Download An integrated geophysical analysis of the upper crust of the southern

Geophys. J. Int. (2001) 147, 543–561
An integrated geophysical analysis of the upper crust of the
southern Kenya rift
Silas M. Simiyu1 and G. Randy Keller2
1
2
Geothermal Project, The Kenya Electricity Generating Company, PO Box 1202, Naivasha, Kenya. E-mail: [email protected]
Department of Geological Sciences, University of Texas at El Paso, El Paso, Texas 79968, USA. E-mail: [email protected]
Accepted 2001 June 18. Received 2001 June 13; in original form 1999 May 27
SUMMARY
Previous interpretations of seismic data collected by the Kenya Rift International
Seismic Project (KRISP) experiments indicate the presence of crustal thickening within
the rift valley area beneath the Kenya dome, an uplift centred on the southern part of
the Kenya rift. North of the dome, these interpretations show thinning of the crust and
an increase in crustal extension. To the south near the Kenya /Tanzania border, crustal
thinning associated with the rift is modest. Our study was aimed at further investigating
crustal structure from this dome southwards via a detailed analysis focused on upper
crustal structure. We used results from surface geological mapping, drill hole data from
water wells and geothermal exploration wells, KRISP 85 seismic data for a profile across
the rift, KRISP 85 and 90 seismic data for a profile along the rift axis and KRISP 94
seismic data for a profile crossing southernmost Kenya to constrain gravity modelling
and construction of integrated models of crustal structure.
Our integrated analysis produced the following results concerning the structure and
evolution of the southern Kenya rift: (1) the graben master faults are consistently located
along the western margin of the rift valley, and there is no evidence for half-graben polarity
reversals for a distance of about 300 km; (2) there is no axial (north–south) crustal
symmetry with respect to the apex of the Kenya dome, and the crustal thickness may be
as much related to pre-rift crustal type and thickness as it is to crustal thickening and
modification by magmatic processes; (3) the pre-existing lithospheric contrast between
the Archaean and Proterozoic basement terranes played a significant role in the location
and structural geometry of the rift; (4) south of latitude 1uS, low velocities and densities
observed under the western flank of the rift probably represent reworked Archaean
Tanzanian craton; (5) magmatic modification of the upper crust is modest except near
the major Quaternary volcanic centres that produce a series of isolated axial gravity
highs; (6) the other element of the axial gravity high is an intrarift horst block that
extends along the axis of most of the rift valley.
Key words: continental rifting, crustal structure, gravity, refraction seismology.
INTRODUCTION
The East African rift zone is the classic example of an active
continental rift, and its eastern arm extends from southernmost
Ethiopia, through Kenya and into Tanzania and is variously
called the Eastern, Gregory or Kenya rift, the term we will
employ. The southern portion of the Kenya rift (Fig. 1) has been
the target of a number of recent geophysical and geological
investigations, many of which are summarized in two special
volumes (Prodehl et al. 1994, 1997). This is due in part to
the Kenya Rift International Seismic Project (KRISP), which
has focused attention on the region, but it is also the result
of the exploration efforts of the Kenya Electricity Generating
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RAS
Company. These studies have created a large new database,
which has not been fully exploited to elucidate the structure
of the upper crust. Here, new gravity measurements and drill
hole data from geothermal areas are combined with the KRISP
seismic data, data from water wells and geological constraints
in an integrated analysis of upper crustal structure. Our specific
goals were to investigate basement structures beneath the rift
valley and the extent of magmatic modification of the upper
crust.
Our focus on the upper crust is different from the previous
studies motivated by the KRISP effort that have targeted
lithospheric-scale structure. The combination of all of these
data allows us to pursue a level of detail that has not been
543
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S. M. Simiyu and G. R. Keller
340E
380E
10 N
NA
I
Mt Elgon
L. Baringo
L
AN
ELDORET
L. Bogoria
Mt Kenya
KISUMU
0
Aberdare
Range
Menengai
Mf
L. Victoria
0
z
Eburru
L. Naivasha
KENYA
Ol
Ok
Kf
Sf
Longonot
10 S
Suswa
NA
I
NAIROBI
TANZANIA
L. Magadi
20 S
AN
L
Tanzania
Craton
Chyulu
Range
L. Natron
Kilimanjaro
Tanzania craton
Rift volcanics and sediments
Off–axis major volcanoes
Mobile belt quartzites
with reworked Archean
Mobile belt gneisses and
other metasediments
Outline of the
Kenya dome
NAI - Nyanega-Athi-Ikutha shear zone & ANL - Aswa-Nandi-Loita shear zone
Volcanic ranges
Rift axis caldera volcanoes
Figure 1. Tectonic map of the southern Kenya rift showing the surface distribution of the major lithologic units (after Smith 1994). Mfz: Mau fault
zone; Sf: Sattima fault; Kf: Kinangop fault zone; Ol: Oloololo Escarpment; Ok: Olkaria volcanic centre
possible previously. Our approach included simultaneous 2-D
modelling of first arrivals from KRISP seismic lines in the
area (Fig. 2) and 2.5-D modelling of gravity profiles along
these lines and across key features where drill holes provided
important constraints. In addition, we used updated gravity
maps to provide a view of the lateral extent of features revealed
in the modelling. Throughout this process, all supporting data
available were employed as constraints. Our integrated approach
produced geophysical models with some details that cannot be
defended on the basis of a single data set. Finally, we produced
interpretative crustal-scale cross-sections along the three profiles
where we had the most constraints.
Geological setting
Baker (1986), Baker et al. (1988), Smith & Mosley (1993)
and Smith (1994) provide recent reviews of the geology and
evolution of the southern Kenya rift region (Fig. 1). Besides the
rift and associated volcanics, two major geological provinces
are present. The Archaean Tanzania craton to the west is characterized by exposures of granitoid gneisses and metavolcanics.
