro (Socorro fracture zone). While the Jemez lineament is marked by a large volcanic complex penetrating the rift, major recent surface volcanic features are missing at Socorro (Sanford and Lin, 1998). However, an extensive mid-crustal magma body which is responsible for intense earthquake swarms has been doubtlessly identified at Socorro. The Kenya rift is intersected by major shear zones near Lake Magadi and in the Lake Bogoria/Lake Baringo section of the rift. Both locations are characterized by the occurrence of earthquake swarms, and similar to the Rio Grand rift, there is strong evidence that the swarms are related to midcrustal magma bodies. Also in the Eger rift, the most seismically active segment is found at the intersection of the rift with a prominent shear zone (Mariánské Lázně fault zone). Here, intense earthquake swarms have been observed for centuries and various geophysical and geochemical observations indicate crustal magmatic activity. A general comparison of the Rio Grande and Kenya rifts demonstrates major similarities in their evolution, tectonism, and their present-day lithospheric structure (Keller et al., 1991). For both rifts there is strong evidence that rifting is still ongoing today. These similarities are also reflected in the occurrence of earthquake swarms observed in the two rifts. The swarm activity occurs in the rift centres and research finding indicates that this is controlled by magmatic activity related to the rifting process. Additional evidence for magmatic activity includes observations such as local surface uplifts and clear migration of a microearthquake front. In contrast, the evolution of the European rift system including the Eger rift was impeded by compressional stresses emanating from the Alpine orogen (Ziegler, 1992). Reduced active rifting processes in the Eger rift may be one reason that most of the seismic activity observed is not situated in the rift proper, but located at the newly found N–S trending Počatky–Plesná zone at the eastern margin of the Cheb basin. While for the Kenya rift, the emplacement of dikes along the rift axis has been confirmed by geological and geophysical studies (e.g. Swain, 1992), there is evidence for sill-shaped crustal intrusions in the Rio Grande rift. The differences in the shape and the depth of the assumed magmatic intrusions might also be reflected in the spatial distribution of the swarm earthquakes: A sill-shaped intrusion can lead to sources of seismic activity that are scattered in space as observed in Socorro, while dike intrusions will preferably lead to planar vertical arrangements of hypocenters like in the Magadi area. Here, also the unusual shallow depth of the swarm activity and a related opening of a surface crack points to the place-taking of an intrusion along the rift axis. The Eger rift swarm activity is characterized by several foci of activity distributed over an area of at least 500 km2. A great part of the activity is arranged along northerly-trending fault lines and shows a clear vertical planar hypocenter arrangement. Seismic and geochemical observations (e.g. Geissler et al., 2005) speak for presently active magmatic underplating beneath the NWBohemia/Vogtland region, which act as sources for fluids penetrating the crust along pre-existing zones of weakness. The statistical and magnitude properties of the swarm activity analyzed in the three rifts show some similarities. The

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maximum magnitudes observed during swarms range from 4.2 to 4.7 and the maximum seismicity rates for M N 1.3 earthquakes lie between 20 and 100 events per day. However, these values may be biased due to different seismic monitoring durations. The b-values for all three swarm areas are between 0.8 and 1.0, hence not deviating from a normal non-volcanic intraplate environment. This result is considered important because it indicates a clear difference between earthquake size distributions for swarms in continental rift zones and those in areas of active volcanism. The latter typically shows b-values significantly higher than 1, which is generally attributed to a highly fractured crust in connection with a heterogeneous stress system (Mogi, 1963). Focal mechanism studies of swarm earthquakes in the three rifts give similar and uniform pictures for stress field orientation and faulting style. All studies indicate that the dislocation patterns of swarm earthquakes in continental rifts correspond to the general regional stress regime of the respective areas, which is extensional in the Rio Grande and Kenya rifts and strike slip in the Eger rift. According to the stress field orientations, the faulting styles in the three rifts are dominated by normal fault and strike slip mechanisms. Thus, there is no indication that swarm earthquakes in continental rifts can be characterized by a local and heterogeneous stress system as is observed for active volcanoes. The results of our study suggest that the occurrence of earthquake swarms in continental rift zones is restricted to deepreaching zones of weakness that allow intrusion of upper mantle material into crustal layers. Magma-related fluids or fluctuations in the magma bodies themselves seem to cause the generation of earthquake swarms. Geophysical, especially seismological studies of swarm earthquakes contribute not only to a better understanding of the phenomenon itself, but also, records of swarm earthquakes are unique data sets that can serve to retrieve detailed information on the local crustal structure and stress field. For example, they can be used to identify magma bodies, as input for local earthquake tomography studies, to determine the spatial orientation of fault planes, and to improve stress field models. Swarm earthquakes are therefore an interesting tool for geophysical investigations of the Earth's crust and in active rifts in particular. Acknowledgements The authors wish to thank the reviewers Randy Keller and Wolfram Geissler for important comments that helped improve the manuscript, and Richard Aster who provided additional information on the Rio Grande rift for us. This work was supported by the Deutsche Forschungsgemeinschaft (DFG, KL 776/4-1). References Ake, J.P., 1984. An analysis of the May and July, 1983, Socorro mountain microearthquake swarms. Geophysics Open-file Report, vol. 49. New Mexico Institute of Mining and Technology. 74 pp. Ake, J.P., Sanford, A.R., 1988. New evidence for existence and internal structure of a thin layer of magma at mid-crustal depths near Socorro, New Mexico. Bull. Seismol. Soc. Am. 78, 1335–1359.

