An integrated model for the deep structure of the Chyulu Hills volcanic field, Kenya

TECTONOPHYSICS ELSEVIER

Tectonophysics 278 (1997) 187-209

An integrated model for the deep structure of the Chyulu Hills volcanic field, Kenya O. Novak a~¢, , J.R.R. Ritter b,l p R. Altherr c, V. Garasic c, F. Volker d, C. Kluge b, T. Kaspar b, G.F. Byrne b, S.V. Sobolev b,2, K. Fuchs b a Dublin Institute for Advanced Studies, School of Cosmic Physics, 5 Merrion Square, Dublin 2, Ireland b Geophysikalisches lnstitut, Universitiit Karlsruhe, Hertzstrafle 16, D-76187 Karlsruhe, Germany ~'Mineralogisches Institut, Universitiit Heidelberg, lm Neuenheimer Feld 236, D-69120 Heidelberg, Germany d lnstitutfiir Geowissenschafien und Lithosphdrenforschung, Universitiit Giessen, Senckenbergstrafie 3, D-35390 Giessen, Germany

Accepted 29 April 1997

Abstract

The Chyulu Hills, a 1.4 Ma B.E to Holocene volcanic field located about 150 km to the east of the Kenya rift, is one of the few locations on Earth for which detailed geochemical (volcanic rocks), thermobarometric (xenoliths), seismological and gravity data are available. This paper combines these data to achieve an integrated seismic-petrological model for the deep structure of this volcanic field. Results of a wide-angle reflection and refraction experiment reveal an average crustal thickness of 40 km and a thickness of 20 km for the lower crust. Beneath the volcanic field, the crust thickens to about 44 km. In this region a low-velocity body (LVZ) is modelled which extends from about 30 4- 5 km depth to the Moho. The LVZ is characterised by an increased vp/vs-ratio ranging from 1.81 to 1.93 depending on the possible extents of this body. This is in contrast to the surrounding crust where a ratio of only about 1.76 is observed. In the same area, the results of a teleseismic tomography study show a P-wave low-velocity anomaly of - 3 % . The seismic data can be explained by either an anorthositic body directly above the Moho in the region of the Chyulu Hills or by the presence of partial melt. Directly beneath the Chyulu Hills, a P-wave velocity of 7.9 km/s is determined for the uppermost mantle; this velocity is 0.2-0.3 km/s lower than that of the surrounding mantle region. The teleseismic tomography model suggests a P-wave low-velocity anomaly of -2.5 to -3.5% in the uppermost mantle (<70 km depth). Widespread garnet-bearing pyroxenitic and lherzolitic mantle xenoliths are mostly well equilibrated and suggest an apparent lithospheric thickness of about 105 kin. Most garnet-free spinel harzburgitic xenoliths and some garnet pyroxenitic xenoliths were significantly heated before they were sampled and erupted by the host magmas. Heating events lasted for less than 210 ka as indicated by chemical diffusion profiles observed in orthopyroxene grains. It is suggested that heating was caused by stagnating magmas in the uppermost lithospheric mantle. At the same depth P-wave velocity perturbations of the tomographic model show a low-velocity zone directly underneath the youngest part (SE) of the volcanic field. At depths greater than about 70 km, this low-velocity zone is shifted towards the east, away from the volcanic field. Keywords: interdisciplinary study; seismology; petrology; lithosphere; Kenya; partial melt; petrophysics; volcanism

* Corresponding author. Present address: GeophysicalInstitute, Universityof Karlsruhe, Hertzstrasse 16, D-76187 Karlsruhe, Germany. E-mail: [email protected] i Present address: Institute of Geophysics, University of G6ttingen, Herzberger Landstrasse 180, D 37075 G6ttingen, Germany. 2 Present address: GeoForschungsZentrumPotsdam (GFZ), TelegrafenbergA 17, D-14473 Potsdam, Germany. 0040-1951/97/$17.00 © 1997 Elsevier Science B.V. All rights reserved. PlI S0040- 195 1 (97)00 104-2

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1. Introduction Continental rifting and related magmatism are caused by a favourable interplay of mantle convection, plume activity, and plate boundary forces (e.g. White and McKenzie, 1989). For any given dynamic setting, the styles of rifting and magmatic activity will depend on the thermomechanical state of the lithosphere. The Kenya rift system offers an excellent scenario for studying the relations between the thermal and structural state of the lithosphere and the volcano-tectonic evolution. As this rift system is located in the transition zone between the Tanzanian craton and the Pan-African Mozambique belt (Fig. 1) (Smith and Mosley, 1993; Shackleton, 1993), the rifting process affected lithospheric domains of different thermal and structural states. Furthermore, the Kenya rift is characterised by extensive volcanic activity which started in northern Kenya in the Oligocene (Morley et al., 1992, and references cited therein). Subsequently, volcanism propagated southwards and was mainly restricted to the area of the actual rift and its western flank. Since the Late Miocene, large basaltic fields started to form on the eastern flank of the rift and Quaternary volcanic activity is restricted to the rift proper and its eastern flank (Baker, 1987; Karson and Curtis, 1989; Class et al., 1994; Macdonald, 1994). The origin and evolution of the large volcanic fields on the eastern rift flank are still an enigma because of the high rate of magmatism occurring about 150 km to the east of the rift where the actual extension of the lithosphere takes place (Achauer et al., 1994). Bosworth (1987) explained this off-axis volcanism as due to significant lithospheric thinning above the intersections of low-angle detachment systems with the base of a regionally thinned lithosphere. Thermobarometric data on mantle xenoliths (Henjes-Kunst and Altherr, 1992) as well as seismic studies (Keller et al., 1994) indicate that in NW Kenya both the lithosphere and the crust have been extensively thinned. Geochemical and isotopic data obtained for the Quaternary volcanic fields of NW Kenya have been interpreted in terms of a mantle plume interacting with an already thinned lithosphere (70-80 km) (Class et al., 1994). Such thinning of the lithosphere is, however, not observed in the area of the Chyulu Hills, a 1.4 Ma B.E to