The northeastern boundary of this craton is the Aswa–Nandi–
Loita NW–SE-trending shear zone (Fig. 1), and in southern-
most Kenya, geophysical data place the boundary between
the Oloololo escarpment and Lake Magadi (Birt et al. 1997;
Simiyu & Keller 1998). Between these areas, the boundary is
not visible due to volcanic cover. Rocks of the Neo-Proterozoic
Mozambique belt are well exposed east of the rift valley and
consist mainly of gneisses, schists and migmatites that contain
the Nyangea–Athi–Ikutha shear zone that crosses the rift valley
(Fig. 1). Near the Kenya–Tanzania border, this mobile belt
contains many areas composed of highly fractured quartzites
and gneisses, which may have been overthrust onto a reworked
portion of the Archaean craton (Smith 1994).
Recent geological studies (Shackleton 1986; Smith & Mosley
1993; Stern 1994) suggest that the southern Kenya rift
developed entirely above thick cratonic lithosphere and that
the Mozambique belt rocks seen on the surface are thrust over
a reworked craton that extends at least 100 km east of the
craton /mobile belt outcrop boundary. Smith & Mosley (1993)
discussed the evidence for Archaean crust underlying surface
exposures of the Mozambique belt, which include zones of
thrust faulting containing slices of Archaean greenstones, gneissic
xenoliths in overlying Tertiary lavas east of the mapped
boundary that yield Archaean Nd model ages, and a variety
of metamorphic indicators. Smith (1994) suggested that the
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2001 RAS, GJI 147, 543–561
S. Kenya rift upper crustal structure
L. Baringo
ELDORET
L. Bogoria
0
CHE
o
KISUMU
Menengai
Mt Kenya
SOL
ELM
VIS
Eburru
NAI
L. Naivasha
Olkaria
OLO
EWA
NTU
MOR
TANZANIA
Longonot
SUS MAR
Suswa
NAIROBI
MAG
KAJ
o
2 S
Shotpoint
KRISP 85 }
KRISP 90
KRISP 94
KRISP 85 (EW)
KIB
L. Natron
Kilimanjaro
Figure 2. Index map of the southern Kenya rift showing the location
of the KRISP 85, KRISP 90 and KRISP 94 seismic refraction lines used
in this study. The locations of the shotpoints used in this study are
indicated by stars and are named as follows: Victoria (VIS), Oloololo
(OLO), Morijo (MOR), Magadi (MAG), Kajiado (KAJ), Ewaso Ngiro
(EWA), Ntulelei (NTU), Suswa (SUS), Margaret (MAR), Naivasha
(NAI), Elmenteita (ELM), Solai (SOL), Chepkererat (CHE) and
Baringo (BAR).
margin for the reworked craton trends broadly NW–SE and
that its northern limit is marked by the Aswa–Nandi–Loita
shear zone (Fig. 1). He also argued that in the Magadi–
Oloololo area (Fig. 1), the Archaean basement is buried by
gravitationally collapsed nappe structures.
The rift valley is a 50–70 km wide volcanic filled depression
bounded by major normal faults (e.g. Baker et al. 1972). There
has been considerable discussion about the role of the boundary
(suture or shear zone) between the Tanzania craton and the
Mozambique belt controlling the location of the rift at both
the regional and the local level (Nyblade & Pollack 1992; Smith
& Mosley 1993; Smith 1994; Hetzel & Strecker 1994; Birt et al.
1997; Tesha et al. 1997; Simiyu & Keller 1998). These authors
all feel that this boundary exerts some level of control on the
location of the rift valley. However, without seismic reflection
profiles, we cannot trace structural features from the surface to
depth.
The highest portion of the rift valley (y1.9 km) is the
Nakuru–Naivasha–Suswa region (Fig. 1), which contains three
large volcanoes with calderas (Menengai, Longonot and Suswa)
and large volcanic fields at Olkaria and Eburru (Fig. 1). These
volcanoes are located near the axis of the rift valley, but basalt–
trachyandesite volcanic cones and flows occur in the intervening low areas. Exposed volcanic rocks include the Mau and
Kinangop tuffs, which have been dated at 2.6–3.68 Ma (Baker
et al. 1988). The Kinangop tuffs /ignimbrites are exposed by the
eastern boundary fault while the Mau tuffs are exposed along
the western fault scarp. Deep drilling at Olkaria and Eburru
has shown that the Mau tuffs are downthrown to depths of
more than 2.5 km by major N–S-trending faults (Omenda
1994). Limuru trachytes that have been dated at 1.68–2.01 Ma
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2001 RAS, GJI 147, 543–561
(Baker et al. 1988) are exposed to the southeast of Suswa and
are thought to underlie most of the inner graben of the rift
floor. Plateau trachytes that have been dated at 0.98–1.23 Ma
overlie the Limuru trachytes. The plateau trachytes are exposed
in the southern part of Suswa and as horst blocks in the Eburru–
Elmenteita areas (Baker et al. 1988). Overlying these trachytes
are Quaternary volcanics that include basalts, trachytes, rhyolites,
tuffs and pyroclastics. These materials form a thickness of
about 1.5 km within the graben in the Olkaria–Eburru area
(Omenda 1998)
Late Quaternary volcanic activity within the rift valley was
concentrated in the axial region of the rift valley and resulted in
the formation of caldera volcanoes and volcanic cones (Fig. 1).
Menengai, which is the northernmost volcano in the Nakuru–
Naivasha basin, is located about 60 km north of Eburru and
consists almost exclusively of strongly peralkaline oversaturated
trachytes (Clarke & Woodhall 1987). Eburru is dominated on
the surface by pantellerites, and pantelleritic trachytes, but
trachytes are more abundant with depth. Syenite intrusives were
encountered at the bottom of some of the holes drilled in the
volcanic complex for geothermal exploration. The Elmenteita
and Ndabibi basaltic fields are located northeast and south
of Suswa, respectively, and both lie along the N–S axial fault
system (Fig. 1).