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Aldrich Jr., M.J., Chapin, C.E., Laughlin, A.W., 1986. Stress history and tectonic development of the Rio Grande rift, New Mexico. J. Geophys. Res. 91, 6199–6211. Atmaoui, N., Hollnack, D., 2003. Neotectonics and extension direction of the Southern Kenya rift, Lake Magadi. Tectonophysics 364, 71–83. Bachmann, G.O., Mehnert, H.H., 1978. New K–Ar dates and the late Pliocene to Holocene geomorphic history of the central Rio Grande region, New Mexico. Geol. Soc. Am. Bull. 89, 283–292. Baker, B.H., 1987. Outline of the petrology of the Kenya rift alkaline province. Geol. Soc. Lond. Spec. Publ. 30, 293–311. Baker, B.H., Mohr, P.A., Williams, L.A.J., 1972. Geology of the eastern rift system of Africa. Geol. Soc. Am. Spec. Pap. 136, 1–67. Balch, R.S., Sanford, A.R., Hartse, H.E., 1994. Focal mechanisms for two microearthquake swarms beneath Socorro Mountain, New Mexico, in May and July 1983. Geophysics Open-file Report, vol. 73. New Mexico Institute of Mining and Technology. 20 pp. Balch, R.S., Sanford, A.R., Jaksha, L.H., Hartse, H.E., House, L.S., 1995. Exotic seismic sources observed in the Central Rio Grande rift. Seismol. Res. Lett. 66 (2), 25. Balch, R.S., Hartse, H.E., Sanford, A.R., Lin, K.W., 1997. A new map of the geographic extent of the Socorro midcrustal magma body. Bull. Seismol. Soc. Am. 87, 174–182. Baldridge, W.S., Perry, F.V., Vaniman, D.T., Nealey, L.D., Leavy, B.D., Laughlin, A.W., Kyle, P., Bartov, Y., Steinitz, G., Gladney, E.S., 1991. Middle to late Cenozoic magmatism of the southeastern Colorado Plateau and central Rio Grande rift (New Mexico and Arizona, USA): a model for continental rifting. Tectonophysics 197, 327–354. Bankwitz, P., Schneider, G., Kämpf, H., Bankwitz, E., 2003. Structural characteristics of epicentral areas in Central Europe: study case Cheb basin (Czech Republic). J. Geodyn. 35, 5–32. Benoit, J.P., McNutt, S.R., 1996. Global volcanic earthquake swarm database 1979–1989. Open-file Report, vol. 69. U. S. Geological Survey. 31 pp. Braile, L.W., Keller, G.R., Wendland, R.F., Morgan, P., Khan, M.A., 1995. The East African rift system. In: Olsen, K.H. (Ed.), Continental Rifts: Evolution, Structure, Tectonics. Developments in Geotectonics, vol. 25, pp. 213–231. Bräuer, K., Kämpf, H., Strauch, G., Weise, S.M., 2003. Isotopic evidence (3He/4He, 13CCO2) of fluid-triggered intraplate seismicity. J. Geophys. Res. 108 (B2), 2070. doi:10.1029/2002JB002077. Dahm, T., Horalek, J., Sileny, J., 2000. Comparison of absolute and relative moment tensor solutions for the January 1997 West Bohemia earthquake swarm. Stud. Geophys. Geod. 44, 233–250. Déverchère, J., Petit, C., Gileva, N., Radziminovitch, N., Melnikova, V., San'kov, V., 2001. Depth distribution of earthquakes in the Baikal rift system and its implications for the rheology of the lithosphere. Geophys. J. Int. 146, 714–730. Fialko, Y., Simons, M., 2001. Evidence for on-going inflation of the Socorro magma body, New Mexico, from interferometric synthetic aperture radar imaging. Geophys. Res. Lett. 28, 3549–3552. Fialko, Y., Simons, M., Khazan, Y., 2001. Finite source modeling of magmatic unrest in Socorro, New Mexico, and Long Valley, California. Geophys. J. Int. 146, 191–200. Fischer, T., 2003. The August–December 2000 earthquake swarm in NWBohemia: the first results based on automatic processing of seismograms. J. Geodyn. 