MIRAMARE – TRIESTE September 2009

Available at: http://publications.ictp.it IC/2009/074

United Nations Educational, Scientific and Cultural Organization and International Atomic Energy Agency

THE ABDUS SALAM INTERNATIONAL CENTRE FOR THEORETICAL PHYSICS

STRUCTURE OF THE CRUST BENEATH CAMEROON, WEST AFRICA, FROM THE JOINT INVERSION OF RAYLEIGH WAVE GROUP VELOCITIES AND RECEIVER FUNCTIONS

Alain-Pierre K. Tokam Department of Physics, University of Yaoundé 1, P.O. Box 812, Yaoundé, Cameroon,

Charles T. Tabod1 Department of Physics, University of Yaoundé 1, P.O. Box 812, Yaoundé, Cameroon and The Abdus Salam International Centre for Theoretical Physics, Trieste, Italy,

Andrew A. Nyblade, Jordi Julià Department of Geosciences, Pennsylvania State University, University Park, PA, USA,

Douglas A. Wiens Department of Earth and Planetary Sciences, University of Washington in St. Louis, St. Louis, USA

and

Michael E. Pasyanos Earth Science Division, Lawrence Livermore National Laboratory, Livermore, California, USA.

MIRAMARE – TRIESTE September 2009

1 Regular Associate of ICTP.

Abstract

The joint inversion of Rayleigh wave group velocities and receiver functions was carried out to investigate the crustal and uppermost mantle structures beneath Cameroon. This was achieved using data from 32 broadband seismic stations installed for 2 years across Cameroon. The Moho depth estimates reveal that the Precambrian crust is variable across the country and shows some significant differences compared to other similar geologic unitsin East and South Africa. These differences suggest that the setting of the Cameroon Volcanic Line (CVL) and the eastward extension of the Benue Trough have modified the crust of the Panafrican mobile belt in Cameroon by thinning beneath the Rift area and CVL. The velocity models obtained from the joint inversion show at most stations, a layer with shear wave velocities ≥ 4.0 km/s, indicating the presence of a mafic component in the lower crust, predominant beneath the Congo Craton. The lack of this layer at stations within the Panafrican mobile belt may partly explain the crustal thinning observed beneath the CVL and rift area. The significant presence of this layer beneath the Craton, results from the 2100 Ma magmatic events at the origin of the emplacement of swarms of mafic dykes in the region. The CVL stations are underlain by a crust of 35 km on average except near Mt-Cameroon where it is about 25 km.. The crustal thinning observed beneath Mt. Cameroon supported by the observed positive gravity anomalies here, suggests the presence of dense astenospheric material within the lithosphere. Shear wave velocities are found to be slower in the crust and uppermost mantle beneath the CVL than the nearby tectonic terrains, suggesting that the origin of the line may be an entirely mantle process through the edge-flow convection process.

1. Introduction The Cameroon Volcanic Line (CVL) is a major geologic feature that cuts across Cameroon from the south west to the north east. It is a unique volcanic lineament which has both an oceanic and a continental sector and consists of a chain of Tertiary to Recent, generally alkaline volcanoes stretching from the Atlantic island of Pagalu to the interior of the African continent (figure.1). The oceanic sector includes the islands of Bioko (formerly Fernando Po) and São Tomé and Príncipe while the continental sector includes the Etinde, Cameroon, Manengouba, Bamboutos, Oku and Mandara mountains as well as the Adamawa and Biu Plateaus. Three other major tectonic features are present in the region: the Benue Trough located west of the line, the Central African Shear Zone (CASZ) trending N70°E along the line and The Congo Craton located in southern Cameroon. The origin of the CVL is still a subject of debate between the plume and non-plume models (Deruelle et al., 2007; Ngako et al, 2006; Burke, 2001; Ritsema and Van Heijst, 2000; Ebinger and Sleep, 1998; Lee et al, 1994; Dorbath et al., 1986; Fairhead and Binks, 1991; King and Ritsema, 2000). The crustal structure beneath Cameroon has been investigated using active (Stuart et al, 1985) or passive (Dorbath et al., 1986; Tabod, 1991; Tabod et al, 1992; Plomerova et al, 1993) source seismic data in different regions of Cameroon. These studies reveal a crust of about 33 km or more at the south-western part of the continental CVL (Tabod, 1991), a crust of about 33 km beneath the Adamawa Plateau and a thin crust of about 23 km beneath the Garoua Rift in the north (Stuart et al, 1985). However, a map showing the variations of crustal thickness within Cameroon was obtained using gravity data (Poudjom et al., 1995). This map reveals a particular feature of a thin crust beneath the Garoua Rift associated to relative positive Bouguer gravity anomalies, while low Bouguer gravity anomalies beneath Adamawa Plateau is also related to a thin crust. Spectral analysis applied to the same dataset revealed rather a normal crust beneath the Plateau (Nnange et al., 2000). More recently, a tomographic study provided the first surface wave group velocity maps of Africa (Pasyanos, 2005) giving way later to its lithospheric structural maps (Pasyanos and Nyblade, 2007). These studies delineate major geologic features within Cameroon but do not fill in the gap in the knowledge of its crustal structure due to the scale of the studies. One approach that has been used successfully to image crustal and upper mantle structure in East African Rift region and Parana basin (southern Brazil) utilizes a joint inversion of Rayleigh wave group velocities and P wave receiver functions (Julià et al., 2005, 2008; Dugda et al., 2007). From the 1-D Shear wave velocity models obtained, estimates for the composition and structure of the crust and upper mantle have been derived, leading to a better knowledge of the subsurface lithology in these areas. In this study, we investigate the crustal structure beneath the CVL and the adjacent geological provinces using 1-D shear wave velocity models obtained from the joint inversion of Rayleigh wave

Group velocities and P- Receiver functions for 32 broadband stations located in Cameroon (figure 2). After briefly reviewing the geological framework of Cameroon, we describe the data and technique used for the joint inversion, discuss crustal structure revealed by the velocity models. We also examine the crustal heterogeneities observed beneath the CVL in relation to the adjacent geologic structures.