Previous gravity studies
The southern portion of the Kenya rift has been the focus of
many gravity studies, and gravity anomalies across the region
reflect the classic pattern of a long-wavelength minimum due
to mantle structure and shorter-wavelength anomalies due to
crustal structure (Fig. 3). These studies are summarized by
Swain et al. (1994) and provided an invaluable database that
laid the foundation for our study.
Bouguer gravity (mGal)
BAR
Suswa
* * ** * * *
*
* *
*
* * Observed Gravity
Estimated regional
–160
*
–180 * ** *
***
*
–200
Menengai
*
*** *
* *
*
**
**
**
* * *
* * * *
*** * *
Olkaria **
*
**
** * * *
*
South
North
(a)
Bouguer gravity (mGal)
36 o E
Mt Elgon
545
0
o
o
o
0.5 S
1.0 S
East
West
–160
Regional (Swain, 1992)
Regional (Fairhead, 1976)
Olkaria
–180
Observed gravity
–200
(b)
Regional Searle (1970)
0
50
100
150
200
DISTANCE (KM)
Figure 3. Observed gravity profiles and interpreted regional anomaly
profiles. (a) Axial (N–S) gravity profile along the Kenya rift valley
through Menengai, Olkaria and Suswa volcanoes (after Simiyu 1990).
(b) E–W gravity profile crossing the rift valley and passing through
Olkaria (after Simiyu 1990).
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S. M. Simiyu and G. R. Keller
Swain (1992) re-interpreted the axial gravity high from Lake
Baringo to Suswa using the KRISP 85 seismic results as constraints. He showed that the axial high in the southern Kenya rift
is sometimes associated with volcanoes, which produce local
maxima. He also showed that a minor densification of the basement by dyke injection could explain much of the axial high.
The KRISP 90 and 94 results were not available at the time he
did his analysis, and he did not have access to the new gravity data
and drill hole information presented here. However, we used his
analysis as a starting point for the gravity portion of this study.
Most of the earlier studies concentrated on the axial high
anomaly that is superimposed on three different wavelength
negative anomalies: (1) the primary regional negative anomaly
associated with the East African Plateau; (2) the secondary
regional negative anomaly associated with the mantle anomaly
centred on the Kenya dome; (3) local anomalies associated with
upper crustal features (e.g. Simiyu & Keller 1997). Attempts
to separate these anomalies using different methods [smoothing (Searle 1970), tying to outcrops of Precambrian basement
(Baker & Wohlenberg 1971), a combination of these approaches
including balancing of the thickness and density of the rift
volcanics (Fairhead 1976), and linear filters (Swain 1992)] have
yielded different residual anomaly wavelengths and amplitudes
(Fig. 3) and thus differing models of crustal structure. These
studies lacked seismic and drill hole control, and gravity data
coverage was sparse in many key areas at the time.
In an effort to integrate recent seismic studies with regional
gravity anomalies, Simiyu & Keller (1997) looked at the entire
East African Plateau region crossing both the eastern and
western arms of the East African rift. Their models of deep
(150 km) structure provide the regional framework for this
study. Simiyu & Keller (1998) analysed KRISP 94 seismic data
and gravity readings to derive a detailed model of the upper
crust at Lake Magadi, where the KRISP 94 line crossed the rift
valley. A goal of this paper is to extend this detailed analysis
northwards as far as the Menengai region.
Previous seismic studies
In the first major seismic study of the crust of the southern
Kenya rift, Henry et al. (1990) interpreted the KRISP 85
seismic refraction profiles, one of which partially followed
the axis of the rift from Lake Baringo to Lake Magadi (Fig. 2).
The overall data quality was fair, and the lower crustal phases
were not reversed on this profile. Interpretation of these data
showed that there were significant differences in crustal structure
between the northern and southern parts of the rift valley. The
rift fill thickness was found to vary between 1.5 and 5 km and
to be underlain by basement material of velocity 6.05 km sx1.
They found no evidence for an extensive axial mafic intrusion
in the crust suggested by some of the early studies.
The model of Henry et al. (1990) shows the presence of highvelocity material lying directly beneath the Menengai gravity
high. However, this model does not explain the corresponding
increase in the axial positive Bouguer gravity anomalies in the
area between Elmenteita and Suswa, where several thousand
metres of Cenozoic volcanics of relatively low density (inferred
from their low seismic velocity of 2.8–5.1 km sx1) blanket the
area. Their data were sparse south of the Elmenteita region,
and they did not suggest the presence of any crustal highvelocity /density bodies beneath this area, although it contains
a large positive gravity anomaly.
The KRISP 90 (Fig. 2) experiment was aimed at studying the
variation in crustal and uppermost mantle structure across and
along the axis of the Kenya rift. The analysis of the axial
profile by Mechie et al. (1994) and Keller et al. (1994a) was of
particular significance to this study. This analysis shows that
the crystalline upper crust, interpreted to be the Precambrian
basement, has velocities of 6.1–6.3 km sx1 and is 5–6 km below
surface in the Menengai–Suswa region. Mechie et al. (1994)
converted their velocity structure to densities using standard
relationships and showed that the regional Bouguer anomaly
can be explained well by southwards crustal thickening along
the rift axis. However they did not attempt to fit shortwavelength upper crustal anomalies in the Elmenteita, Naivasha
and Suswa areas. In this paper, we attempt to complete the
picture of the upper crust by modelling the local gravity
anomalies using well data and seismic constraints.