Holocene volcanic field located in southern Kenya about 150 km to the east of the rift and 40 km to the north of Mt. Kilimanjaro (Fig. i). Based on thermobarometric data on mantle xenoliths, Henjes-Kunst and Altherr (1992) suggested an apparent lithospheric thickness of about 105 km for this part of the eastern rift flank. Encouraged by the promising geothermobarometric results, a multi-disciplinary study was planned and conducted from 1992 to 1995. More xenolith samples were collected and investigated and seismological as well as gravity studies were performed in order to image the lithosphere and uppermost asthenosphere underneath the Chyulu Hills region. Technical details about the refraction seismic and teleseismic tomography experiments are reported in (Prodehl et al., 1997). The 420-km-long refraction seismic line F (Fig. 1) extending from Athi River (30 km southeast of Nairobi) to the Indian Ocean near Mombasa was designed to gain the best resolution in the area of the Chyulu Hills volcanic field. The modelling technique and the complete P-wave model are described in (Novak et al., 1997). In a passive tomography experiment teleseismic P-wave arrivals were recorded at 31 stations (Fig. 2) in order to obtain structural information down to a depth of 140 km. Following Aki et al. (1977), the delay times were inverted into relative velocity perturbations within a block model (Ritter et al., 1995; Ritter and Kaspar, 1997). In this contribution, we try to combine the seismological, gravity and petrological data into an integrated model of the deep structure under the Chyulu Hills that may help to better understand the origin of volcanic activity in this part of the East African rift system.

2. Geological setting The Chyulu Hills volcanic field consists of several hundred volcanic cones and lava flows resting on a basement peneplain with an average elevation of about 1000 m. Dominant basement lithologies are quartz and feldspar-rich gneisses with minor garnet-beating amphibolites and marbles (Saggerson, 1963; Omenge and Okelo, 1992; Shackleton, 1993). The NW-SE-striking chain of the Chyulu Hills has a maximum altitude of 2187 m and is located directly on a major shear zone within the Pan-African

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M o z a m b i q u e belt (Smith and Mosley, 1993). This shear zone can be regarded as a pre-existing structure that could have facilitated m a g m a ascent to the surface. Based on various geomorphological criteria, the volcanic deposits have been subdivided

into five mapping units (groups 1-5) ranging in age from about 1.4 M a B.P. to H o l o c e n e (Saggerson, 1963; Haug and Strecker, 1995). The oldest four groups seem to be confined to the northwestern part o f the volcanic field near Sultan H a m u d and E m a l i

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(Fig. 2). Here, volcanic activity started with the eruption of highly silica-undersaturated lavas (foidites according to the lUGS nomenclature as given by Le Bas et al., 1986) (groups 1 and 2). Volcanic rocks of this initial phase are anomalously rich in TiO2 (up to 8.5 wt.%) and may contain abundant perovskite (Haug and Strecker, 1995). Basanitic and alkali basaltic pyroclastic deposits of groups 3 and 4 contain abundant mantle and lower crustal xenoliths (Goles, 1975; Henjes-Kunst and Altherr, 1992) indicating rapid ascent of the magmas from the mantle to the Earth's surface. The southeastern part of the Chyulu Hills forms a steep ridge of closely spaced volcanic edifices which is about 50 km long and was built up between Late Pleistocene and Recent time (group 5). The volcano Shaitani at the southeastern end of the main ridge (Fig. 2) was active in the last century (Saggerson, 1963; Nyamweru, 1980; Omenge and Okelo, 1992). The group-5 volcanic rocks are fractionated alkali basalts that lack mantle and lower crustal xenoliths but contain abundant crustal-derived xenocrysts like plagioclase and minor quartz. These characteristics suggest magma stagnation within the lower crust accompanied by fractionation and assimilation of lower crustal material. 3. Seismic structure of the lower crust

Along the refraction seismic line F (Figs. I and 2) the Mozambique belt crust has an average thickness of 40 km and thickens locally to 44 km underneath the Chyulu Hills volcanic field (Novak et al., 1997, fig. 8). The lower crust is remarkably thick and reaches a maximum thickness of 24 km below the Chyulu Hills, with P-wave velocities (Vp) increasing from 6.8 km/s at the top to 7.00-7.15 km/s at the base. This P-wave model (Novak et al., 1997), in the following called the 'original' model, is exclusively based on the refraction seismic P-wave dataset. It contains only positive vp-gradients. Using the results of the teleseismic tomography and

191

the refraction seismic S-wave modelling as further a priori constraints, the 'original' model was modified. The updated P-wave model will be referred to as the 'revised' model. In contrast to the 'original' refraction seismic P-wave model, the teleseismic tomography model of Ritter and Kaspar (1997) reveals a low-velocity anomaly in the lower crust underneath the Chyulu Hills (Fig. 3a). The calculated P-wave velocity perturbations ranging from - 1 . 6 % to +1.7% are the deviations from the starting model (table 1b in Ritter and Kaspar, 1997) taken from Novak et al. (1997, fig. 8) with an initial Vp of 6.9 km/s. The observed NW-SE-striking low-velocity zone in the tomography model correlates with the southern part of the Chyulu Hills, an area characterized by Holocene volcanic activity. The southeastern part of the model with positive velocity deviations is interpreted as undisturbed Mozambique belt lower crust (Ritter and Kaspar, 1997). Therefore, the total contrast between the undisturbed region and the P-wave low-velocity anomaly beneath the volcanic field is as much as 3%. The layer thickness of the teleseismic tomography model is 20 km and the velocity perturbation values are averaged along the whole vertical ray path. Thus, smaller volumes of even lower seismic velocity could also be present but would be smeared along the total layer thickness. Tests with various parameterizations of the tomography model and careful analyses of the full resolution matrix confirmed the existence of the low-velocity anomaly (Ritter and Kaspar, 1997). Likewise the interpretation of the refraction seismic S-wave data also requires a velocity decrease at the bottom of the lower crust beneath the Chyulu Hills. The S-wave starting model contains the layer boundaries derived from the 'original' P-wave interpretation. Further the P-wave velocities devided by 1.73 were used and adjusted until a final S-wave model providing travel time agreement with the observed S-wave data was obtained. At the top of Fig. 4a,b and Fig. 5a,b the P- and S-wave record

Fig. 2. Map displaying the distribution of volcanic rocks in the area of the Chyulu Hills and Kilimanjaro. Symbols show the locations of the xenolith localities (rhombs), the seismological network (triangles) and a part of the refraction seismic line F (dots and stars). The Middle Miocene Yatta phonolite flow is shown as a geographical landmark only, since it is not directly related to the Chyulu Hills volcanism. The inset shows Kenya with the rift valley.