2. Tectonic and geologic settings 2.1. The Cameroon volcanic Line Much of the continental sector of the CVL is underlain by Pan-African basement rocks while oceanic crust underlies the oceanic parts. The continental sector is represented by rocks of Precambrian, Cretaceous, Tertiary and Quaternary ages. On this sector of the Cameroon line, the basement, which represents the Central African mobile zone, consists of Precambrian rocks that were remobilized by the Pan-African episode (600-500Ma). These rocks are mainly schists and gneisses which have been intruded by granites and diorites. Cretaceous sediments, mostly sandstones and small amounts of limestone and shales, are found in the coastal areas. Along the whole line, several occurrences of mantle-derived (ultramafic) xenoliths have been found in basaltic lavas (figure 1; Deruelle et al., 2001, 2007; Princivalle et al., 2000). The presence of xenoliths gives some information about the composition of the lower crust and indicates that the formation of the CVL could be the result of a mantle process. In fact, evidence of metasomatism in the mantle source of the CVL parental magmas has been deduced partially from mantle xenoliths (Deruelle et al., 2007). The continental sector volcanic rocks range from basalts to trachytes. Mt. Manengouba is made up of basalt, trachyte and rhyolite lavas. The lavas of Mt. Cameroon, the largest of the continental volcanoes, are mainly of alkaline basic type (Hedberg, 1968) while Mt. Etinde, an older volcanoe is covered by the nephelinite lavas (Nkoumbou et al, 1995; Hedberg, 1968). The Mt. Bamboutos lavas are mainly alkali basalts and trachytes while Mt. Oku is composed of transitional basalt, quartz trachyte and rhyolite (Fitton and Dunlop, 1985). The Mandara Mountains along the northern Cameroon-Nigeria border represent trachyte and rhyolite plugs with the presence of alkali basalt flows. Some of the oldest volcanic rocks (aged up to 34Ma) of the CVL are found in the Mandara Mountains (Fitton and Dunlop, 1985).

2.2. The Benue Trough The Benue trough is a NE-SW trending basin that extends from the Niger delta basin (Gulf of Guinea) to Lake Chad. Its origin is thought to be linked to the opening of the South Atlantic Ocean in the cretaceous times (Guiraud et Maurin, 1992). The Yola-Garoua or Garoua Rift and Mamfe basin are known as its eastward extensions in Cameroon. The Y-shape of the (CVL) is similar to the shape of the adjacent Benue trough (Fitton, 1987) and suggests an interaction during the formation of both

structures. In fact, the alkali basalts of the Benue Trough are geochemically and isotopically similar to those of the CVL (Coulon et al., 1996). The orientation of the trough seems to be controlled by northeast trending dextral shear zones of late Pan-African age (Guiraud and Maurin, 1992).

2.3. The Central African Shear Zone The Central African Shear Zone (CASZ) is a major tectonic feature of Africa extending from the Darfur in Sudan across central Africa to the Adamawa Plateau region (Dorbath et al., 1986). In this region, the CASZ runs south-westwards, constituting the Foumban Shear Zone which is considered to be a continuation of the Pernambuco lineaments in Brazil (Burke, 1971; Browne and Fairhead, 1983). This shear zone is masked in south western Cameroon by the widespread Tertiary to Recent volcanic cover.

2.4. The Congo craton The Congo craton is an ancient Precambrian structure that occupies a large part of Central Africa. Its northern portion (in Cameroon) is referred to as the Ntem Complex (Vicat et al., 1996). The Congo Craton in Cameroon consists predominantly of Archean rocks with some reworked and resedimented material formed in the Paleoproterozoic (Tchameni et al., 2001). The Archean period began before 3.14 Ga during which Paleoarchaean protocrusts were formed. Evidence of the existence of these protocrusts have been pointed out in the Ntem complex and they represent the protholiths of the Archean greenstone belts of the Congo Craton (Nsifa, 2006; Tchameni et al, 2001). Large segments of the preserved Archaean continental crust are made up of an association of mafic to ultramafic extrusive rocks and metasediments that form the classic greenstone belts surrounded by magmatic rocks of the tonalite-trondhjemite-granodiorite (TTG) suite (Tchameni et al., 2000). The PaleoProterozoic evolution of the Ntem Complex is equivalent to the Eburnean orogenic cycle, characterized by intrusion of mafic doleritic dykes; this cycle ended with a thermal or hydrothermal event at around 1800 Ma (Nsifa, 2006; Tchameni et al., 2001). The boundary between the Panafrican belt and the Congo Craton occupies part of southern Cameroon and progresses towards the North of Central African Republic (Boukeke, 1994). In this region the Congo Craton, underthrusts the Panafrican block, along an intracrustal discontinuity. At depth, there are E–W striking dense bodies corresponding to lower crustal and/or upper mantle rocks setting by tectonic compression (Boukeke, 1994).

3. Data and methodology The data used in this study were recorded between January 2005 and February 2007 by the Cameroon Broadband Seismic Experiment, which consisted of 32 portable broadband seismometers installed across the country (Figure 2). Each station was equipped with a broadband seismometer (Guralp 3T

and Streckheisen STS-2), a 24-bit Reftek digitizer and a GPS clock. Data were recorded continuously at 40 sps. Eight stations were installed in January 2005 and operated for two years; the remaining 24 stations were operated only for the second year of the experiment. The global station spacing was of about 50 to 150 km. The use of Rayleigh wave group velocities and Receiver Functions depends on the fact that both datasets are complementary and consistent (Julià et al., 2000). Receiver functions are time series that represent the radial impulse response of the shallow structure of the Earth in the vicinity of the seismic station (Langston, 1979). They are a powerful tool to resolve velocity contrasts across discontinuities and the S-P travel times between the discontinuity and the free surface (Owens et al., 1984; Ammon et al., 1990). Rayleigh wave group velocities can be used to constrain the average shear wave velocity with depth and considerably reduce the intrinsic non-uniqueness during receiver inversion (Julià et al., 2000).