In 1994, an experiment was carried out across the southern
Kenya rift extending from Lake Victoria to the Indian Ocean
(Fig. 2). The KRISP 94 experiment showed only slight crustal
thinning under the rift at Lake Magadi, with the Moho being
about 35 km deep (Birt et al. 1997). However, Thybo et al.
(2000) studied the lower crust and concluded that significant
underplating occurred during rifting and that the amount of
thinning of the pre-rift crust was probably significant. East of
the rift valley, the crust thickens to 42 km in the Chyulu Hills
area with increased crustal velocities (Novak et al. 1997). Based
on the KRISP 94 results, a detailed interpretation of the upper
crustal structure near Lake Magadi was presented by Simiyu &
Keller (1998) and their results showed that the rift axial high
anomaly was caused by a basement horst structure at its axis.
SEISMIC INTERPRETATION
Seismic and gravity data tend to be complementary due to the
general relationship between seismic velocity and density. We
chose to model seismic data first and then use these results as
constraints for the integrated modelling of gravity profiles. The
three seismic profiles modelled are shown in Fig. 2. They are
the KRISP 85-90 axial profile, the KRISP 85 profile that
crossed the rift valley just north of Mt Suswa, and the KRISP
94 profile that crossed the rift valley at Lake Magadi.
Although picking first arrivals for these data (shown in Henry
et al. 1990; Mechie et al. 1994; Birt et al. 1997) was generally
straightforward, we fine-tuned this process using a commercial
seismic processing package. Frequency filtering, deconvolution,
trace normalization and other amplitude enhancement techniques were employed with some success using this interactive
program, which includes an algorithm for the automatic
picking of first arrivals. In order to be as objective as possible,
this algorithm was utilized to pick the Ps (direct wave) and Pg
(head /diving wave travelling in the crystalline upper crust) first
arrivals that we used for modelling the velocity structure. These
traveltime picks were confirmed manually and found typically
to differ from handpicked arrivals by less than 0.05 s.
Record sections and the first arrival traveltime picks were
used to determine apparent velocities for the Ps and Pg phases
that were used in the starting models. These velocities were
determined by least-squares straight-line fitting in which the
traveltimes were weighted by the reciprocal of the observational
error estimates. The starting model for the shallow velocity
structure was based on previous velocity results from KRISP
85, KRISP 90 and KRISP 94 (Mechie et al. 1994; Henry et al.
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2001 RAS, GJI 147, 543–561
S. Kenya rift upper crustal structure
sediments and volcanics. The KRISP 90 data were then used
to model the basement structure, with the exception that the
data from the 1985 Magadi shotpoint were used to reverse the
KRISP 90 Naivasha shotpoint. The Baringo (85,90), Naivasha
(90) and Magadi (85) shotpoints and the KRISP 94 EW model
at Magadi (Birt et al. 1997; Simiyu & Keller 1998) provided the
most rigorous constraints on the velocity model.
The model shows that the depth to Precambrian basement
along the axis of the rift varies from about 3.5 km beneath
Lake Baringo to a maximum of 6 km below Lake Naivasha
and back to about 3 km to the south at Lake Magadi (Fig. 4).
The rift fill is modelled as being composed of three layers. The
first layer has a velocity of 2.8 km sx1 and extends from model
coordinates 0–270 km, which is just south of Suswa (Fig. 4).
The second layer (average velocity of 4.2 km sx1) is 3 km thick
in the north at Lake Baringo and thins southwards to 1 km
beneath Lake Magadi. The velocity of this layer is variable,
ranging from 3.9 km sx1 beneath Lake Baringo to 4.6 km sx1
beneath Lake Magadi, with a localized low-velocity area
beneath Lake Naivasha (3.6 km sx1).
1990; Birt et al. 1997; Simiyu & Keller 1998) and knowledge
of the geology of the area. The modelling of the velocity
structure was carried out using the MacRay ray tracing
package (Luetgert 1992). For refracted phases to be generated,
layers were assigned positive velocity–depth gradients, typically
of around 0.01 km sx1 kmx1. The starting model was adjusted
by manual iteration until the observed traveltimes closely
matched the calculated traveltimes. The average difference
between calculated and observed traveltimes for the three models
is about 0.06 s, which is on the order of the picking error.
Axial seismic model
There were numerous sources along the axial profile (Fig. 2),
so the model derived from these data (Fig. 4) is well constrained except for a small area south of Suswa where there is a
data coverage gap. The KRISP 85 shots, most of which only
produced Ps arrivals, were modelled first as they were more
closely spaced and most useful for constraining the rift fill of
3
NORTH
Predicted
Observed
BAR
CHE SOL
547
SOUTH
T–X/6.00
2
1
0
ELM
NAI
MAG
–1
0
40
80
120
160
200
240
280
0
40
80
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Depth (km)
0
3
6
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NORTH
CHE SOL
BAR
Depth (km)
0
Kenya Dome
ELM
NAI
4.5
3.6
4.2
2.8
5.3
5.0
6.2
6.3
6
4.6
4.4
4.9
5.0
6.0
6.0
SOUTH
2.8
MAG
3.8
2.8
3.9
3
2.8
5.6
5.5
6.0
6.4
Precambrian Basement
9
6.6
6.2
6.3
12
6.45
6.7
6.7
6.45
6.70
15
0
40
80
120
160
200
240
280
Distance (km)
Figure 4. Upper crustal seismic velocity model (lower diagram) along the rift valley axis extending from Lake Baringo to Lake Magadi (Fig. 2). The
numbers are P-wave velocities in km sx1. Shotpoint locations are shown by stars (Fig. 2). The middle diagram shows the ray coverage provided by the
observed data and the velocity model. The upper diagram shows the fit between the observed traveltimes and those predicted by ray-trace calculations
for the velocity model.
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2001 RAS, GJI 147, 543–561
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S. M. Simiyu and G. R. Keller
The third and lowermost layer of the rift fill is thin at Baringo.
