Interactive comment on “ Seismicity at the Rwenzori Mountains , East African Rift : earthquake distribution , magnitudes and source mechanisms ”

This manuscript addresses a heterogeneous distribution of seismicity and focal mechanism within the Rwenzori Mountains area in the western branch of the East African Rift. Deploying temporary seismic network, many earthquakes seem to be relocated with higher accuracy than previous studies. Vertical sections of the seismicity indicate that the maximum focal depths vary significantly within the Rwenzori area. It is interesting that there is no seismicity within the uppermost 10 km beneath the rift valley, whereas beneath the rift shoulders seismicity range from the surface to the Moho depth. The authors discuss this lateral variation of seismicity linking with that of Moho

The East African Rift System (EARS) separates the Nubian and the Somalian plates and represents one of the world's largest continental rift structures (Bendick et al., 2006;Chorowicz, 2005;Ebinger, 1989) it extends approximately in N-S direction through East Africa consisting of a series of connecting rifts segments (Fig. 1).South of the Main Ethiopian Rift the system splits up into two branches, the Albertine Rift in the west and the Kenya Rift in the east.They enclose the Tanzania Craton which is surrounded by Proterozoic mobile belts (Nyblade and Brazier, 2002).Further south the system continues as the Malawi Rift and terminates at the coast of southern Mozambique.Globally recorded earthquakes (Fig. 1) as well as compilations on regional and local earthquake studies have shown that the western branch of the EARS exhibits significantly higher seismic activity than the eastern branch (Midzi et al., 1999;Twesigomwe, 1997).Numerous publications focus on the hypocentral depth distribution, which turns out to extend down to the crustal base -in contrast to other seismically active continental regions where earthquakes are restricted to the upper crust (e.g.Albaric et al., 2009;Camelbeek and Iranga, 1996;Foster and Jackson, 1998;Nyblade and Langston, 1995;Shudofsky et al., 1987;Young et al., 1991).
The 5000 m high Rwenzori Mountains are situated within the Albertine rift, in an area with the highest seismic activity of the whole rift system.They represent a non-volcanic basement block whose origin and relation to the evolution of the EARS are focus of the RiftLink project (www.riftlink.org).Despite of the high activity, there is only little knowledge on the local seismicity in the Rwenzori region.An early microearthquake survey in the Rwenzori area (Maasha, 1975) was based on four portable seismic stations.
They were operated at 28 different sites during June to September 1973 moving the instruments after a period of 4 to 20 days to new locations.An average of 5 to 20 events per day was recorded in the area with magnitude estimates between -2 and 4 and focal depth reaching down to 25 to 40 km.However, these depth estimates exhibit large uncertainties because of the limited number of stations.A small number of composite fault plane solutions indicated normal faulting in the Rwenzori area.A more recent study by Tugume and Nyblade (2009)  depth extent and its implication on the heat flow in the area.The observed focal depth distribution exhibits a maximum number of earthquakes at 16 km and falls off sharply at 22 km extending down to 31 km.In all of the above investigations the depth extent of seismicity is restricted to the crust.However, Lindenfeld and R ümpker (2011) have reported the detection of mantle earthquakes to the northeast of the Rwenzori Mountains, which are likely related to magmatic intrusions in relation to lithospheric break up.
In the present paper we provide a comprehensive characterization of the seismic activity within the whole Rwenzori region, covering a substantially larger area and time period than previous investigations.For the first time we present magnitude distributions and reliable source mechanisms of a large set of earthquakes, together with a detailed discussion of the complex hypocentral depth distribution.These data are essential for a better understanding of the rifting processes in this part of the EARS and will improve our general insight into the mechanisms of active rifts.

Seismic network and data analysis
The data we present in this study were moved to new places so that the network in total consisted of 35 station locations (Fig. 2).