3.1. Rayleigh wave group velocities The measurement of fundamental mode Rayleigh-wave Group velocities has been performed on more than 100 events of magnitude 5 and above originated at epicentral distance range below 40° and recorded by the Cameroon array. The single station method was used for this purpose and provided several event-to-station dispersion data. The measurements were later included in the database for Africa by Pasyanos (2005) to carry out a tomographic inversion that increased the resolution of group velocity estimates within Cameroon. Local dispersion data were then acquired, at period range 7 to 100 s, for the tomographic cells including each of the stations under study. Figure 3 shows the group velocity curves for some stations lying along the CVL. At first sight, the group velocities show an overall decrease from the South (CM09) to the North (CM32) (figure 3a) and an increase from West (CM18) to the East (CM24) (figure 3b) for the period range 20 to 100 s. The dispersion curve for station CM06 located on the Congo Craton is plotted below for comparison with those along the CVL. It can be seen that the group velocities are generally higher under the Craton than the CVL. In fact, this is an indication that the thickness of the crust along the continental part of the CVL varies from north to south and from west to east and exhibits a relatively slow lower crust and upper mantle compared to the Congo Craton. This feature has been revealed in the regions of Lake Nyos and Mt. Cameroon belonging to the CVL (Tabod et al., 1992) and along the whole volcanic line (Euler et al, 2007). The variation of group velocities for the period range 7-20 s is more scattered and displays the complexities of the shallow structures of the upper-crust.

3.2. Receiver Functions Receiver Functions are computed using data from 69 teleseismic events that occurred between 30° and 95° epicentral distance with magnitude greater or equal to 5.5. They are computed at two overlapping

frequency bands corresponding to Gaussian widths of a = 1.0 and a = 2.5 (corner frequencies of 0.5 Hz and 1.2 Hz respectively) since they contain complementary information on the receiver structure under the station (Julià, 2007). To compute the receiver functions, the selected waveforms were decimated to 10 s.p.s., windowed between 10 s before and 100 s after the leading P arrival, de-trended, tapered, and high-pass filtered above 50 s to remove low-frequency, instrumental noise. Radial and transverse receiver functions were then obtained from the filtered traces by rotating the original horizontal components around the corresponding vertical component into the great-circle path, and applying the iterative, time domain deconvolution procedure of Ligorría and Ammon (1999) to the rotated traces, with 200 iterations. The percentage of recovery of the original radial waveform was evaluated from the RMS misfit between the original radial waveform and the convolution of the radial receiver function with the original vertical component, and the events that were recovered to less than 85% were rejected. The remaining waveforms were visually inspected for coherence and similarity, and were later stacked and clustered by ray parameter and back azimuth. In addition, at least 3 waveforms were required to perform the stacks with the exception of the station CM27 at which only two waveforms were successfully recovered. The transverse receiver functions (not shown here, see Tokam, 2010) are computed to check the degree of lateral heterogeneity and isotropy of the propagating medium (Cassidy, 1992). Theoretically, we expect small or zero amplitudes for transverse amplitudes to conclude on a laterally homogeneous and isotropic media beneath the CVL stations. In general, compared to radial receiver functions, the transverse waveforms are of small amplitudes apart from the waveforms of stations CM09 and CM15 which are comparable. Therefore, both stations are probably close to the regions of exceptionally high heterogeneities. The deconvolved waveforms are in general a succession of peaks and troughs that can be regarded as interactions of the incoming P-wavefront with subsurface discontinuities. The interactions generally result into P to S converted (Ps) phases. The receiver functions also include effects of reverberation between the free surface and discontinuities resulting from multiple conversions of P to S in the propagation path. The Ps phase generated by the Moho is generally apparent in all waveforms as peaks at about 5 s while some other apparent phases (peaks or troughs) between 1 and 4 s could be interpreted as the interaction of the P-wave front with sediments or intra-crustal discontinuities.

3.3. Joint inversion The joint inversion is performed using the method developed and updated in Julià et al. (2000, 2003). Recall that both receiver functions and dispersion curves are sensitive to the shear wave velocity of the lithosphere. The method is then based on a linearized inversion procedure that minimizes a weighted

combination of least squares norms for each data set, a model roughness norm, and a vector-difference norm between inverted and preset model parameters. The velocity models obtained are consequently a compromise between fitting the observations, model simplicity and a priori constraints. For the joint inversion, the two data sets must sample the same study area, which is achieved by combining receiver functions with tomographic dispersion velocities. To make the contribution of each data set to the joint least squares misfit comparable, a normalization of the data set is necessary, and this is done using the number of data points and variance for each of the data sets. An influence factor can then be used to control the trade-off between fitting the receiver functions and the group velocity curves. The initial model for the joint inversion assumed an isotropic medium with a 37.5 km thick crust and a linear shear wave velocity increase from 3.4 to about 4.0 km/s overlying a flattened PREM model (Dziewonski and Anderson, 1981). The poisson’s ration was set to 0.25 for the crust and crustal densities were deduced from P-wave velocities through the empirical relationship of Berteussen (1977). The initial model consisted of constant velocity layers that increase in thickness with depth. The thickness of the first and second layers are respectively 1 and 2 km while the thickness increases as 2.5 km between 3 and 60.5 km depth, 5km between 60.5 and 260.5 km, and 10km below a depth of 260.5 km. Shallow layers thinner than 1 km could have been resolved by the joint inversion if a priori information such as geotechnical data were available in the vicinity of our stations. In fact, the existence of geotechnical data allows including in the starting model, shallow layers as thin as 0.25 km, but required that the joint inversion closely matches the a priori geotechnical values (Julià et al., 2008). The inverted models generally show a good fit with group velocity curves and receiver functions with the exception of stations CM09, CM15 and CM27. In general, the group velocities below 10 s period are sensitive to shallow crustal structures and sedimentary basins. For period up to 40 – 50 s, they are sensitive to the whole crust and the Moho and for period 50-100s, they sample the uppermost mantle (Pasyanos, 2005). Consequently, the depth of investigation was fixed to 200 km to account for the sensitivity of our group velocities dispersion curves to the above depth range. The models were then constrained to the PREM velocity model by Dziewonski and Anderson (1981) for depth below 200 km while inverting at the same time the velocity structure above the 200 km depth. The Poisson’s ratio of 0.25 was assumed for the crust while the Poisson’s ratio for to the PREM model was used for the mantle.