O. Novak et al./Tectonophysics 278 (1997) 187-209

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Fig. 3. (a) Lower crustal layer 3 (23-43 km depth) of"the teleseismic tomographymodel with velocity perturbations in percent. Areas covered by volcanicrocks are outlined in order to facilitate comparison with Fig. 2. (b) Uppermostmantle layer 4 (43-66.5 km depth) of the teleseismic tomographymodel. sections of shotpoints ATH and CHS are shown (for location see Figs. 1 and 2). The P-wave record sections (Fig. 4a and Fig. 5a) are plotted with a reduction velocity of 6 km/s, the S-wave record sections (Fig. 4b and Fig. 5b) with a reduction velocity of 6 km/s: 1.73 = 3.46 krrds based on a Vp/vs-ratio of 1.73. Focusing on the lower crust only, the ray-traced phase Pi3 (Si3) travelling through the lower crust and the ray-traced phase PMP (SMS) reflected from the Moho (bottom of Fig. 4a,b and Fig. 5a,b) are plotted in the record sections. These phases result already from the 'revised' P-wave model and the final S-wave model which will be discussed below. In order to obtain a qualitative first estimate of deviations from the Vp/vs-ratio of 1.73, the phases Pi3 (Si3) and PMP (SMS) are plotted in the S-wave (P-wave) record section by multiplying (dividing) the travel times by 1.73. On the P- and S-wave record sections of shotpoint ATH (Fig. 4) phases Pi3 and Si3 a r e almost coincident and therefore confirm the assumption of

a Vp/vs-ratio of about 1.73 along their ray paths. In contrast, phase SMS, which samples the Moho underneath the Chyulu Hills area (bottom Fig. 4b), arrives about 0.6-1.3 s later than expected for a Vp/vs-ratio of 1.73. A similar delay of SMS relative to PMP does not appear on the P- and S-wave record sections of shotpoint CHS (Fig. 5), where PMP and SMS-reflection points are located outside the Chyulu Hills area (bottom of Fig. 5a,b). Detailed 2D-modelling of Pi3 and Si3 , which are almost coincident on record sections ATH and CHS, reveals a Vp/vs-ratio of 1.76 for the upper part of the lower crust. This Vp/vs-ratio, however, is not high enough to cause the observed delay of 0.6-1.3 s of SMS relative to PMR To match these travel times, a zone of decreased S-velocity (S-LVZ) is required. Its location must be restricted to the lower part of the lower crust beneath the Chyulu Hills area in order to explain both the coincidence of Si3 and Pi3 and the observed delay of SMS relative to PMP on record section ATH.

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Fig. 4 (continued). (b) Top. Trace-normalised radial component S-wave record section for shotpoint ATH, reduced at 3.46 km/s. Plotted are the ray-traced phases Sis and SMS (solid lines) based on the S-wave model described in Fig. 6. In addition the ray-traced phases Pi3 and PMP are plotted (dashed lines) by multiplying their travel times by 1.73 based on the P-wave model described in Fig. 6. Bottom. Ray tracing of phase Si3 and SMS through the S-wave model (see Fig. 6).

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Fig. 5. (a) Top. Trace-normalised vertical component P-wave record section for shotpoint CHS, reduced at 6 km/s. See Fig. 4a for further explanation. Bottom. Ray tracing of phases Pi3 and PMP through the P-wave model (see Fig. 6),

As the determination of the S-wave velocity (Vs) and the vp/vs-ratio in the S-LVZ is only based on the observed delay of SMS relative to PMP, we need to

determine the dimensions of this body at the base of the crust underneath the Chyulu Hills. Unfortunately, the poor quality of the S-wave dataset excludes a

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refinement of the S-LVZ based on the S-wave refraction data alone. For example the determination of the maximum epicentral distance up to which phase

Si3 is recorded is not possible. The knowledge of this particular epicentral distance, however, would be needed to determine the corresponding penetra-

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tion depth of Si3 which depends on the vs-gradient in the upper part of the lower crust. This maximum penetration depth defines the possible upper boundary of the S-LVZ because refracted rays penetrating into the S-LVZ are not observed at the surface. A further investigation of this anomaly is only possible through a combined interpretation of the teleseismic tomography and the refraction seismic Pand S-wave datasets. As the teleseismic tomography yields a - 3 % P-wave anomaly in the lower crust beneath the Chyulu Hills, we consider first whether the refraction seismic P-wave dataset could be explained by a similar feature. Assuming that a possible P-LVZ has the same vertical extension as the S-LVZ, the high-quality P-wave data provide more detailed information on the dimension of the LVZ than the S-wave data. On the P-wave record section ATH phase Pi3 can be clearly observed up to an epicentral distance of 330 km (top Fig. 4a). On the reverse profile from RUK phase Pi3 can be observed out to the maximum epicentral distance of 310 km (Novak et al., 1997, fig. 6). 2D-modelling of these Pi3 phases (ATH ATH

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and RUK) establishes a lower crustal vp-gradient of 0.009 4- 0.004 1/s underneath the Chyulu Hills area. Considering the observed maximum epicentral distance of phase Pi3 and the lower crustal vp-gradient, the Vp-structure is well determined down to 30 4- 5 km underneath the Chyulu Hills (bottom Fig. 4a). For this part of the lower crust a P-LVZ is not compatible with the seismic observations and can safely be excluded. Below 30 -4- 5 km depth Pi3 cannot provide constraints for the P-velocity structure. Hence the introduction of a P-LVZ in the refraction seismic model is only feasible in the lower part of the lower crust. The vertical restriction of the P-LVZ does not conflict with the observed - 3 % anomaly of the teleseismic tomography model because the tomographic inversion algorithm averages the travel time delays along the whole lower crustal layer from a depth of 23 to 43 km (Fig. 3a). This implies larger perturbation values than 3% for the vp-anomaly as the P-LVZ does not extend vertically over the whole lower crust. Assuming an upper boundary of the P-LVZ at about 30 km depth, a reduction of Vp of about 4% is needed