Hypocentre location
In order to detect and extract the numerous local events we applied a STA/LTA trigger algorithm to the continuous data streams.Prior to this procedure the data were filtered with a Butterworth high pass at 1 Hz.Tests were performed to determine the optimum set of trigger parameters, i.e. to identify a maximum number of earthquakes without too much false trigger events.Time windows for short-and long-term average (STA, LTA) were set to 3 s and 30 s respectively, a ratio of STA/LTA = 3.0 was chosen to trigger an event.The routine data analysis was done with the SEISAN software package (Havskov and Ottem öller, 1999), including an implementation of HYPOCENTER, a location algorithm by Lienert and Havskov (1995).Arrival times were manually picked on the unfiltered seismograms to avoid phase shift effects, P waves on the vertical component and S on one of the horizontal components.For earthquake location we used the velocity model UVI-N74 of Bram (1975) with a fixed vp/vs ratio of 1.74 (Table 1).The Bram model was derived from local and regional P n phases travelling predominantly parallel to the Albertine rift.To test the limits of such a relatively simple 1-D velocity model we relocated the hypocentres using a synthetic 3D model with several crustal velocity anomalies between -7 % and +7 % (Jakovlev, 2010).The average deviation of the relocated hypocentres compared to the 1-D locations was about 1.30 km, which is smaller than the estimated 1-D location error (<3 km, see below).Furthermore, 3D velocity anomalies from travel-time tomography are only well constrained in the centre of the seismic station network with relatively poor vertical resolution (Jakovlev et al., 2011).For these reasons we think that the use of a 3-D velocity model would not improve the location precision of the Bram model.
To give an impression of the data quality we present seismograms of two earthquakes (Fig. 3).The first event (top row) was relatively weak (M L = 0.9).The plotted seismograms were recorded at stations KARU (∆ = 12.9 km) and KISA (∆ = 49.2 km).Figures

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Full Both stations show clear P and S-onsets with good signal/noise ratios that enable reliable arrival time pickings (dashed lines).The second event (bottom row) was relatively strong (M L = 3.8).Records are from stations KASS (∆ = 2.8 km) and KISA (∆ = 49.3 km).
In general it was possible to determine P-arrival times with a precision of ±0.05 s.
We assume a slightly reduced accuracy for S-wave arrivals of ±0.1 s.This is due to the lower frequency content of the S-phases and because in many cases the true Sonset is blurred by the P-wave coda.During the location process the SEISAN software provides estimates of location errors based on travel time residuals and the station network geometry.The RMS residuals range between 0.1 s and 0.2 s in most cases.
The majority of the events (ca.60 %) were located with an error better than 3 km.In the following analysis we use only events with location errors ≤5 km which is 76 % of all located events.

Magnitude determination
Local magnitudes were determined using the formula of Hutton and Boore (1987): dependence of the calculated magnitudes.This suggests that the calibration parameters of Alsaker et al. (1991) give an appropriate description of the attenuation in the Rwenzori area.For comparison we re-calculated the magnitudes using calibration parameters that were derived for Tanzania (Langston et al., 1998) and the original values for Southern California (Hutton and Boore, 1987).The results are also plotted in Fig. 4 and show no significant deviation from the magnitudes calculated with the default parameters of SEISAN.

Temporal variations of seismic activity
The analysis of the recorded data revealed high microseismic activity in the Rwenzori region.Figure 5 presents the monthly numbers of located earthquakes which correlates fairly well with the number of operating stations.Starting in February 2006 the number of events increases until July when the expansion of the station network was finished.Between August 2006 and September 2007, when the network was running in a more or less steady state, we located on average 766 events per month, 613 of them with errors ≤ 5 km (light and dark bars, respectively).We are not able to identify any clear seasonal variations of the seismicity rate.However, we observe a distinct dependence on the time of day (Fig. 6).There is an apparent breakdown in seismic activity between 04:00 and 17:00 UT which corresponds to 07:00-21:00 LT.The number of located events decreases to less than 25 % of the night time values.We associate this observation with an increased noise level during day time.During this period the sensitivity of the seismometer network is considerably reduced as a result of human activity, such as field work and traffic, which starts shortly after sunrise (ca.06:00) and ends a few hours after sunset (ca.18:00).However, this affects only events with magnitudes smaller than 1.5 whereas the number of larger earthquakes is more or less constant over the whole day.