3.4. Azimuthal dependence and Model uncertainties The approach of Julià et al. (2008) was applied to check for laterally varying structures around each station. Recall that the Receiver functions were stacked clustered according to backazimuth and ray- parameter and were computed for each event at overlapping frequencies corresponding to Gaussian width of a= 1.0 and a= 2.5. Thus, there are actually two receiver function averages for each cluster. Each receiver function cluster was then jointly inverted with the corresponding dispersion velocities to

obtain a shear velocity model for that specific group. An average S-velocity model is obtained by inverting all the receiver functions groups with the same dispersion velocities and compared to the single group velocity models. The procedure, applied on stations that have at least four clusters of receiver functions, is illustrated on figure 4 for the station CM07. In general, there should not be significant variation between the average and single cluster models in the case of a stratified isotropic media since the receiver functions bear the imprint of the latter stages of the propagation process in the vicinity of the receiver through, for example, P to SV conversions (Kennett, 2002). Any variations between both models should result from the interaction of the P-waves with some small scale lateral heterogeneities. Then, to estimate the uncertainties in the S- velocity models, the approach to compute the standard deviation between the single cluster and the average models was adopted. From this analysis, estimates of uncertainties in shear velocity for selected stations are generally less than 0.2 km/s for the crust. Given that for most stations, we got less than four clusters of receiver functions, we place the overall uncertainty in shear velocity at no more than 0.2 km/s for any given crustal layer in the model. This implies an uncertainty of 2 to 3 km in Moho depth for most stations where a rapid increase of velocity can be observed at crust mantle- boundary, and no more than 5 km where a smoothly varying shear velocity is found, characteristics of a gradational Moho.

4. Results Two stations (CM08 and CM14) among the 32 stations constituting the Cameroon Array did not provide enough waveforms to compute valuable receiver functions and so were not used in the joint inversion. Results from the joint inversion for the remaining stations are shown on figures 5 and 6 respectively. Figure 6 shows the shear wave velocity models clustered by tectonic terrains and Table 1 provides a summary of Moho estimates from this study. The criteria used to determine the crustal thickness is based on the following two assumptions. Firstly, from their difference in composition, there should be a rapid velocity increase between the crust and the Mantle. Secondly, Shear wave velocity above 4.3 km/s indicates the presence of lithologies with mantle compositions (Christensen, 1996; Christensen and Money, 1995). In fact using poisson’s ratio of 0.25 for the crust and the experimental P-wave velocities for various tectonic terrains, it is found that Shear wave velocity for typical lower crust lithologies is less than 4.3 km/s. Therefore, for the stations where rapid increase of velocity is observed, the Moho is placed at the corresponding depth with the condition to have shear velocity at least equal to 4.3 km/s, while for the gradational Moho, it is placed at depth where Shear wave velocity exceeds 4.3 km/s. The lower crustal velocity structure can be viewed as the part of the velocity model where the change in velocity becomes gradual and almost uniform. This gradual change is observed at most stations at depth of 15 - 20 km where velocity increases from 3.6-3.8 km/s to about

3.8-4.0 km/s. Therefore, the portion of the crust between 20 km depth and the Moho is here considered as the lower crust. Using the above considerations, we found good agreement between our Moho estimates and those reported by other studies at various areas in Cameroon using either seismic or gravity data (see table 2 for comparison). A detailed description of shear wave velocity models obtained is commented below for stations located at each geological terrain as described in the section 2 and displayed on figure 6.

4.1. CVL shear wave velocity models CVL stations consist of stations located at the Southern CVL and Adamawa plateau terrains. Using the above criteria, The Moho is found on average at depth of 35.5 km beneath the Adamawa uplift (figure 6). Beneath the southern CVL terrain, the crustal thickness estimates are on average about 35.5 km in high lands, 28 and 30.5 km respectively in the Kumba graben and Mamfe basin. Beneath Mt. Cameroon (CM09), the shear wave velocity model shows two depths with shear wave velocity above 4.3 km/s including a zone of velocity less than 4.3 km/s in between. The attempt to estimate the Moho depth here will take into account gravity data. The average crustal shear wave velocity of 3.7 km/s is found at most CVL stations except in the basin or graben areas where it is about 3.6 km/s. The models show within the top 15 km of the upper crust, the presence of a fast layer intruded into slower ones, giving the appearance of low velocity zone in the mid-crust. This anomalous feature is remarkable for most stations of the CVL with shear wave velocities between 3.6-3.8 km/s. Another feature is the presence of a thin low velocity layer (Vs < 3 km/s) observed at stations CM18 and CM21. The CVL lower crust as defined above has an average shear wave velocity of about 3.9 km/s and a thin higher velocity layer (Vs ≥ 4.0 km/s) component in its lowermost part. The thickness of this high velocity layer is 2.5 km at most stations while it is thicker at few stations within the line (CM19, CM23 and CM27). This layer is absent at stations CM24, CM26 of the Adamawa plateau. The crust-mantle transition is investigated by computing the average shear wave velocity for layers in the zone between the Moho to depths of about 60 km. This part of the mantle referred to here as uppermost mantle, could help at least to speculate on the mantle contribution to the origin of the CVL. The average shear wave velocity found here varies from 4.4 km/s beneath Adamawa Plateau and grabens to 4.5 km/s beneath the high lands CVL (table 1).