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to be consistent with the - 3 % anomaly of the teleseismic tomography model. Translated to absolute velocities this results in a vp-reduction to about 6.8 km/s at about 30 km depth from 7.05 km/s, representing the average P-velocity at 30-44 km depth in the 'original' model (Novak et al., 1997). The introduction of a P-LVZ at the bottom of the crust of the 'original' refraction seismic model requires the crustal thickness to be decreased in the area of the P-LVZ to rematch the observed P~P phase. Travel time calculations for PMP were made for various thicknesses of the P-LVZ which correspond to an upper boundary at 30 + 5 km depth and for various vp in the P-LVZ which correspond to the - 3 % vp anomaly of the teleseismic tomography model. The calculations show that the required decrease in crustal thickness in the area of the P-LVZ is only of second order as the maximum required decrease is less than 0.3 kin. Assuming that the vertical extension of the P-LVZ and S-LVZ is identical, a combined 2D-modelling of phases PMP and SMS on record sections ATH (Fig. 4), CHN and CHS (Fig. 5) yields further restrictions on the shape of the LVZs which allow to determine the S-velocity and the vp/vs-ratio in the LVZ. PMP and SMS on record section ATH sample the northwestern and central part of the LVZ. On record section CHN, PMP and SMS are restricted to the central and southeastern part of the LVZ, whereas on record section CHS, PMP and SMS reflection points are located outside the LVZ (southeast). Both LVZs have an oval shape, however the P-LVZ has a smaller horizontal extension than the S-LVZ. If we assume a LVZ with an upper boundary at 30 km depth, a P-velocity of 6.8 km/s is required to match PMP on record section ATH (see discussion above). Based on these values, 2D-modelling of phase SMS yields an S-velocity of 3.7 km/s in the LVZ which gives a vp/vs-ratio of 1.84. The alternative models with the upper boundary of the LVZ at 25 km and 35 km depth, based on the error bar of the established lower crustal vp-gradient, yield vp/vs-ratios of 1.81 and 1.93, respectively. If the upper boundary of the LVZ is located deeper than 35 km, which the seismic data cannot exclude, the resulting vp/vs-ratio is even higher. Fig. 6 shows the 'revised' refraction seismic model with the P and S-LVZ at the base of the lower crust beneath the Chyulu Hills area. Fig. 7 shows the

same 'revised' model at the same scale as the crosssection through the teleseismic tomography model (see Figs. 3 and 13). We point out that the refraction seismic and tomographic P-LVZs occur in the same region. A further constraint for the existence of a lowvelocity body at the base of the lower crust is given by the gravity model along line F (Fig. 8). In a first approximation the density structures of the 'original' (Novak et al., 1997) and the 'revised' P-velocity models (upper boundary of P-LVZ at depth of 30 kin; Fig. 6) were calculated as simple linear functions of the P-velocities (Luetgert, 1992). Fig. 8 shows the calculated gravity response of the 'original' and 'revised' P-velocity models. The slightly thicker crust in the area beneath the Chyulu Hills in the 'original' P-model compensates to some extent the reduced density in the LVZ in the 'revised' P-model. Nevertheless the 'revised' model achieves a better fit to the observed data than the 'original' model (Fig. 8). The increase of the vp/vs-ratio and the decrease of the seismic velocities in the LVZ are the seismological bounds for a combined seismic-petrological interpretation. We will discuss two end-member models for the lower crust. The first one assumes a solid crust (without melt) and the second one assumes magma chambers with partial melt.

4. Seismic-petrological model of the lower crust 4. ]. Plagioclase-rich nature of the lower crust The low-velocity body at the bottom of the lower crust beneath the Chyulu Hills is characterised by high vp/vs-ratios of 1.81-1.93. If a completely solid crust is assumed, the combination of such high vp/vs-values with a P-wave velocity of 6.8 -4- 0.1 km/s requires a plagioclase-rich modal composition like, for example, that of an anorthosite. We use a typical anorthosite composition (SiO2, 51.05%; TiO2, 0.63%; A1203, 25.57%; Fe203, 0.99%; FeO, 2.07%; MgO, 2.14%; CaO, 12.76%; Na20, 3.18%; K20, 0.62%; Le Maitre, 1976) to calculate equilibrium modal compositions and seismic velocities at P - T conditions expected in the lower crust. In our calculations we followed the procedure of Sobolev and Babeyko (1994). Thermodynamic data on wellequilibrated mantle and lower crustal xenoliths from

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CHN

CHYULUHILLS

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I I

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¢.

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0.0

100

!

-0.6 -1.2

120

-1.8

140

. 100

NW

150 Distance [km]

200

'

-z~

b) SE

Fig. 7. (a) Part of the 'revised' refraction model along line F (see Fig. 6) compared at the same scale to (b) cross-section through teleseismic tomography model (along line A-B, Fig. 2) along line F (see also Figs. 3 and 13). Displayed are perturbations of the compressional wave velocity in percent relative to the starting model (see text).

the Chyulu Hills volcanic rocks suggest temperatures of 700-800°C for the depth range of 30-40 km (Fig. 9). For T = 700°C and P = 1 GPa we obtain a plagioclase-rich lithology composed of 83 vol.% plagioclase (An64), 8 vol.% clinopyroxene, 5 vol.% garnet (PY43Alrn41Gr16) and minor quartz, kyanite, alkali feldspar and magnetite. For this garnet-bearing meta-anorthosite seismic velocities calculated by the method of Hashin and Shtrikman (1963) are

vp = 6.83 4- 0.05 km/s and vs = 3.78 4- 0.03 km/s resulting in vp/vs = 1.8l 4- 0.03. If a depth of about 25 km to 30 km is assumed for the upper boundary of the LVZ, these calculated values agree with the seismic model. However, if the LVZ is thinner (its upper boundary is deeper than 30 km), which cannot be excluded from the seismic data (see discussion above), the LVZ cannot be explained by an anorthositic composition as the corresponding Vp/vs-ratiois too high.