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Full 3 Results and discussion

Hypocentre distribution
In total we located almost 14 000 earthquakes from February 2006 to September 2007 in the Rwenzori area.The seismicity map (Fig. 2) includes ca. 10 600 events with location errors ≤5 km.Only few earthquakes are located within the Rwenzori block.The majority of the epicentres lie within fault zones to the east and west of the mountain range.This may possibly indicate that the mountains can be considered as a decoupled rigid block that is surrounded by two rift segments and resistant to seismogenic deformations as proposed by Koehn et al. (2008Koehn et al. ( , 2010)).However, between 0.5 • and 0.6 • latitude, we detected a narrow band of seismicity, separating the main Rwenzori block in the south from its northern ridge, see dotted ellipse in Fig. 2.This is probably an indication of stresses and relative movements acting between the two mountain segments.The highest seismic activity is observed within a region east of the Rwenzoris between Lake Albert and Lake George, where the mountains are connected to the eastern rift flank.The majority of these events are grouped in clusters rather than in planar structures.Only in few cases it is possible to associate the epicentre distribution with known faults.At the north-eastern tip of the mountains, between 0.8 • -1.0 • latitude and 30.2 • -30.5 • longitude, we recorded high seismicity aligned with the N-E trending eastern rift flank.These events are probably located at the southern segment of the Toro-Bunyoro fault (Fig. 2), which runs in NE direction along the eastern side of the Lake Albert Basin (e.g.Ring, 2008).However, at about 1.0 • N seismic activity on the fault ends abruptly and continues more diffuse into the Lake Albert Rift.In the south-western parts of the Rwenzori region we observed only weak seismic activity.As will be shown in Sect.seismic activity in this region is actually weak or has a similar high level as observed in the north-eastern parts of the Rwenzori area.

Depth distribution
In Fig. 7 we present a histogram of the general hypocentral depth distribution in the Rwenzori area.Seismicity extends from the surface down to a depth of 32 km with a clear maximum of activity at 15 km.This agrees well with earlier observations leading to the conclusion that in East Africa earthquakes occur throughout the crust, but not deeper than the Mohorovicic discontinuity (Albaric et al., 2009;Brazier et al., 2005;Foster and Jackson, 1998).While in most other continental regions seismicity is restricted to the upper crust, similar focal depth distributions throughout the complete crust are also reported for the Baikal Rift (Chen and Molnar, 1983;Emmerson et al., 2006;Jackson et al., 2008;Maggi et al., 2000).D éverch ère et al. ( 2001) find a seismogenic lower crust with significant earthquake activity down to 35-40 km and only few earthquakes below the Moho in the Baikal Rift system, very similar to our results.From modelling of the brittle-ductile transition in the crust they infer a quartz rheology at 0-10 km depth and a diabase rheology at 10-45 km depth.For a more detailed analysis of the hypocentral depth distribution we compiled several vertical sections of seismicity transecting the area in different places and directions (Fig. 8).Four profiles (A-D) are oriented more or less perpendicular to the rift axis, whereas profile E runs parallel to the rift from NNE to SSW.Black dots in Fig. 8 mark the background seismicity; coloured symbols represent hypocentres that were projected onto the respective planes.The vertical sections are presented in Fig. 9. Additionally to the hypocentres we have marked the Moho depths determined at several stations by teleseismic receiver functions (W ölbern et al., 2010).The locations of these stations are also indicated in Fig. 8.In contrast to the overall depth distribution of -Profile A: This profile crosses the Albertine Rift north of the Rwenzori mountains and ends up at the rift shoulder to the east.Beneath the shoulder (35-45 km on horizontal profile axis) seismicity is concentrated along a vertical oriented plane reaching from the surface down to 30 km.In this area the Moho depth was determined at 34 km (station KABE) indicating that all events occur within the crust.
On the western side however (5-30 km), beneath the rift valley, hypocentres are restricted between 10 km and 25 km depth.The Moho is also raised up to 25 km depth at SEML. (Moho depth at SEML was originally published with 34 km.After reinterpretation of the receiver functions the authors modified the Moho depth at this site to 25 km, (W ölbern, personal communication, 2011)).We observe this pattern of hypocentral depth distribution in all profiles.There are two general features.
(1) Beneath the rift shoulders or in areas where the Rwenzoris are connected to the shoulders seismicity extends from the surface down to ca. 30 km depth.
(2) Beneath the rift valley we observe no seismic activity within the uppermost 10 km together with reduced depth extension.The absence of shallow seismicity beneath the rift graben is also observed at Lake Bogoria (Young et al., 1991) and Lake Magadi (Ibs-von Seht et al., 2001), both located at the southern end of the Kenya rift.They attributed the activity deeper than 10 km to deep faults in the basement, buried beneath the rift.
- between 53 km and 60 km (dashed circle) which is clearly below the Moho.These events occurred during a period of 14 days with local magnitudes between 1.4 and 2.4.We are not aware of similarly deep earthquakes in other regions of the EARS.A detailed study of these events can be found in Lindenfeld and R ümpker (2011).
-Profile C: This profile traverses the Rwenzoris in an area where the main block is connected with the northern ridge.Here we observe a strip of seismicity crossing the mountains in E-W direction (30-40 km on horizontal profile axis) with hypocentres between 10 km and 20 km depth.On the eastern side seismicity again extends from the surface down to 25 km.Moho depths at NGIT and KASS are 25 km and 24 km, respectively.
-Profile D: This depth section starts beneath the central peaks of the Rwenzori block and crosses the Edward rift towards the eastern rift shoulder.Again, there is no seismic activity in the uppermost 10 km beneath the rift graben -similar as already observed beneath the Albert rift in profiles A-C.Moho depths close to the mountain (KILE, RUBO) are approximately 22 km which is about 5 km deeper than the maximum depth of seismicity in this area.Further to the east seismicity deepens systematically down to more than 25 km.At the end of the profile, beneath the eastern rift shoulder the Moho discontinuity is located at 32 km depth (NYAN, RWEB).
-Profile E: This profile runs east of the Rwenzoris, roughly in N-S direction, parallel to the northward penetrating Edward rift.Again, there is no seismic activity within the uppermost 10 km beneath the rift valley.At its northern end (E2), where the profile leaves the tip of the rift, seismicity extends from the surface down to about 25 km depth.Further south the hypocentres concentrate at 15 km depth.Then the seismic activity splits up into two horizontal levels at 15 km and 25 km depth, respectively.There is a remarkable coincidence between Moho positions and the focal depth extension, although seismicity fades out and has a more diffusive 575 Introduction