4.2.2. Panafrican terrain velocity models Stations CM03, CM10, CM12, and CM17 are located on the part of the Proterozoic Panafrican terrain not affected by the CVL. The crustal thickness here varies from 35.5 to 43 km with an average crustal shear wave velocity of about 3.8 km/s for the region. The Moho is therefore found deeper here than beneath the CVL. The upper crustal shear wave velocity displays the similar feature of the presence of

a fast layer within the top 15 km of the crust. Meanwhile, the lower crustal shear velocity is in range 3.9 - 4.0 km/s for the region. High velocity layer is also observed here with the thickness of about 5 – 7.5 km in the area. The uppermost mantle here is characterized by shear wave velocity of 4.5 km/s similar to the uppermost mantle velocity of the high lands part of the southern CVL terrain.

4.2.3. Congo craton terrain velocity models The Congo craton stations consist of stations CM02, CM04, CM06, CM03 and CM06. It is characterized by a crustal thickness of 43 – 48 km with an average crustal shear wave velocity estimated at 3.9 km/s. However, the feature of a fast layer in the upper crust remains visible in the velocity models for all stations. Below this fast layer, there is a gradual increase of velocity with depth up to Moho. The lower crustal shear velocity for the craton stations is in range 4.0 -4.2 km/s which is faster compared to the entire Panafrican lower crust. Meanwhile, the high velocity layer is found beneath the Congo craton with thickness varies between 18 km and 27.5 km across the region. This layer is the dominant component of the lower crust across the region. The Congo craton crust is underlain by a faster uppermost mantle with average shear wave velocity of 4.6 km/s.

4.2.4. Coastal terrain velocity models The Moho at stations CM01 and CM05 is found at depth of 28 km, showing that the crust is thinner at coastal terrain. This thin crust is associated to an average slower shear wave velocity of 3.7 km/s. The fast layer in the upper crust remains pervasive in the velocity models for this region. The average lower crustal shear wave velocity of 4.1 km/s is found beneath both stations. The lower crust here is entirely made of this higher velocity layer which has a thickness here of about 10 km. The fast lower crust is underlain by an uppermost mantle with shear wave velocity of about 4.4 km/s.

4.2.5. Rift area shear wave velocity models This area is composed of stations located in and around the Garoua rift, north to the Adamawa uplift. The crustal thickness is variable across the region. The Moho is found at depth of 25.5 km beneath station CM29, 30.5 km beneath CM28 located south to the rift and above 30 km beneath stations CM31 and CM32 located north to the rift. The average crustal shear wave velocity is about 3.4 – 3.5 km/s across the region. The upper crustal structure reveals the presence of a shallow thin layer with low shear wave velocity (Vs < 3km/s) at most stations. The lower crustal velocity is variable with an average shear wave velocity of 3.7 – 3.9 km/s across the area. The crust is thinner and slower beneath stations CM29 within the rift. The average uppermost mantle shear wave velocities is about 4.3 – 4.4 km/s across the region which is slower compared to the uppermost mantle velocity for terrains south of the region.

5. Discussion The joint inversion was performed using receiver functions clustered by ray parameter and back azimuth (Julia, 2008). In general our inverted models are averages of the crustal structure within at least two back azimuths close to 80 and 240 degrees, except at stations CM02 and CM09. These two back azimuths are almost opposite providing a good constrain to our Moho depth estimates. The velocity-depth profile for most stations shows the presence of higher shear wave velocity layer (Vs ≥ 4.0 km/s) in the lower crust. This layer is consistent with the presence of mafic rocks characteristic of the lower crust in Precambrian terrains globally as pointed out by Rudnick and Gao (2003). The Precambrian crust in Cameroon is represented mainly by Congo craton in the south and Panafrican mobile belt in the north. The average thickness of this mafic layer is about 22.8 km for the Congo craton and less than 10 km for the entire Panafrican crust. The presence of a fast layer in the upper crust is well marked on high frequency band receiver functions as phases (peaks) between 1 and 4s. Given that group velocities acquired from the revised model of Pasyanos and Nyblade (2007) are better resolved within Cameroon, this layer cannot be neglected and may be regarded as an anomalous upper crustal feature. The signature of known sediment coverage within Cameroon also appears on velocity profiles as layers with Vs < 3 km/s and thickness range 1 – 3 km. The sediment thickness in the Garoua rift (CM29) is about 3 km, comparable to the thickness of 4 km found using gravity data (Kamguia et al., 2005), and showing that the joint inversion technique is able to resolve small scale structures. A strong variability in Panafrican mobile belt crustal structure is observed in the velocity models while crustal structure beneath the Congo craton is more stable. In the following subsections, we discuss on the variabile structure of Panafrican mobile belt, compare the Precambrian crust in Cameroon to that for similar aged terrains in East and South Africa and determinetheir contribution to the origin of the CVL.