35, 59–81. Fischer, T., Horálek, J., 2003. Space-time distribution of earthquake swarms in the principal focal zone of the NW Bohemia/Vogtland seismoactive region: period 1985–2001. J. Geodyn. 35, 125–144. Geissler, W.H., 2004. Seismic and petrological investigations of the lithosphere in the swarm-earthquake region and CO2 degassing region Vogtland/NWBohemia. Doctoral thesis, Freie Universität Berlin. Geissler, W.H., Kämpf, H., Kind, R., Bräuer, K., Klinge, K., Plenefisch, T., Horálek, J., Zedník, J., Nehybka, V., 2005. Seismic structure and location of a CO2 source in the upper mantle of the western Eger (Ohře) rift, central Europe. Tectonics 24, TC5001. doi:10.1029/2004TC001672. Grünthal, G., 1989. About the history of earthquake activity in the focal region Vogtland/Western Bohemia. In: Bormann, P. (Ed.), Monitoring and Analysis of the Earthquake Swarm 1985/86 in the Region Vogtland/Western Bohemia. Akad. der Wissensch. der DDR, Potsdam, pp. 30–34.

Hainzl, S., Fischer, T., 2002. Indications for a successively triggered rupture growth underlying the 2000 earthquake swarm in Vogtland/NW-Bohemia. J. Geophys. Res. 107, 2338. doi:10.1029/2002JB001865. Hartse, H.E., Sanford, A.R., 1992. A new map of the Socorro magma body. Seismol. Res. Lett. 63, 69. Horálek, J., Šílený, J., Fischer, T., 2002. January 1997 earthquake swarm in West Bohemia (Czech Republic): moment tensors and their time and space variations. Tectonophysics 356, 65–85. Ibs-von Seht, M., Blumenstein, S., Wagner, R., Hollnack, D., Wohlenberg, J., 2001. Seismicity, seismotectonics and crustal structure of the southern Kenya rift — new data from the Lake Magadi Area. Geophys. J. Int. 146, 439–453. Ibs-von Seht, M., Gaciri, S.J., Barongo, J.O., 2002. The seismological network of Kenya. In: Nyamai, C., Maimba, M.M. (Eds.), Proceedings of the 8th and 9th Regional Conference on the Geology of Kenya. University of Nairobi, Nairobi, pp. 5–7. Ibs-von Seht, M., Plenefisch, T., Schmedes, E., 2006. Faulting style and stress field investigations for swarm earthquakes in NE Baveria/Germany — the transition between Vogtland/NW-Bohemia and the KTB-site. J. Seismol. 10. doi:10.1007/s10950-005-9008-5. Jarmon, C., 1989. Anatomy of earthquake swarms in the San Acacia, New Mexico Area. Geophysics Open-file Report, vol. 60. New Mexico Institute of Mining and Technology. 29 pp. Jarpe, S., 1984. Characteristics and possible cause of an earthquake swarm in the Central Rio Grande rift 28 km North of Socorro, New Mexico: February and March, 1983. Geophysics Open-file Report, vol. 50. New Mexico Institute of Mining and Technology. 54 pp. Kämpf, H., 1994. Vogtland earthquake swarms and their relation to reflection seismic indications. In: Behr, H.-J., Dürbaum, H.-J., Bankwitz, P. (Eds.), Crustal Structure of the Saxothuringian Zone: Results of Deep Seismic Profile MVE-90 (East). Z. Geol. Wiss., vol. 22, pp. 696–700. Karnik, V., Schenkova, Z., 1988. Comparison of the1985/86 events in western Bohemia with earlier Vogtland swarms. In: Prochazka, D. (Ed.), Induced Seismicity and Associated Phenomena, Geophys. Inst. of Czechosl. Acad. Sci., Praha, pp. 328–342. Karpin, T.L., Thurber, C.H., 1987. The relationship between earthquake swarms and magma transport: Kilauea Volcano, Hawaii. Pure Appl. Geophys. 