O. Novak et al. / Tectonophysics 278 (1997) 187-209

200

100

200

300

400

0 -40 E

-8o

-120 -i60 i

CHYULU,

HILL:~

I

t

I-

,

J

0

10 ~,

20

--,

30

C3

4O 50 0

1

0 NW

~

I

50

'

100

1

150

I

" !

.... ] . . . . . . q

200 250 Distance [km]

I

300

I

I

350

'

I

400 SE

Fig. 8. Density models along line F, derived from the seismic P-wave model (Fig. 6). Densities in g/cm 3. Solid line: gravity response of the converted 'revised' P-velocity model. Dashed line: gravity response of the converted 'original' P-velocity model. Note that the crust is thicker by about 0.2-0.3 k m at the base of the LVZ if the gravity response is calculated for the converted 'original' P-velocity model.

The calculated modal composition for the lowvelocity body (possible upper boundary between 25 km and 30 km) at the bottom of the lower crust beneath the Chyulu Hills is also compatible with the compositional range of lower crustal xenoliths erupted in the northeastern part of the Chyulu Hills. These xenoliths show a progressive change from igneous-textured anorthosites consisting of plagioclase, clinopyroxene, and minor orthopyroxene to foliated granuloblastic granulites made up of pla-

gioclase, clinopyroxene, garnet, minor kyanite, futile and quartz. In the igneous-textured anorthosites, reaction between coarse-grained primary plagioclase and orthopyroxene resulted in complex corona structures around relict orthopyroxene grains indicating arrested reactions. Even the foliated granuloblastic granulites show relict corona structures (HenjesKunst and Altherr, 1992; R. Altherr, unpubl, results). This textural evidence confirms the transformation of primary igneous mineral assemblages to meta-

201

O. Novak et al./Tectonophysics 278 (1997) 187-209

Temperature [°C]

500 0,

600

700

800

900

1000

1100

1200

1300

1400 0

) ~7

[]

"2 a.

40

Wlrn2

[]

-

~ 2

~3 Q- 4

70.~

............ , 100 3 !

Chyulu Hills: Grt-bearing xenoliths O equilibrated grt Iherzolites [] equilibrated grt pyroxenites • equilibrated grt granulites ', ~ unequilibrated grt pyroxenites

- 130

- 160

Temperature [°C]

600 0

700

800

900

1000

1100

1200

1300

1400

40 "-'2 Q. O

~3 V Q- 4

7o

100 3

130 150

Fig. 9. Pressure-temperature (P-T) data for xenoliths from the northwestern part of the Chyulu Hills volcanic field. (a) P - T estimates for garnet-beating pyroxenites, lherzolites, and granulites were obtained by graphical best fit combining the two-pyroxene thermometer and the Al-in-orthopyroxene barometer of Brey and Ktihler (1990). Except for two garnet pyroxenites which show minor compositional zoning in pyroxene grains, all the other samples are well equilibrated and yield P - T estimates plotting near to a 70 mW/m 2 steady-state model geotherm (Chapman, 1986). (b) For garnet-free harzburgites, a pressure range is given for each xenolith. Minimum pressure limits are constrained by the absence of plagioclase in peridotite systems (Gasparik, 1987) and are generally lower than pressures at the Moho (~1.3 GPa) which is at a depth of 40-44 km according to refraction seismic model (Novak et al., 1997). Maximum pressures for harzburgites were obtained from spinel compositions according to the formula given by Webb and Wood (1986). In both diagrams, the position of the apparent lithosphere-asthenosphere boundary (LAB) is shown, according to thermobarometric data (Henjes-Kunst and Altherr, 1992; V. Garasic and R. Althem unpubl, results). AAC refers to the adiabatic upwelling curve for normal temperature asthenospheric mantle in the absence of significant amounts of melting (potential surface temperature Tp = 1280°C) from McKenzie and Bickle (1989).

202

O. Novak et al. / Tectonophysics 278 (1997) 187-209

morphic assemblages of the granulite facies. Many unequilibrated pyroxenitic xenoliths from the Chyulu Hills also show a transformation from an earlier (igneous) high-temperature stage to a later low-temperature metamorphic stage (Henjes-Kunst and Altherr, 1992; V. Garasic and R. Altherr, unpubl, results). In order to explain both the pyroxenitic and the plagioclase-rich xenoliths we speculate that mafic melts underplated at the Moho were differentiated during crystallisation whereby pyroxenitic cumulates and anorthositic differentiates were formed. During subsequent cooling the igneous-textured rocks were partially transformed to metamorphic granulitic mineral assemblages. In such a scenario the low-velocity body at the base of the lower crust beneath the Chyulu Hills would correspond to the plagioclaserich differentiates. 4.2. Partial melt in the lower crust?

The model of a plagioclase-rich completely solid lowermost crust beneath the Chyulu Hills is not unique. The observed seismic characteristics of this body may also be explained by the presence of melt. Our interpretation of seismic velocities (Vp and Vs) in partially molten rocks is, however, limited since (1) the refraction seismic dataset does not allow for a precise determination of shear attenuation, (2) literature mostly focuses on models based on mantle rocks, and (3) the topology of the partially molten rock system ('melt geometry') is not known. If velocity reductions are interpreted in terms of partial melt the influence of the melt geometry is significant. We refer to Mavko (1980) who calculated the effective bulk and shear moduli for olivine and pyroxene at 20 kbar as a function of melt fraction. The similarity of results for these two minerals suggests that the result may apply at least qualitatively to a range of modal compositions. The refraction seismic model (Fig. 6) shows a decrease in ve of about 4% (7.05 to 6.80 kin/s) and a concomitant decrease in Vs of about 8% (4.00 to 3.70 km/s, corresponding to a Vp/Vs-ratio of 1.84) in the low-velocity body beneath the Chyulu Hills if we assume the upper boundary of the LVZ at a depth of 30 kin. The value of Vp = 7.05 km/s is the average P-wave velocity at the depth range of about 30 km to 44 km according to the 'original' refraction seismic