It thickens around Chepkererat, where it has a velocity of
5.0 km sx1. The velocity increases to 5.1 km sx1 beneath Lake
Naivasha, to about 5.4 km sx1 beneath the Suswa volcano,
and to 5.6 km sx1 beneath Magadi. A 5.6 km sx1 velocity
layer on the KRISP 94 EW profile was interpreted to represent
fractured Precambrian basement (Simiyu & Keller 1998). The
velocity of the uppermost portion of the crystalline crust varies
from 6.0 to 6.2 km sx1 at Baringo, increasing to a maximum of
6.4 km sx1 between Lake Naivasha and Suswa Volcano.
Suswa seismic model
The Suswa profile (Fig. 2) crosses the rift valley just north
of the Suswa volcano, which is capped by one of the largest
calderas in the rift valley. The original seismic profile was
designed as a cross-line that intersected the axial profile and is
in an area that has the highest gravity values along the entire
southern rift valley. The model (Fig. 5) is well constrained for
the volcanics within the rift valley because two shots produced
energy transmission effective enough to probe depths down to
T– X/ 6.00
3
WEST
about 6 km. Constraints from the N–S axial line also helped in
determining the starting model parameters. The western rift
flank volcanic thickness and basement velocities are also well
constrained by the Ntulelei (NTU) and Ewaso Nyiro (EWA)
shots, which are reversed. This model includes a y1 km thick
low-velocity (2.8 km sx1) surface layer that thickens to 2.2 km
just east of the Ntulelei (NTU) shotpoint at the western rift
boundary fault. Further east (Fig. 5), this surface layer thins
to less than 0.2 km at the Suswa (SUS) shotpoint. At the
Mt Margaret (MAR) shotpoint, this layer thickens to 0.5 km.
The next layer has a velocity of about 4 km sx1. Its thickness
is 1.2 km west of Ntulelei (NTU), increasing to a maximum of
about 2.0 km at the Suswa (SUS) shotpoint, then decreasing to
less than 1.7 km towards the eastern part of the rift floor at the
Mt Margaret (MAR) shotpoint. The lowermost layer of the rift
fill has velocities ranging from 5.0 to 5.5 km sx1. The layer we
interpret as crystalline basement has a velocity of 5.8 km sx1
west of the rift valley. This low value possibly indicates the
presence of quartzitic basement rocks, and based on the axial
line results, the velocity of this layer increases towards the rift
axis.
EAST
Predicted
Observed
2
1
0
EWA
NTU
SUS
MAR
–1
20
0
40
60
Rift
80
100
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Depth (km)
0
3
Precambrian Basement
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9
12
WEST
15
EAST
20
0
40
Distance (km)
60
NTU
–3 EWA
100
80
MAR
SUS
2.8
Depth (km)
3.8
5.0
0
5.8
5.8
3.8
5.2
6.0
3
2.7
3.9
5.0
3.9
4.5
5.2
5.0
5.5
6.0
Precambrian Basement
6.0
6
9
12
15
0
20
40
60
100
80
Distance (km)
Figure 5. Shallow seismic velocity model (lower diagram) along the KRISP 85 EW profile that extends from Ewaso Nyiro through Suswa to Mt
Margaret (Fig. 2). The numbers are P-wave velocities in km sx1. Shotpoint locations are shown by stars (Fig. 2). The middle diagram shows the ray
coverage provided by the observed data and the velocity model. The upper diagram shows the fit between the observed traveltimes and those predicted
by ray-trace calculations for the velocity model.
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2001 RAS, GJI 147, 543–561
S. Kenya rift upper crustal structure
velocity of the crystalline crust from an average of 5.9 km sx1
on the western flank to 6.0 km sx1 beneath the rift and
6.3 km sx1 on the eastern flank.
Magadi seismic model
The 1994 KRISP profile crossed the rift valley near Lake Magadi
(Fig. 2). Details of the seismic data and modelling results
are discussed by Birt et al. (1997), Simiyu & Keller (1998)
and Thybo et al. (2000). The most significant aspect of this
model (Fig. 6) is a 4 km thick layer of relatively low-velocity
(5.6 km sx1) material overlying a normal (6.0 km sx1) basement. This layer starts from the eastern rift boundary and
extends to the western limit of the model. Beneath the rift, the
layer extends to a depth of 5.5 km.
The derived geometries of the rift basin show that the
sedimentary and volcanic fill is deepest (3.3 km) to the west.
The fill is composed of a 2.5 km sx1 surface layer underlain by
a 4.2–4.9 km sx1 layer at the base. Below the Magadi shotpoint, the model includes a basement upwarp suggesting a buried
horst block separating two sub-basins. The eastern basin
is shallower than the western one and does not contain the
2.5 km sx1 layer found at the surface in the western basin.
From west to east, the model shows a gradual increase in the
2
WEST
T–X/ 6.00
1
0
549
GRAVITY ANALYSIS AND MODELLING
Our first task in the gravity analysis was to determine the
average density of the topography for the area and to use this
value to calculate Bouguer anomaly values for the data base
that we constructed by combining our new data and published
data provided by Leicester University. We then used these data
to construct anomaly maps and integrated crustal models.