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Full character towards the southern profile end.This is certainly due to the fact that we have to expect less location accuracy for events at the border of the station network.The seismicity seems to trace the upper and lower edge of a low velocity zone (LVZ) that was identified by receiver function inversion at station MWEY below 15 km depth and can be seen also beneath KASE and NYAN at similar depth but less pronounced (W ölbern et al., 2010).From modelling the authors derive a strong velocity degrease which is unlikely caused by compositional changes of the crustal rock (see S-velocity graph in Fig. 9).W ölbern et al. ( 2010) interpret this channel as a possible zone of partial melt within the crust related to local volcanic fields in the area.Reduced brittleness of the lower crust caused by increased temperature could explain the observed decrease of seismic activity in this area.

Magnitude distribution
Local magnitudes were calculated as specified in Sect.2.2.The results for the complete data catalogue range from M L = -0.5 up to M L = 5.1 with a b-value of ∼1.1 and a magnitude of completeness M C = 1.3, as derived from the magnitude-frequency distribution.However, the map in Fig. 10 presents a highly heterogeneous spatial distribution of the calculated magnitudes, which is obviously the result of the unevenly spaced station network.The smallest events with magnitudes M L ≤1.5 (green and blue symbols) could only be located in the north-eastern area where the network has its highest density.In contrast, events with M L > 1.5 are located more evenly over the whole region.
This implies that the apparently reduced seismic activity southwest of the Rwenzoris is probably due to the lack of stations in this area.The data of a recently deployed seismic network west of the mountains will prove if there is indeed similar high activity here as in the north-eastern region.
The heterogeneous spacing of the station network implies also that the completeness of the recorded data set strongly depends on the respective region.Here, the inter-station distances are relatively large and M C is approximately 1.5 (triangles).The b-value for both magnitude distributions is 1.1 which is slightly larger than the values published by Kebede and Kulh ánek (1994) for the western branch of the East African Rift (0.9-1.0).According to Mogi (1967) b-values between 0.5 and 1.0 are indicative of tectonic processes in non-volcanic intraplate areas.Relatively large b-values (1.0-2.5) are usually observed in regions with active volcanism where seismicity is associated with eruptions and magma movements (e.g.Karpin and Thurber, 1987).Our results may suggest that at least a part of the recorded seismicity in the Rwenzori Region has its origin in volcanic-magmatic processes within the crust.