model of (Novak et al., 1997). The corresponding Vs = 4.0 km/s results from a Vp/vs-ratio of 1.76 as observed in the upper part of the lower crust. Following Mavko (1980) who calculated seismic velocities and attenuation in partially molten rock systems with a tube-like geometry of the melt phase, the observed velocity reductions result in about 5% melt for Vp and about 4% for vs. Faul et al. (1994) found that approximately 75% of melt occurs in low aspect ratio inclusions (a < 0.1, ratio of the minor to the major axis of an ellipse). These low aspect ratio inclusions reduce seismic velocities much more efficiently than melt in triple junction tubes. Following Faul et al. (1994) the observed reductions of 4% (Vp) and 8% (Vs) result in values of 2.2% (vp) and 2.4% (vs) melt for the case of low aspect ratio inclusions and 4% (Vp) and 3.5% (Vs) melt for a triple junction geometry. We conclude that on the basis of available experiments relating melt and seismic velocities (Mavko, 1980; Faul et al., 1994) the observed reductions in seismic velocities allow for not more than 2-5% of melt in the lower part of the lower crust beneath the Chyulu Hills. If present, this melt either represents in-situ partial melt or was intruded from greater depths. Petrological arguments may help to resolve this question. Volcanic rocks from the Chyulu Hills display pronounced chemical and isotopic evolutionary trends (Haug and Strecker, 1995). Groups 1 to 4 show a progressive evolution from highly silica-undersaturated Ti-rich foidites to less undersaturated basanites and alkali basalts (Fig. 10). This trend in chemical compositions is accompanied by a systematic variation in isotopic signatures (Fig. 11). The oldest lavas (group 1) have very radiogenic Pb isotope compositions suggesting the contribution of a HIMU-type mantle component. With decreasing age, the isotopic signatures of the younger lavas (groups 3 and 4) shift towards less radiogenic Pb and Sr compositions. This evolution is compatible with an increasing contribution of a DMM-type mantle component. In Pb-Pb isotopic variation diagrams most samples of groups 1 to 4 plot near to the Northern Hemisphere Reference Line (NHRL) as defined by Hart (1984). The near-primitive nature of all the older (groups 1-4) volcanic rocks (i.e. high Mg# and high Ni and Cr abundances), their highly fractionated

O. Novak et al. / Tectonophysics 278 (1997) 187-209

203

0.7041

a'ane SIc.Y

6

~

Foidite

~5

•

++4 +

0

0.7039

+ basalt 0.7037

+ ¢10

=

•

.

+

~4 l+Gro.p51 z

/1° Group4 | 3 [-[1:3Group 3 | [ [ • Group 2 I

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+ 1:3- o

Basalt

basalt

,

HIMU

n

D

Picro-

,

~pe

0.7035

o o

i

,

i

43 45 41 SiO 2 [wt%]

,

i

47

,

i

49

t

0.7033

/

51

Fig. 10. Nomenclature diagram for total alkali vs. silica (TAS) (Le B a s e t al., 1986) with data points of representative lavas from the Chyulu Hills. Note the chemical evolution from the oldest foidites (groups 1 arid 2) to the youngest alkali basalts and trachybasalts (group 5). For further explanation see section on magma chambers in the continental crust in the text.

chondrite-normalized rare earth element patterns (i.e. high La/Yb and Tb/Yb ratios), their isotopic compositions, and the widespread occurrence of mantle xenoliths suggest that the lavas were erupted without being significantly modified by continental crust. In marked contrast, the younger group-5 alkali basalts are more silica-rich (Fig. 10). The ubiquitous presence of crustal-derived xenocrysts like feldspar and quartz indicates contamination of these magmas by crustal material. Furthermore, the isotopic compositions of most group-5 rocks are characterised by significantly elevated S7Sr/S6Sr and 2°Tpb/2°4pb values (Fig. 11). Such a component which is not seen in the older lavas, suggests either the presence of an older sedimentary component in the mantle source of the parent magmas or, more likely, contamination of the ascending magmas by crustal materials. Such a contamination points to a stagnation of the rising magmas at some crustal level and would support the hypothesis of melt pockets in the lowermost crust. 5. The uppermost mantle In the uppermost mantle, P-wave velocities revealed from the refraction seismic experiment are

15.70 L

g

15.65

15.60

15.55 19.0

19.5

20.0

20.5

206pb/204pb

Fig. 11.87Sr/S6Sr vs. 2°6pb/2°4pb and 2°7pb/2°4pb vs. 2°6pb/2°apb variation diagrams with data points for representative volcanic rocks from the Chyulu Hills. The oldest lavas (group 1) have very radiogenic Pb compositions suggesting the contribution of a HIMU-type mantle component. With decreasing age of the rocks (groups 3 and 4) their isotopic signatures shift towards less radiogenic Pb compositions, compatible with an increasing contribution of a depleted mantle (DMM-type) component. Except for one sample, all volcanic rocks of groups 1 to 4 plot near to the Northem Hemisphere Reference Line (NH R L, Hart, 1984). Late Pleistocene to Recent lavas (group 5; shaded fields) are characterised by variably elevated 87Sr/S6Sr and 2°7pb/2°6pb ratios suggesting the involvement of crustal material.

lower (7.9 • 0.05 km/s) beneath the Chyulu Hills area than in surrounding areas (8.1-8.2 kin/s; Figs. 6 and 7a). This observation is well in accordance with the results from teleseismic tomography. Fig. 3b and Fig. 7b show that the low-velocity anomaly seen in the lower crust underneath the Chyulu Hills continues into the uppermost mantle (layer 4; 43-66.5 km). The velocity contrast between this low-velocity anomaly and the surrounding uppermost mantle is as high as 3.5%. In the cross-section through the