125 (6), 971–991. Keller, G.R., Khan, M.A., Morgan, P., Wendland, R.F., Baldridge, W.S., Olsen, K.H., Prodehl, C., Braile, L.W., 1991. A comparative study of the Rio Grande and Kenya rifts. Tectonophysics 197, 355–371. Klinge, K., Plenefisch, T., Stammler, K., 2003. The earthquake swarm 2000 in the region Vogtland/NW-Bohemia-earthquake recording at German stations and temporal distribution of events. J. Geodyn. 35, 83–96. Larsen, S., Reilinger, R., Brown, L., 1986. Evidence for ongoing crustal deformation related to magmatic activity near Socorro, New Mexico. J. Geophys. Res. 91, 6283–6292. Mogi, K., 1963. Some discussions on aftershocks, foreshocks and earthquake swarms — the fracture of a semi-infinite body caused by an inner stress origin and its relation to the earthquake phenomena, 3. Bull. Earthq. Res. Inst. Tokyo Univ. 41, 615–658. Morgan, P., Seager, W., Golombek, M., 1986. Cenozoic thermal, mechanical and tectonic evolution of the Rio Grande rift. J. Geophys. Res. 91, 6263–6276. Neunhöfer, H., 2000. The catalogue of the Vogtland/Western Bohemia earthquakes and VOCATUS as an auxiliary code. Stud. Geophys. Geod. 44, 549–555. Neunhöfer, H., Meier, T., 2004. Seismicity in the Vogtland/Western Bohemia earthquake region between 1962 and 1998. Stud. Geophys. Geod. 48, 539–562. Olsen, K.H. (Ed.), 1995. Continental Rifts: Evolution, Structure, Tectonics. Developments in Geotectonics, vol. 25. Elsevier, Amsterdam. 466pp. Olsen, K.H., Baldridge, W.S., Callender, J.F., 1987. Rio Grande rift: an overview. Tectonophysics 143, 119–139. Petrilla, C.C., 1987. Analysis of recent swarm activity along the Hot Springs– Montosa fault East of Socorro, New Mexico. Geophysics Open-file Report, vol. 58. New Mexico Institute of Mining and Technology. 53 pp. Plenefisch, T., Klinge, K., 2003. Temporal variations of focal mechanisms in the Novy Kostel focal zone (Vogtland/NW-Bohemia) — comparison of the swarms of 1994, 1997 and 2000. J. Geodyn. 35, 145–156.

M. Ibs-von Seht et al. / Tectonophysics 452 (2008) 66–77 Sanford, A.R., Lin, K.W., 1998. Evidence for a 1400 km long Socorro fracture zone. Geophysics Open-file Report, vol. 89. New Mexico Institute of Mining and Technology. 18 pp. Sanford, A.R., Alptekin, Ö., Toppozada, T.R., 1973. Use of reflection phases on microearthquake seismograms to man an unusual discontinuity beneath the Rio Grande rift. Bull. Seismol. Soc. Am. 63, 2021–2034. Sanford, A.H., Jaksha, L.H., Cash, D.J., 1991. Seismicity of the Rio Grande rift in New Mexico. In: Slemmons, D.B., Engdahl, E.R., Zoback, M.D., Blackwell, D.D. (Eds.), Neotectonics of North America: Geological Society of America, Decade Map 1, pp. 229–244. Sanford, A.R., Lin, K.W., Tsai, I.C., Jaksha, L.H., 2002. Earthquake catalogs for New Mexico and bordering areas: 1869–1998. New Mexico Bureau of Geology and Mineral Resources, Circular 210. Sanford, A.R., Mayeau, T.M., Schlue, J.W., Aster, R.C., Jaksha, L.H., 2006. Earthquake catalogs for New Mexico and bordering areas II: 1999–2004. N. M. Geol. 28 (No. 4), 99–109. Schlue, J., Aster, R., Meyer, R., 1996. A lower crustal extension to a midcrustal magma body in the Rio Grande rift, New Mexico. J. Geophys. Res. 101 (B11), 25,283–25,291. Schlüter, T., 1997. Geology of East Africa. Beiträge zur Regionalen Geologie der Erde, vol. 27. Gebrüder Bornträger, Berlin-Stuttgart. 