5.1. Panafrican mobile belt structure across Cameroon The Panafrican mobile Belt contains the Panafrican, CVL and Rift terrains (table 1). The mafic component of the lower crust beneath the Panafrican terrain has a thickness of,about5 - 7.5 km which agrees with the range of 5 – 10 km found beneath many proterozoic terrains (Durrheim and Mooney, 1991). Thus, the quasi-absence of this component beneath the CVL and the rift area suggests both geologic structures being particular features of the Panafrican mobile belt in the sense that their mafic layers seem to be removed, given that a layer of 2.5 km thick is not strongly resolved on models. As a matter of fact, these observations suggest that the presence of the CVL has modified the crustal structure of the Panafrican mobile belt. An early stage of continental rifting has been suggested by Poudjom et al. (1997) to explain the lithospheric thinning observed beneath the Adamawa uplift and could also explain the removal of the mafic layer beneath the uplift.

The rift area is known as the eastward extension of the Benue Trough., which has been accompanied by crustal thinning (Kamguia et al., 2005; Stuart et al., 1985). The subsequent removal of the lower crustal mafic layer in the area occurred through the delamination process due to the successive phases of magmatic activity at the origin of the opening and the infilling of the Trough. The main extensional tectonic regime affected the Trough in the Early Cretaceous time (Baudin, 1991), followed by the emplacement of basaltic magma in the Tertiary with the greatest concentration in the North East (Wilson and Guiraud, 1992). As a result, many hypovolcanic dykes are found within the Garoua basin with basaltic lavas similar to those of the whole CVL (Ngounouno et al., 1997). The similar presence of tertiary basaltic lava coverage is found within the Mamfe basin (Dumort et al, 1968) another arm of the Benue Trough in Cameroon, where crustal thinning has been suggested in a previous study by Kamguia et al. (2008).

5.2. Interpretation of Precambrian crustal structure in Cameroon with respect to Precambrian crustal structure in East and South Africa. Table 3 below compiles results of Moho estimates by Julia et al. (2005) and Kgaswane et al. (in press) using the joint inversion technique under Precambrian terrains from two regions in Africa –East Africa and South Africa. The East African region consists of the Archean Tanzanian craton, bordered to the east by the Mozambique belt and to the southwest by the Kibaran belt. The South African region on the other hand consists of the Archean Kaapval Craton welded to the Archean Zimbabwe cratons by the Limpopo Paleoproterozoic mobile belt.

5.2.2. Congo Craton Versus East and South African cratons The Tanzanian craton is underlain by a crustal thickness ranging between 38 – 42 km while the crustal thickness beneath Kaapval and Zimbabwe cratons is between 36 and 40 km. Both craton crusts are thinner than the crust beneath Congo craton (43 – 48 km). A quick inspection of the lower crustal structure indicates that the southern African shield has a predominant mafic composition of the lower crust except the southwest part of the Kaapval craton and the western part of the Zimbabwe cratons, where the lower crust is intermediate-to-felsic in composition. The mafic layer thickness is in general less than 14 km beneath these cratons, which is thinner than the average thickness of about 23 km found beneath the Congo craton. The mafic layer is estimated to be less than 10 km thick beneath the Tanzania craton. Geologic evidence of the mafic composition of the Congo craton lower crust could be found in the past magmatic activities recorded in the region. The last mafic event is dated at the period before 2100 Ma and marked by the rifting of the Archean Ntem crust (Nsifa, 2006; Vicat et al., 1996) and resulted in the emplacement of swarms of mafic doleritic dykes (Maurizot et al., 1986; Vicat et al., 1996; Tchameni et al., 2000; Nsifa, 2006). In addition, the presence within the Ntem complex of mafic

rocks such as gabbros (Maurizot et al., 1986) and the greenstone belts (Maurizot et al., 1986; Shang et al., 2004, Tchameni et al. 2001) indicate that a major part of the Archean protocrust is preserved. Accordingly, these observations could justify why the Congo craton crust is thicker compared to the other African craton crusts. However, the existence of suture zones between the Congo Craton and the Panafrican mobile belt has been suggested by several studies (e.g. Penaye et al., 1993; Tadjou et al., 2009), although, the position of the main suture zones and the character of the different tectonic units are still a matter of debate. In fact, continental collision like the transition between the Congo craton and the Panafrican mobile belt is generally accompanied by a crustal thickening while the crust squashes horizontally. In that hypothesis, the Congo craton might be underplated by the mobile belt, which is not the actual case. Therefore, the Congo craton might have not experienced crustal thickening, and the thicker crust observed here is typical for the region. Further estimates of the crustal thickness at other areas of the craton are needed to support this finding. The finding of 50 ± 10 km for the crustal thickness by Poujom et al. (1995) is not strongly resolved due to the lack of coverage of gravity data across the region.

5.2.2. Proterozoic mobile belts The crustal thickness beneath the Mozambique and Kibaran belts (East Africa) are about 36 – 40 km and 40 – 44 km respectively. Beneath the Limpopo belt, it is found at depth of about 38 – 44 km while the crustal thickness beneath the preserved Panafrican terrain in Cameroon is found in a similar range of about 35.5 – 43 km. Inspection of the lower crustal structure indicates the presence of a mafic component beneath all the mobile belts. The thickness of 5 – 7.5 km for this mafic layer beneath the Panafrican terrain in Cameroon is found in the range of typical Proterozoic terrains (Durrheim and Mooney, 1991). For these reasons, the typical mobile belt structure in Cameroon is comparable to that in East and South Africa and therefore can be considered normal regardless the suture zone producing thickened crust in the Limpopo belt (Kgaswane et al., in press). The thickness of the mafic component is between 11 and 14 km beneath the Limpopo Belt in the lower crust. Accordingly, the crust under the CVL and the rift area is thinner than expected. Typical CVL stations are characterized by the quasi- absence of the lower crustal mafic layer as mentioned above and an average thickness of 35.5 km. These observations are evidence that the settings of the CVL and the Benue Trough have modified the Panafrican mobile belt structure in Cameroon.