204

o. Novak et al./Tectonophysics 278 (1997) 187-209

tomographic model (Fig. 7b) it can be seen that the anomaly is limited to the upper 70 km of the lithosphere. No continuation to greater depths is visible. The seismic features of the uppermost mantle are further constrained by petrological evidence. Mantle xenoliths erupted in the northern part of the Chyulu Hills comprise garnet lherzolites, spinel harzburgites (some of which contain minor amounts of clinopyroxene), garnet pyroxenites, and pyroxenites devoid of garnet. According to thermobarometric data, these xenoliths are either derived from the lithosphereasthenosphere boundary (LAB) region or were sampled from the upper part of the lithospheric mantle. Whereas xenoliths from the LAB region are generally well equilibrated, those from the upper part of the lithospheric mantle are either equilibrated or show the effects of a heating event after equilibration (Fig. 9), as suggested by the compositional zoning patterns of their pyroxenes (Henjes-Kunst and Altherr, 1992; V. Garasic and R. Altherr, unpubl. results). The fact that some xenoliths are well equilibrated while others have been significantly heated indicates that heating was a local rather than a regional phenomenon and was probably related to the intrusion of magmas into the uppermost mantle. Using the observed compositional zoning patterns of orthopyroxene grains in unequilibrated harzburgites in combination with the estimated ranges of diffusion coefficients for AI and Ca in pyroxenes (Brady and McCallister, 1983; Sautter et al., 1988; Opper, 1989; Sautter and Harte, 1990) result in maximum and minimum durations of about 210 ka and 1.3 ka, respectively, for the heating event. These estimates clearly indicate that heating suffered by the xenoliths prior to sampling by the host magmas is related to the Quaternary volcanic activity and not due to conductive regional heating of the lithosphere. The wide-angle reflection data contain evidence for upper mantle reflectors in the area beneath the Chyulu Hills. One example from record section RUK is displayed in Fig. 12. At a distance range from 260 km to 310 km a correlated phase (d) can be identified after the first arrivals. Preliminary modelling suggests a depth of approximately 75 -4- 3 km based on reversed observations on record section ATH (Byrne, 1997).

6. The lithosphere-asthenosphere boundary region Deep mantle anomalies in the teleseismic tomography model are displayed in Fig. 13. In the depth range between 66.5 and 90 km (layer 5), velocity perturbations from - 2 . 2 % to +1.7% are observed. Compared to layer 4, where a low-velocity zone is located beneath the eastern and southeastern part of the Chyulu Hills (Fig. 3b), the position of the low-velocity zone in layer 5 is shifted about 20 km to the east of the central ridge of the Chyulu Hills (Fig. 13a). Compared to the reference model (Ritter and Kaspar, 1997), the mantle beneath the northern part of the Chyulu Hills consists of material with slightly higher P-wave velocities. The low-velocity zone at the southern boundary of the model seems to be an artefact of the inversion. In layer 6 (90-115 km; Fig. 13b), the velocity perturbations range from - 2 . 3 % to +2.2% with the extreme values occurring at the less well resolved borders of the model. In accordance with layer 5 above we find reduced velocities to the east of the Chyulu Hills and increased velocities around this zone, especially farther to the east. As in layer 5, the northwestern part of the Chyulu Hills volcanic field is characterised by increased velocities, a feature consistent down to a depth of 140 km (Fig. 13c). The extreme velocity contrast in the western part of the model (Fig. 13b) is regarded as an inversion artefact. The depth range between 115 and 140 km in the teleseismic tomography model (layer 7; Fig. 13c) corresponds to the asthenosphere as suggested by geothermobarometric data on equilibrated garnet lherzolite xenoliths (Fig. 9a). Equilibration conditions of these xenoliths are close to the adiabatic upwelling curve for normal temperature asthenospheric mantle material (AAC). In summary, the lower part of the lithosphere and the uppermost part of the asthenosphere are characterised by two interesting features. (1) The low-velocity anomaly beneath the southeastern part of the Chyulu Hills volcanic field is limited to the upper 70 km (Fig. 7b). (2) There is another low-velocity anomaly at greater depths which appears about 20 to 30 km further east and reaches down at least to the bottom of the tomography model (140 km depth).

O. Novak et al./Tectonophysics 278 (1997) 187-209

205

12 11 10 9 8 7

CO ,

E

.m

t-

6 5

4

3 2

0 -300

NW

-200

-100

Distance [km]

0

SE

Fig. 12. Trace-normalisedvertical componentP-wave record section for shotpoint RUK, reduced at 8 km/s. The inset shows an enlarged part of the record section displaying a mantle reflector (phase d). The resolution of the tomography model does not allow us to decide whether these two low-velocity anomalies are connected. At the present state of interpretation we tentatively assume the presence of two distinct anomalies. The absence of a deep lowvelocity root directly beneath the Late Pleistocene to Recent part of the volcanic field seems to rule out an upwelling of the asthenosphere as proposed for the Kenya rift proper. This is in line with the small magnitudes of the velocity perturbations which also exclude a major heating from the asthenosphere in form of a mantle diapir under this part of the eastern rift flank. The results of the teleseismic tomography and the geothermobarometry correspond well with the gravity model along line F (Fig. 8). No density variations

at depths greater than 60 km are necessary to fit the observed data (Fig. 8). This result is in contrast to the gravity model across the southern part of Kenya rift 150 km to the west (Birt et al., 1997), where gravity modelling suggests that compensation of the Bouguer anomaly occurs below the Moho, invoking the need for deep density contrasts. There the regional gravity gradient has been used as supporting evidence for mantle-plume type circulation beneath the uplifted East African Plateau to the west of the Kenya rift. For the Chyulu Hills we favour a model where small magma bodies segregate from the asthenosphere-lithosphere boundary and ascend to the Moho where they stagnate and assimilate lower crustal material (Fig. 14). Such a scenario could

O. Novak et al. /Tectonophysics 278 (1997) 187-209

206

100 50 50 0

0

-50

a)

-50

b) -50

-50 0 50 distance from center [kin]

Layer 7 ( 115-140 km)

0

50

100

velocity perturbation [%]

100

-2.5 -1.8-1.2 -0.6 0.0 0.6 1.2 1.8

50

2.5

projection center: 37 45'E - 2 20'S KRISP 93: Teleseismic Group

0

c)

-5o

-I00

-50

0

50

I00

Fig. 13. Layers 5 to 7 of the teleseismic tomographymodel. See caption to Fig. 3 and text for further explanation. Displayed are relative variations of the compressional wave velocityin percent. Depth ranges are indicated above seismic images. well explain the upper mantle and lower crustal low-velocity anomaly observed in the teleseismic tomography model. The confinement of this anomaly to a depth of less than 70 km does not necessarily mean that the area of stagnated magmas is completely disconnected from the asthenosphere. Due to their limited size, existing magma channels cannot be resolved by teleseismic tomographic methods. It is speculated that the deeper (> 70 km) low-velocity feature in the eastern part of the tomographic model

area (Fig. 13) could represent recent upwelling of magmas, as no surface expression is observed.