484pp. Smith, M., 1994. Stratigraphic and structural constraints on mechanisms of active rifting in the Gregory rift, Kenya. In: Prodehl, C., Keller, G.R., Khan, M.A. (Eds.), Crustal and Upper Mantle Structure of the Kenya rift. Tectonophysics, vol. 236, pp. 3–22. Smith, M., Mosley, P., 1993. Crustal heterogeneity and basement influence on the development of the Kenya rift, East Africa. Tectonics 12, 591–606. Špičák, A., 2000. Earthquake swarms and accompanying phenomena in intraplate regions: a review. Stud. Geophys. Geod. 44, 89–106. Špičák, A., Horálek, J., Bouskova, A., Tomek, Č., Vanek, J., 1999. Magma intrusions and earthquake swarm occurrence in the Western part of the Bohemian Massif. Stud. Geophys. Geod. 43, 87–106. Swain, C.J., 1992. The Kenya rift axial gravity high: a re-interpretation. Tectonophysics 204, 49–70. Tongue, J.A., Maguire, P.K.H., Young, P.A.V., 1992. Seismicity distribution from temporary earthquake recording networks in Kenya. Tectonophysics 204, 71–79. Tongue, J.A., Maguire, P.K.H., Burton, P., 1994. An earthquake study in the Lake Baringo basin of the central Kenya rift. Tectonophysics 236, 151–164.

77

Ulrych, J., Lloyd, F.E., Balogh, K., 2003. Age relations and geochemical constraints of Cenozoic alkaline volcanic series in W Bohemia: a review. Geolines 15, 168–180. Vavryčuk, V., 2001. Inversion for parameters of tensile earthquakes. J. Geophys. Res. 106, 16,339–16,355. Vavryčuk, V., 2002. Non-double-couple earthquakes of 1997 January in West Bohemia, Czech Republic: evidence for tensile faulting. Geophys. J. Int. 149, 364–373. Wagner, G.A., Göken, K., Jonckheere, R., Wagner, I., Woda, C., 2002. Dating of quaternary volcanoes Komorni Hurka (Kammerbühl) and Zelezna Hurka (Eisenbühl), Czech Republic, by TL, ESR, alpha recoil and fission track chronometry. Z. Geol. Wiss. 30 (3), 191–200. Weinlich, F.H., Bräuer, K., Kämpf, H., Strauch, G., Tesař, J., Weise, S.M., 1999. An active subcontinental mantle volatile system in the western Eger rift, Central Europe: gas flux, isotropic (He, C, and N) and compositional fingerprints. Geochim. Cosmochim. Acta 63, 3653–3671. Weise, S.M., Bräuer, K., Kämpf, H., Strauch, G., Koch, U., 2001. Transport of mantle volatiles through the crust traced by seismically released fluids: a natural experiment in the earthquake swarm area Vogtland/NW Bohemia, Central Europe. Tectonophysics 336, 137–150. Wieder, D.P., 1981. Tectonic significance of microearthquake activity from composite fault-plane solutions in the Rio Grande rift near Socorro, New Mexico. Geophysics Open-file Report, vol. 37. New Mexico Institute of Mining and Technology. 68 pp. Wilson, D., Aster, R., West, M., Ni, W., Grand, S., Gao, W., Baldridge, W.S., Semken, S., Patel, P., 2005a. Lithospheric structure of the Rio Grande rift. Nature 433. doi:10.1038/nature03297. Wilson, D., Aster, R., Ni, J., Grand, S., West, M., Gao, W., Baldridge, W.S., Semken, S., 2005b. Imaging the seismic structure of the crust and upper mantle beneath the Great Plains, Rio Grande rift, and Colorado Plateau using receiver functions. J. Geophys. Res. 110, B05306. doi:10.1029/2004JB003492. Wirth, W., Plenefisch, T., Klinge, K., Stammler, K., Seidl, D., 2000. Focal mechanisms and stress field in the region Vogtland/Western Bohemia. Stud. Geophys. Geod. 44, 126–141. Young, P.A.V., Maguire, P.K.H., Laffoley, d'A., Evans, J.R., 1991. Implications of the distribution of seismicity near Lake Bogoria in the Kenya rift. Geophys. J. Int. 105, 665–674. Ziegler, P.A., 1992. European Cenozoic rift system. Tectonophysics 208, 91–111.