5.3. Implications to the origin of the CVL The Moho depth for the typical CVL stations is on average about 35 km, in good agreement with previous studies (Fairhead and Okereke, 1987; Tabod, 1991). However, the average shear wave velocity of the crust for typical CVL stations is in general slower than that for other tectonic provinces.

This has also been revealed by preliminary tomographic results using Rayleigh wave phase velocities (Euler et al., 2007). In fact, the CVL crust has in general a Precambrian lower crust (average Vs ≤ 3.9 km/s at typical CVL stations), implying that the average low velocities observed within the line are due either to a thermal effect or to the petrographic composition of the entire crust. Recent heat related activities are recorded along the line denoting that the line is still active: eruptions of Mt-Cameroon in 1999 and 2000; emission of toxic gas in crater Lakes Monoun (1984) and Nyos (1986); a permanent hot spring observed during the field work at Woulde near Tignere (CM27). In addition, the CVL has experienced Cenozoic volcanism and is underlain by Panafrican basement rocks. In the hypothesis that the removal of the mafic layer beneath the Panafrican lower crust did not affect the overall structure of the crust along the CVL, all recent magmatic activities in the region should have occurred from zones of crustal weakening and reached the surface through propagating cracks produced by the interactions between CVL and CASZ (Poudjom et al., 1997). In this hypothesis, the origin of the line probably derived from an entirely mantle process. The uppermost mantle shear wave velocity of 4.4 – 4.5 km/s is slower beneath the line compared to the velocity of 4.6 km/s beneath the Congo craton, suggesting the edge-flow convection mantle process (King and Ritsema, 2000) between the colder craton and the hot line. The edge-flow convection model has been found in a recent study by Reusch (2009) to remain the viable model to explain the linearity of the CVL, the timing of volcanism, and the similarities in the geochemical properties of the rocks. As a result, the remnant heat from Cenozoic to recent magmatic activities could partly explain the lower shear wave velocities observed along the line.

5.4. Correlation of Moho variation with Gravity anomalies To assess the reliability of our Moho estimates, the variation of the Bouguer gravity anomaly was compared to the variation of the crustal thickness along profile P1 crossing the southern CVL terrain and the Rift area (figures 2 and 7). For this purpose, we select as Moho depth beneath CM09, the first depth where Vs > 4.3 km/s. The Moho depth is therefore found to vary almost uniformly with the Bouguer anomaly (figure 8), depicting a negative correlation between the two. Deep Moho corresponds to low anomaly (down to -80 mgal) while the thin crust is concordant with the relatively positive anomaly. The little decrease in the Bouguer anomaly around Garoua (CM29) is due to the residual effect of sediments on the whole regional gravity. The thickness of about 3 km attributed to this sedimentary basin is therefore strongly consistent with depths of 3-4 km thick found by Stuart et al. (1985) from seismic refraction studies. It is worth noting that away from the region around Mt. Cameroon (CM09), the Bouguer anomaly is negative along the whole line as expected for continents. Theoretically, one should expect positive anomalies within the ocean and negative anomalies within the continent. Thus, the positive anomalies within the continent suggest the presence of dense mantle materials at shallower depth than expected.

Although the shear wave velocity model obtained at CM09 suffers from some insufficiencies (large transverse receiver functions, misfit with low-frequency receiver functions in the inverted model), the model could be used as a good approximation to investigate the Moho. In that case, with support from gravity data, the crust is thin in this region and the Moho is estimated at a depth of about 25.5 km. Another supporting argument is the finding by Dugda et al. (2008) of crustal thickness of 22 km near Mt-Cameroon using the H-k stacking method (Zhu and Kanamori, 2000). In addition, the occurrence of regular deep seismicity in the region (depths of about 60 km) has been interpreted to indicate the presence of magma conduit (Ateba and Nteppe, 1997). Consequently, the thin crust beneath the Mt. Cameroon region could reflect the asthenospheric convection implying the lithospheric thinning and/or the presence of asthenospheric material within the lithosphere.

6. Conclusion The joint inversion technique using Rayleigh wave group velocities and Receiver functions has been used to investigate the structure of the crust and the uppermost mantle beneath Cameroon. Using the technique by Julia et al. (2008), better constrained one dimensional shear wave velocity models for the crust and uppermost mantle beneath 30 stations have been derived. From the velocity models, the Achaean Congo Craton is found underlain by an average crustal thickness of about 45 km which is considered thicker than the other African Archaean Cratons. The difference in thickness between those cratons is found in their lower crust composition. Although suture zones could be invoked to explain an eventual thickening of the Congo craton crust, there exists few geologic arguments to support this hypothesis and thus sggests that the 2100 Ma magmatic events contribute to the crustal thickening. The Panafrican mobile belt in Cameroon is found to have been modified by the setting of the CVL and the eastward extension of the Benue Trough. In fact, the Panafrican terrain was found underlain by an African typical mobile belt crust and therefore were considered “normal” while the crust is thinner beneath the CVL and the rift area. The crust beneath the CVL stations is on average about 35 km and is in general slower compared to the nearest Panafrican and Congo craton terrains. The removal of its mafic component together with the uppermost mantle low velocity suggested an edge-flow convection as mantle process in the origin of the line. A gravity profile plotted across the CVL and the Rift area regions depicted a negative correlation between the Bouguer Anomaly and the variation of crustal thickness, therefore provided strong constraints to the thickness of about 25 km found near Mt Cameroon.

Acknowledgments Penn State University and Washington University in St Louis supported APKT in the USA where part of this work was carried out. The Abdus Salam International Centre for Theoretical Physics (ICTP), Trieste, Italy, supported CTT through the Associateship Scheme. Most of the figures were prepared using the GMT software. We would like to thank Charles Ammon, Yongcheol Park, Mulugeta Dugda for assistance with computer codes.