7. Conclusions An integrated seismic and petrological approach has been used to decipher the deeper structure of a young rift-related volcanic field--the Chyulu Hills in Kenya. The combined evaluation of datasets obtained from both refraction seismics and teleseismic

O. Novak et al./ Tectonophysics 278 (1997) I87-209

Northern cones (Xenolith-bearing) CHN Emali / ~ / ~ / ~

207

Shaitani Volcano

Chyulu Hills

CHS

v

}it

upper crust

i.

20

lower crust

........................................

40

___

-22 .......

6O

J¢

c~

Depleted harzburgite lithospheric mantle

100

O

O

o

•

.

O o

m•

Grt-pyroxenites Fertile peridotites

80

~&agnat;on&

assimilation)

_

E

Active magma chambers

0

"plums" (Iow-T melting material)

.

°i O

O

•

O

I O eO o

0a

i

t

.

o

_'

C:,

O

LAB

120

asthenosphere

~

heat

heat + melts

140 100

NW

1

Distance [km]

200

SE

Fig. 14. Cartoon summarising the major features of the deep structure of the Chyulu Hills volcanic field. The northern cones are related to the area of Emali (Fig. 2), the main range extends northwestwards from the volcano Shaitani to shotpoint CHN (Fig. 2).

tomography allows to overcome the usual uncertainties inherent in the individual methods and thus yields better constraints on seismic models. Petrographic and geothermobarometric data on mantle xenoliths provide further information on the compositional and thermal structure of the lithosphere. The main results of these integrated seismic-petrological investigations are presented in Fig. 14. (1) Seismic P- and S-wave velocities determined for the lower crust in the refraction seismic experiment suggest the presence of a low-velocity body with an upper boundary at a depth of 30 4- 5 km extending to the Moho (44 km depth). The low-velocity body is characterised by a high Vp/vs-ratio ranging from 1.81 to 1.93 according to the maximum and minimum possible extension of this body. Two end-member models can explain the observed reduction of vp and Vs as well as the increased vp/vs-ratio in the lower part of

the lower crust beneath the Chyulu Hills: assuming a solid crust the first model suggests a feldspar-rich (anorthositic) nature of this part of the lower crust, which is also compatible with the widespread occurrence of (meta-) anorthositic lower crustal xenoliths. The second model considers the possibility of partial melt in the lower crust. 2-5% partial melt can explain the observed reduction of Vp and Vs, which is compatible with the isotopic signature of erupted volcanic rocks suggesting contamination of the ascending magmas by crustal materials. We favour the 'partial melt' model because it is valid for the whole range of possible extensions of the LVZ (upper boundary between 25 km and 43 km), in contrast to the 'anorthositic' model which is only valid if the upper boundary of the LVZ is located at 25-30 km depth. (2) Geochemical and isotopic characteristics of volcanic rocks and the presence of abundant mantle

208

O. Novak et al. / Tectonophysics 278 (1997) 187-209

and lower crustal xenoliths suggest that the oldest lavas rapidly ascended from the lithosphereasthenosphere boundary region to the surface. Fig. 14 illustrates this process: for the northwestern part of the Chyulu Hills long arrows indicate the immediate ascent of magma from the lithosphereasthenosphere boundary to the surface. In the southeastern part of the volcanic field the youngest magmas, however, stagnate beneath the Moho and cause local heating of the surrounding mantle rocks. This is in accordance with the results of the refraction and teleseismic experiment which reveal a velocity decrease of about 3.5% for the uppermost mantle underneath the Chyulu Hills. The stagnation of the youngest magmas beneath the Moho is also indicated in Fig. 14 by short arrows and symbolic magma blobs. The ubiquitous presence of crustal xenocrysts (plagioclase and minor quartz), comparatively high 2°7pb/2°4pb and 87Sr/86Sr ratios (relative to the NHRL), and the fractionated nature of these magmas suggest contamination by material from the lower crust. Such a contamination might have taken place in crustal magma chambers. (3) The low-velocity anomaly underneath the southeastern part of the Chyulu Hills volcanic field is limited to the upper 70 km of the lithosphere. The absence of a deep low-velocity root and the gravity model seem to exclude an upwelling of the asthenosphere as proposed for the Kenya rift proper. A model is favoured where small magma bodies ascend from the asthenosphere to the Moho where they stagnate and assimilate lower crustal materials (Fig. 14). (4) Thermobarometric data on well equilibrated mantle xenoliths indicate an apparent lithospheric thickness of about 105 kin. The P-wave tomographic model shows almost no velocity perturbations for the depth range between 115 and 140 km (layer 7; Fig. 13c) directly beneath the Chyulu Hills area. There is no evidence that the low-velocity anomaly, which is observed about 20-30 km further east, is related to the magmatism of the Chyulu Hills volcanic field.

Acknowledgements Several Kenyan authorities including the Office of the President, Kenya Wildlife Service and the Kenya Agricultural Research Institute in Kiboko granted

permission and help. Technical assistance by H. Frohna-Binder, M. Jordan, U. Geilenkirchen, G.H. Haug, H.-R Meyer, I. Salzmann, and K. Wacker (?) is gratefully acknowledged. We thank A.W.B. Jacob, J. Mechie, C. Prodehl, U. Achauer and M. Strecker for discussions. We are grateful to M.H.R Bott, G. Fuis and J. Mechie for valuable comments on this paper. This study was funded by the Deutsche Forschungsgemeinschaft within the framework of the Collaborative Research Center 108 in Karlsruhe. This is SFB 108 publication no. 553, Geophysical Institute Karlsruhe contribution no. 715, and Mineralogical Institute Heidelberg contribution no. 2-20.

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