Density determinations
In order to remove the effects of topography, Bouguer and
terrain corrections made routinely as part of gravity data
reduction require knowledge of the average density of the rocks
that constitute the topographic relief of the surveyed area. The
success of these corrections depends on determining the best
Predicted
Observed
EAST
MOR
MAG
KAJ
–1
–2
200
240
Depth (km)
–3
280
320
360
320
360
Rift Valley
0
3
6
9
12
15
200
240
WEST
–3
Depth (km)
0
5.0
MOR
280
Distance (km)
Nguruman
Escarpment
EAST
2.5
MAG
5.6
5.5
4.2
5.8
5.9
3
KAJ
3.9
5.9
4.2
4.9
4.9
6.0
6.0
9
6.2
6.3
5.7
5.6
6
6.1
4.9
Precambrian Basement
6.25
6.3
6.45
12
15
200
240
280
320
360
Distance (km)
Figure 6. Shallow seismic velocity model (lower diagram) along the KRISP 94 EW line G crossing the Lake Magadi area (Fig. 2). The numbers are
P-wave velocities in km sx1. Shotpoint locations are shown by stars (Fig. 2). The middle diagram shows the ray coverage provided by the observed
data and the velocity model. The upper diagram shows the fit between the observed traveltimes and those predicted by ray-trace calculations for the
velocity model.
#
2001 RAS, GJI 147, 543–561
550
S. M. Simiyu and G. R. Keller
average density possible for these rocks. No one density value is
completely satisfactory for a large area, but there are a number
of techniques available to estimate the density of rocks in situ
(e.g. Nettleton 1939). The underlying goal in these techniques is
to minimize the correlation between Bouguer gravity anomalies
and elevation. The approaches we employed were as follows.
(1) Traverses were made across topographic features that
were lithologically similar to their surroundings, and a series
of gravity profiles using various Bouguer reduction densities
were plotted. A density between 2.5r103 and 2.6r103 kg mx3
produced the minimum correlation with topography along
these traverses.
(2) Bouguer gravity values were calculated for different
densities and were plotted against elevation. A value of 2.55r
103 kg mx3 produced the least correlation between Bouguer
gravity and elevation.
(3) More than 2000 density measurements were made on
samples from drill holes to a depth of about 3000 m in the
geothermal areas. Shallow drill holes outside the rift show
densities of less than 2.4r103 kg mx3 up to depths of 2 km,
increasing to more than 2.5r103 kg mx3 as the depth increases.
The mean drill core density value is 2.5r103x2.6r103 kg mx3,
and in situ density measurements carried out by the Kenya
Electricity Generating Company (KenGen) in areas of geothermal exploration show basement density values that range
from 2.58r103x2.88r103 kg mx3.
Based on these results, we chose a density of 2.55r
103 kg mx3 for our reductions. We believe that our density
analysis is a significant contribution, but the use of the traditional
2.67r103 kg mx3 value would not greatly alter our results.
The Bouguer anomaly map (Fig. 7) constructed using this
density value indicates that the regional negative anomaly over
the Kenya rift between 1uN and 3uS is complex and displaced
slightly to the west of the rift valley. This map is dominated by
this regional gravity low due to deep structures, which are not
of interest in this study but which have recently been analysed
by Simiyu & Keller (1997).
In order to delineate major upper crustal features of interest
in this study, a variety of bandpass filters were applied to
the Bouguer anomaly values. The map for wavelengths of
30–150 km (Fig. 8) was the one we found most useful in our
analysis. This map does not include long-wavelength (>150 km)
features associated with deep mantle anomalies under the
Kenya dome (e.g. Green et al. 1991; Achauer 1992; Achauer
et al. 1994; Slack et al. 1994). Wavelengths of less than 30 km
were also removed since such short-wavelength anomalies are
caused by features too local to be of interest in this study. We
feel that this map best depicts the features targeted in this
study. In particular, the axial gravity high is very well defined as
a narrow (y20 km) feature extending from the Lake Baringo
area to Lake Naivasha. In the Baringo area, this high corresponds to the combined effects of a horst block bounding the
Baringo basin to the west in which Precambrian rocks are
exposed at the surface (Swain et al. 1981; Smith & Mosley 1993;
Maguire et al. 1994) and of the anomaly associated with the
Silali volcanic centre. South of Lake Naivasha, this anomaly
widens to a maximum of 40 km at Suswa and then dies out
30 km south of this volcano. At Lake Magadi and further
south, the rift valley occurs within a broad gravity low that is
probably due to one or more of the following effects: (1) more
low-density fill is present than to the north; (2) the rift beneath
Lake Magadi developed on a different crustal type; or (3) the
crust at Lake Magadi has not been magmatically affected in
the same way as in the sectors to the north. At the equator, a
gravity low follows the Nyanza rift eastwards to Lake Victoria,
which implies that it has a simple structure with limited
magmatic modification of the crust. East of the rift valley at
Magadi, a positive gravity anomaly begins at Kajiado (2.1uS,
36.7uE) in an area coincident with outcrops of the Turoka
marble (Shackleton 1986).
Integrated gravity models
Our next step was to extend the seismic results by employing
gravity data for which the geographical coverage was extensive.
Our approach was to model the total Bouguer anomaly, and we
used the models of Simiyu & Keller (1997) to fit the longerwavelength anomalies and then modelled detailed local profiles.
This approach has the advantage of eliminating the need to
perform mathematical or visual regional–residual separation,
which can be a highly subjective process.
A 2.5-D forward modelling program based on the technique
of Cady (1980) was employed. In this study, densities and
depths to particular interfaces were constrained using seismic
and drilling data to the maximum extent possible. The 2.5-D
approach allows bodies of various strike lengths to be included.
For most of the bodies in the models, strike lengths were long
relative to the depth, and therefore, the strike distance was not
very important. However for localized intrusives, such as those
in the Menengai–Suswa area, calculated gravity was sensitive
to the strike lengths.
Five gravity profiles were constructed, and their locations
(Fig. 9) were governed by data coverage and the availability
of seismic and drill hole constraints. The first of the four
E–W profiles is located along y 0.2uS, passing through the
Menengai volcano near Nakuru. The second profile is located
at y 0.5uS, crossing the Mau escarpment to the west and
extending through Naivasha to Mt Kenya. Further south of
Lake Naivasha, the third profile follows a KRISP 85 seismic
line, crossing the rift valley near Suswa. The fourth profile
follows KRISP 94 line G along the Kenya–Tanzania border.