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Table 1. Summary of crustal structure by geological terrains shown in figure 6.

Average Average Average Crustal crustal Average Average Vs thickness of Uppermost Terrain Stations thickness thickness ± crustal Vs below 20 km layers with Vs mantle Vs (km) standard (km/s) depth (km) ≥ 4.0 km/s (km/s) deviation (km) (km) CM01 28 3.6 4.1 10 Coastal terrain 28 4.4 3.7 4.1 10 CM05 28 3.7 4.1 10 CM02 43 3.8 4.0 18 CM04 45.5 44.7 ± 1.5 4.6 3.9 3.9 4.1 4.1 25.5 22.8 Congo craton CM06 45.5 3.9 4.2 27.5 CM07 43 3.9 4.1 20 45.5 ± 2.5 4.6 3.9 4.1 CM11 48 3.9 4.1 22.5 CM03 43 3.8 3.9 5 CM10 38 3.8 4.0 7.5 Panafrican 38.6 ± 3.1 4.5 3.8 3.9 6.3 CM12 38 3.8 3.9 7.5 CM17 35.5 3.7 3.9 5

20 Mt Cameroon CM09 25.5 ? ------CM15 33 3.6 3.9 2.5 CM16 35.5 3.7 3.9 5 High lands CM19 35.5 35.5 ± 3.1 4.5 3.7 3.7 3.9 3.9 5 5 Southern CVL CM20 33 3.7 3.9 2.5 CM23 40.5 3.7 4.0 10 Kumba graben CM13 28 3.6 3.8 3.8 2.5 2.5 29.2 ± 1.8 4.4 3.6 Mamfe basin CM18 30.5 3.6 3.9 3.9 2.5 2.5 CM21 35.5 3.7 3.9 2.5 Adamawa CM22 35.5 3.7 3.9 2.5 Plateau CM24 35.5 3.7 3.8 - 35.5 ± 1.6 4.4 3.7 3.9 2 (northeastern CM25 38 3.7 3.9 2.5 CVL) CM26 33 3.6 3.8 - CM27 35.5 3.7 3.9 5.0 South of CM28 30.5 30.5 4.3 3.5 3.5 3.8 3.8 2.5 2.5 Garoua rift Garoua rift CM29 25.5 25.5 4.3 3.4 3.4 3.7 3.7 - - Rift area CM30 28 3.4 3.7 - North of CM31 30.5 30.5 ± 2.5 4.4 3.5 3.5 3.9 3.8 2.5 2 Garoua rift CM32 33 3.5 3.9 2.5

Table 2. Comparision of crustal thickness estimates from this study with crustal thickness estimates from previous studies.

Average crustal thickness (km) Type of data References This study Other studies used 23 Seismic Stuart et al., 1985 Garoua Rift 25.5 24 Gravity Kamguia et al., 2005 CVL in general - 30-34 Gravity Fairhead and Okereke (1987) CVL southern part 35.5 ~ 33 Seismic Tabod, 1991. 33 Seismic Stuart et al., 1985 Adamawa Plateau 35.5 33 Gravity Nnange et al., 2000 Congo craton 45.0 50 ± 10 Gravity Poudjom et al., 1995

Table 3. Comparison of the thickness of the Cameroon Precambrian crust to the thickness of other Precambrian crusts across Africa.

East Africa South Africa Cameroon Zimbabwe 36 – 40 km Achaean craton Congo Craton 38 – 42 km 43- 45.5 km Tanzania craton Kaapval craton ~ 36 km craton Mozambique 36 – 40 km Panafrican 35 – 43 km belt Limpopo Mobile belt 38 – 44 km Kibaran belt 40 – 44 km belt CVL 28 - 35.5 km Ubedian belt 38 – 42 km References Julia et al. (2005) Kgaswane et al. (in press) This study

Figure 1: The Cameroon Volcanic Line (CVL) is a N30°E lineament of Cenozoic volcanism that extends from the Gulf of Guinea to Lake Chad. From a structural point of view, the continental sector is a succession of horsts and grabens . Mantle-derived xenolith occurrences are indicated by empty circles.

Figure 2: Station locations for the Cameroon Seismic Array experiment.

4 4 CM09 CM18

CM13 CM20

CM15 CM23

3.5 CM16 3.5 CM25

CM19 CM27

CM25 CM26

CM27 3 3 Group velocity (km/s) Group velocity CM28 (km/s) velocity Group

CM29

CM30

2.5 CM32 2.5 0 20406080100 0 20406080100 Period (s) Period (s) (a) (b)

4 CM18

CM20

CM23

CM25 3.5 CM21

CM22

CM24

3 Group (km/s) velocity

2.5 020406080100 (c) Period (s)

Figure 3: Dispersion curves for some selected stations showing the variation of Group velocities across the network. a) profile P1 lying along the CVL. b) profile P2 lying along the Adamawa branch of CVL. The station CM06 belonging to the Congo craton is plotted for comparison.

Figure 4: a) Single Joint inversion predictions for station CM07. The top, middle and bottom panels in each column display receiver functions, group velocities and Shear wave velocity models respectively. Observations are shown in black and predictions in red. b) Shear wave velocity models obtained using the whole group of receiver functions. c) Superposition of single (red) and full joint inversion models to assess the uncertainties.

Figure 5: Joint inversion results for all stations used in this study. Each column show results for a single station. The top, middle and bottom panels in each column display receiver functions, group velocities and Shear wave velocity models, respectively.

Figure 6: Shear wave velocity profiles grouped by tectonic terrain. The dashed (4.3 km/s) lines indicate the crust mantle transition.

Figure 7: Correlation between the Bouguer anomalies along the profile P1 with the Moho depth found in this work. The Bouguer anomaliy profile is extracted from the simple Bouguer gravity anomaly map computed by Kamguia et al. (2008).