Source mechanisms
Fault plane solutions were derived from P-wave polarities and SV/P amplitude ratios.
To achieve good data coverage we processed only events that were recorded by at least 12 stations.In total we analysed 304 events located in the Rwenzori area.P wave polarities were manually picked on the raw seismograms.Then we used a grid search algorithm to determine the orientation of the nodal planes (FOCMEC by Snoke et al., 1984).In general we were able to derive solutions without admitting P polarity errors.However, in many cases the nodal planes covered large angle ranges and it was impossible to derive a unique solution for the respective event.With P-polarities alone we were able to determine reliable source mechanisms for only 40 % of the selected earthquakes.To further constrain the nodal planes, we integrated also SV/P amplitude ratios into the grid search.In order to minimize the effects of different sensor bandwidths and local receiver conditions on the recorded amplitudes we pre-processed the signals with a Butterworth band pass in the 0.1-5.0Hz frequency range.Amplitudes were then picked on the vertical seismogram component.The combination of P-polarities with SV/P ratios enabled us to derive more reliable fault plane solutions than with Ppolarities alone.by 15 stations in total.At 14 stations we were able to determine the P-polarities: 12 compressional (black symbols) and 2 dilatational onsets (open symbols).The result of the grid search offers clear ambiguity (Fig. 12a).The nodal planes are consistent with two normal faulting mechanisms with nearly perpendicular oriented strike directions making it impossible to allocate a single mechanism to this event.However, performing a grid search with P polarities together with SV/P amplitude-ratios restricts the possible source mechanism to normal faulting with NNE-SSW strike direction (Fig. 12b).
With this method we substantially increased the number of reliable solutions to almost 70 % of the analysed events, compared to 40 % using only P-polarities.Furthermore, in many cases we were able to improve the quality and significance of the P-polarity solutions.
A classification of the derived fault plane solutions on the basis of P-and T-axis plunge angles is presented in Fig. 13.Data points that are close to the three vertices of the graph represent pure normal faulting (red symbols), strike-slip (blue symbols), and reverse faulting (green symbols).Additionally, we classified normal faults with strikeslip component (magenta symbols), reverse faults with strike-slip component (cyan symbols), and odd mechanisms (gray symbols).As is expected for a rift structure with extensional stress regime the majority of the derived source mechanisms represent pure normal faults or normal faults with strike slip component (72 %), whereas only 6 % of all analysed events have pure reverse faults or reverse faults with strike-slip component.8 % of the events have pure strike-slip mechanisms.
Figure 14 shows a map with all derived fault-plane solutions.The events are arranged into six groups in order to identify any systematic pattern.For each group the directions of the tension axes (T-axis trends) are plotted in a rose-histogram using sector widths of 15 • .The results reveal a remarkable uniform picture for the orientation of the extensional stress field, particularly in groups 1, 3, and 4 which are located west and east of the Rwenzori Mountains.The average orientation of the T-axes in these groups is WNW-ESE -perpendicular to the rift axis -and in good agreement with kinematic rift models (Stamps et al., 2008(Stamps et al., , 2010)).The two southernmost groups, 5 and 578 Introduction

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Full 6, show the same characteristic, but additionally there is a second prevailing T-axis orientation in WSW-ENE direction in both histograms.Group 2 is located northeast of the Rwenzori, immediate at the eastern rift boundary fault, and reveals a complete different behaviour.T-axis trends are predominantly oriented NEN-SWS, which is more or less vertical to the T-axis orientation of all other groups.This might be an indication that the stress field in this area is controlled by local disturbances.

Conclusions
From February 2006 until September 2007 we have monitored the local seismicity in the Rwenzori Region with a network of 29 seismic stations.During this period a total of 13 964 events were located with local magnitudes ranging from -0.5 up to 5.1.
Magnitude-frequency relations show a b-value of 1.1 which is slightly above the values reported by other authors for the western rift.The magnitude of completeness, M c , was derived for several subregions and exhibits a significant dependence on station density.M c = 1.5 in the southern parts of the area and M c = 0.8 in the northern parts where the average inter-station distance is considerably smaller.
The general focal-depth distribution shows that there is considerable seismic energy release down to 32 km depth with a pronounced peak of activity at 15 km.Typically, there is no seismicity within the uppermost 10 km beneath the rift valley, whereas beneath the rift shoulders seismicity extends close to the surface.Vertical sections indicate that the maximum focal depths vary significantly within the Rwenzori area.However, they correlate fairly well with Moho depths that were derived by teleseismic receiver functions and range between 21 km and 34 km.This confirms that the observed seismicity is restricted to the crustal part of the lithosphere, with the exception of seven events at a depth between 53 and 60 km as reported by Lindenfeld and R ümpker (2011).There is no indication for a crustal root beneath the Rwenzori Mountains, neither from receiver functions nor from focal depth distribution.Fault plane solutions were derived for 304 events using P-polarities and SV/Pamplitude ratios.The majority of the source mechanisms are normal faults, only 2 % of the analysed events are pure thrust faults.The T-axis trends give a uniform picture for the pattern of the extensional stress field.They are aligned in WNW-ESE direction, more or less vertical to the rift axis which is in good agreement with the expected rift process and also with kinematic rift models.Seismicity southeast of the Rwenzoris is less resolved due to the inhomogeneous station distribution.Essential information on microseismic activity will be provided in the future by a recently deployed seismic network which covers also the area west and southwest of the Rwenzori Mountains.Introduction