The last profile is a north–south profile along the KRISP 85-90
axial line and extends from Lake Baringo to Lake Magadi. The
analysis of these profiles focuses on upper crustal structure and
the rift fill. An attempt was made to determine and separate the
gravity anomalies related to basement topography from those
related to magmatic intrusions within the basement.
In the seismic models, the uppermost layer was only constrained directly near the shotpoints. However, geological and
drill hole data and gravity anomalies show that this layer of
tuffs, sediments and other volcaniclastics can vary rapidly in
thickness due to faulting and topography. In order to fit both
gravity and seismic data, it was necessary to adjust the thickness
of this layer using drill hole data as constraints. The starting
gravity model in each case was based on the seismic results, drill
hole and geological data, and densities from previous studies
(Henry et al. 1990; Hay et al. 1995; Simiyu & Keller 1997). The
density of the uppermost portion of the crust was allowed to
vary laterally based on typical seismic velocity–density conversions (Barton 1986). It was assumed that since the cover
rocks are several kilometres thick, individual dykes within the
basement will not produce detectable gravity anomalies but
could increase the overall basement density (e.g. Swain 1992).
#
2001 RAS, GJI 147, 543–561
S. Kenya rift upper crustal structure
35˚E
37˚E
36˚E
38˚E
1˚N
-13
34˚E
1˚N
551
0
0
8
-1
50
0
0
-140 -15
-16
-1
-1
70
00
-2
-190
0˚
0˚
-23
0
-17
0
0
-210
1˚S
-180
-200
-23
-160
0
-210
-180
-170
-22
-190
-130
1˚S
-1
-1 20
10
-1
50
-1
80
-1
60
-160
-170
90
-1
-170
0
-17
00
-1
2˚S
-1
80
2˚S
-15
0
0
-17
-160
-150
3˚S
34˚E
35˚E
36˚E
3˚S
38˚E
37˚E
Figure 7. Bouguer gravity anomaly map of the southern Kenya rift using a reduction density of 2.55r103 kg mx3. The contour interval is 5 mGal.
The dashed white lines show the approximate limits of the rift valley.
34˚E
1˚N
35˚E
37˚E
36˚E
38˚E
1˚N
0
10
-50
-10
-1
0
20
0
-10
-20
0˚
0
20
10
-20 10
20
0
-20
-10
0˚
0
-10
0
10
1˚S
-2
20
0
-1
10
10
0
0
-10
0
20
-1
0
1˚S
0
10
10
-10
0
2˚S
0
2˚S
10
20
0
-1
0
0
3˚S
34˚E
35˚E
36˚E
37˚E
3˚S
38˚E
Figure 8. Bandpass filtered gravity map of the southern Kenya rift. Wavelengths passed by the filter were 30–150 km. The contour interval is 5 mGal.
The dashed white lines show the approximate limits of the rift valley.
#
2001 RAS, GJI 147, 543–561
552
S. M. Simiyu and G. R. Keller
o
1 N
Mt. Elgon
M
Kenya
L. Baringo
ELDORET
L. Bogoria
0
o
Equator regional profile
CHE
SOL
Menengai
KISUMU
A
Eburru
1 S
B'
L. Naivasha
ELM
B
VIS
o
A'
Mt. Kenya
D
Olkaria
OLO
EWA
Longonot
C
C'
MAR
NTU
SUS
Mt Suswa
NAIROBI
MOR
L. Magadi
o
2 S
Tanzania
KAJ
D'
KIB
M'
Southern regional profile
o
34 E
CHN
L Natron
o
o
36 E
38 E
Figure 9. Index map showing locations of the gravity profiles that were modelled in this study (A–Ak to M–Mk). Shotpoints from seismic profiles are
shown as stars (Fig. 2) to indicate the close correspondence between the seismic and gravity models. Also shown are locations of the regional gravity
profiles that were modelled by Simiyu & Keller (1997).
In order to emphasize crustal features, the density contrast
between the rifted and cratonic mantle is large relative to what
the seismic velocities suggest (Ravat et al. 1999), so that the
depth of the models could be limited to 100 km. By extending
the models to 150 km or more based on the teleseismic results
(e.g. Achauer et al. 1994; Slack et al. 1994), this density contrast
could be reduced. In the seismic models, a depth of 0 indicates
sea level, but in the gravity models a depth of 0 corresponds to
the lowest elevation of the rift valley floor on a particular
profile.
Profile A–Ak (Figs 9 and 10) is located near the triple
junction between the northern Kenya, Nyanza and southern
Kenya rifts. This profile and those that follow sample large
topographic variations across the rift in which the surface
rocks have widely varying densities based on surface and drill
hole density measurements. The profile starts 10 km west of
the surface contact between the Quaternary volcanics and the
Mozambique belt rocks. Further east it crosses upper Miocene
phonolites and then lower Miocene mixed lavas. At the 90 km
point, the profile crosses into the rift valley at the northern end
of the Mau escarpment and fault zone (Fig. 1). East of this
fault system, the profile crosses the rift valley and mainly
Pleistocene to Recent pyroclastics. Near the eastern flank
of the rift valley, the profile crosses an area with 1–2 km of
low-density volcanic deposits and wind-deposited lapilli tuffs
(Williams 1978). A hole drilled to a depth of 475 m in this area
encountered volcanic soils and tuffs (Mwaura 1990). Further
east, the profile crosses the Mount Kenya volcanic edifice and a
gravity low that was modelled to be the result of at least 2.5 km
of Pleistocene trachytes and tuffs.