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Full  Full Geosciences, Goethe-University, Frankfurt am Main, Germany 2 Department of Geology, Makerere University, Kampala, Uganda 1 Introduction . Starting at the Afar triple junction in the north, Discussion Paper | Discussion Paper | Discussion Paper | was restricted to the northern end of the Rwenzori Mountains.The authors present the seismic activity between 2006 March and June based on a network of eight stations and they investigate the hypocentral Discussion Paper | Discussion Paper | Discussion Paper | recorded from February 2006 to September 2007 by a temporary network of 29 seismic stations covering an area of 140 × 80 km 2 in the Rwenzori Mountains region.The instruments were installed in the framework of the RiftLink Project, however, some data were also provided by the BGR/GSMD GEOTHERM-Subprogram (German Geological Survey, BGR, Geological Survey and Mines Department, GSMD, Uganda).Until April 2006 a small network of nine 3-component short-period instruments was running in the northern part of the area equipped with Reftek data loggers sampling the data at 100 Hz in triggered mode.Starting in May 2006 the net was expanded by 20 stations equipped with EDL data loggers recording continuously at a sample rate of 100 Hz.Seven of these stations were equipped with broadband instruments (Guralp CMG-3T), the remainder with short-period seismometers (Mark L-4C3D).Occasionally some instruments were Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | where amp is the maximum horizontal ground amplitude(zero-peak)  in nm of a Wood-Anderson simulation of the recorded seismogram, and dist is the hypocentral distance in km.So far, there are no calibration parameters (a, b, c, d) available for the Rwenzori region.We use the presets of the SEISAN software package: a = 1.0, b = 0.91, c = 0.00087, and d = 1.67, which were derived by Alsaker et al. (1991) (Ottem öller, personal communication, 2010).Figure 4 presents the magnitude of an event calculated at twelve different stations as a function of epicentral distance (circles).The average of all stations is M L = 2.3 with a standard deviation of ±0.3.Obviously, there is no systematic distance Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | 3.3., this is probably an artefact due to the absence of stations in that area, which decreases the sensitivity of the station network and enables the detection of only the stronger events.Currently we are operating a new network that covers the whole Rwenzori region including the western and south-western parts situated in the Democratic Republic of Congo.The analysis of the data will answer the question if 572 Discussion Paper | Discussion Paper | Discussion Paper |

Fig. 7
Fig.7we now observe a highly complex and individual picture at the different vertical sections: Profile B: This section is crossing the northern part of the Rwenzoris.Again, there is a distinct difference in the hypocentral depth distribution: west of the mountains (towards the rift valley, 5-25 km) focal depths range from 10 to 20 km.On the eastern side (40-55 km), where the Rwenzori Mountains are connected to the rift shoulder, seismicity extends from the surface down to 30 km depth.In between we observe only weak seismic activity which is typical for the Rwenzori block.Moho depths show the same characteristics.In the rift valley they range from 21 km to 27 km (BURO, SEMP, ITOJ, KARU) and beneath the rift shoulder from 29 km to 33 km (KMTW, KAGO, MIRA).A remarkable feature is the detection of mantle earthquakes in this area.A small group of seven earthquakes is located at depths 574 Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Figure 11 presents the cumulative magnitude-frequency relations calculated for two different locations.Area 1 is situated in a region north of the Rwenzoris with high station density 576 Discussion Paper | Discussion Paper | Discussion Paper | and the magnitude-frequency relation (circles) suggests a magnitude of completeness of M C = 0.8.The second area lies southeast of the mountains within the Edward rift.
Fig. 12 gives an illustration of the procedure.The event was recorded 577 Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper |

Table 1 .
Bram (1975)del UVI-N74 ofBram (1975)used for earthquake location.S-velocities were calculated with a constant v P /v S ratio of 1.74 as derived byBram.