The gravity high at Menengai, with a 40 mGal amplitude
and 35 km width was modelled by a 20 km wide intrusive body
at a depth of 4 km (Fig. 10). This anomaly is bounded by
gravity gradients that we interpret as being caused by sharp
contacts between the dense intrusion and thick low-density
tuff and ash. Henry et al. (1990) noted a traveltime advance at
Menengai that suggests a fairly large high-velocity body, and
the seismic model for the axial line (Fig. 4) indicates high
basement velocities in this area.
Profile B–Bk (Figs 9 and 11) extends from Mau Narok across
the Mau fault zone, the Olkaria geothermal field south of Lake
Naivasha, the Aberdare Range and the southern slopes of
Mt Kenya. West of the rift valley, a 23 mGal amplitude gravity
low is located on the Mau escarpment and is modelled as
being the result of a 2.5 km thick layer of Mau tuffs with
a density of 2.2r103 kg mx3 (Fig. 11). Nyblade & Pollack
(1992) argue that the gravity lows west of the rift valley in this
region could reflect the presence of local crustal thickening due
to the suture contact between the Archaean Tanzania craton
and the Mozambique belt. This particular anomaly is only
about 30 km wide, but Tesha et al. (1997) found wider negative
anomalies to the south that correlated with this possible suture.
#
2001 RAS, GJI 147, 543–561
S. Kenya rift upper crustal structure
A
A'
–50
Bouguer Anomaly
(mGals)
553
Tie
to
Axial Line
Calculated
Observed
–100
Mt. Kenya
Menengai
–150
MFZ
–200
KFZ
–250
0
WEST
50
100
150
200
250
2.4 2.90
2.3
2.74
2.3
300
350
EAST
300
350
2.4
Depth (km)
0
2.70
2.85
20
40
2.70
2.85
2.85
3.05
60
3.12
3.26
3.26
80
100
0
50
100
150
200
250
Distance (km)
Figure 10. Integrated gravity model of Earth structure along profile A–Ak, which crosses the rift valley at Menengai volcano (Fig. 9). Numbers
indicate density valuesr103 kg mx3. In order to emphasize crustal features, the density contrast in the mantle is very large to limit the vertical extent
of the model. MFZ: Mau fault zone; KFZ: Kinangop fault zone.
gravity high beneath the Olkaria geothermal field has been
modelled as being the result of a 9 km wide intrusion that rises
to within 3.5–4.0 km of the Earth’s surface.
Profile C–Ck (Figs 9 and 12) starts at Ewaso Nyiro in the
west, crosses the Mau escarpment, passes through Suswa
caldera and extends to Thika, which is 30 km north of Nairobi.
A unique feature of this profile is that gravity values on the
western flank of the rift valley are uniform (about x200 mGal)
and lower than on the eastern side (Fig. 12). Local gravity lows
coincide with exposures of tuffs that have densities as low
as 2.0r103 kg mx3. East of Ewaso Nyiro, a negative anomaly
can be attributed to 1.5 km of tuffs where a well drilled nearby
encountered continuous tuffs and limited trachyte intercalation
to the bottom at 597 m (Mwaura 1990).
The steep east rift valley boundary fault with a throw of
more than 3.5 km is located at model coordinate 120 km. The
thickness of volcanic fill in the rift valley decreases eastwards
from more than 3.0 km on the western margin against the
Mau fault zone to 1.5 km against the Kinangop fault zone
at the eastern margin. Wells drilled to a depth of about 3 km
on the western part of the rift floor have never reached the
Precambrian basement or encountered any form of sedimentary
rocks (Roessner & Strecker 1997). The thickness of the volcanics
suggests that the master fault for the Naivasha basin region is
on the western side of the rift valley. On the eastern margin, the
Kinangop tuffs form a 1.0 km thick layer (Roessner & Strecker
1997). Further east, south of Mt Kenya, a gravity low has been
modelled as being due to a 2.0 km thick layer of volcanics. The
B'
B
Bouguer Anomaly
(mGals)
–50
Tie
to
Axial Line
Calculated
Observed
–100
Mt. Kenya
Olkaria volcanic
field
–150
KFZ
MFZ
–200
–250
0
WEST
50
2.3
2.3
Depth (km)
100
150
2.90
0
2.4
250
300
2.3
350
EAST
2.4
2.70
2.74
2.70
2.85
20
200
2.85
40
60
3.05
3.26
3.26
3.12
80
100
0
50
100
150
200
250
300
350
Distance (km)
Figure 11. Integrated gravity model of Earth structure along profile B–Bk, which crosses the rift valley at the Olkaria volcanic field (Fig. 9). Numbers
indicate density valuesr103 kg mx3. In order to emphasize crustal features, the density contrast in the mantle is very large to limit the vertical extent
of the model. MFZ: Mau fault zone; KFZ: Kinangop fault zone.
#
2001 RAS, GJI 147, 543–561
554
S. M. Simiyu and G. R. Keller
C
–50
C'
Tie
to
Axial Line
Bouguer Anomaly
(mGals)
Calculated
Observed
–100
Nairobi
Suswa
–150
–200
–250
0
WEST
50
2.3
Depth (km)
0
100
2.3
150
200
300
2.90
2.3
2.74
2.4
350
EAST
2.4
2.70
2.68
20
250
2.85
2.85
40
3.05
60
3.26
3.26
3.12
80
100
0
50
100
150
200
250
300
350
Distance (km)
Figure 12. Integrated gravity model of Earth structure along profile C–Ck, which crosses the rift valley at Suswa volcano (Fig. 9). Numbers indicate
density valuesr103 kg mx3. In order to emphasize crustal features, the density contrast in the mantle is very large to limit the vertical extent of the
model.
On the eastern shoulder of the rift in the Nairobi–Thika area,
a local low in the much broader regional positive anomaly
coincides with low-density surface volcanics that thicken from
longitude 37.20uE towards the rift. Water wells drilled in the
area east of Nairobi (Mwaura 1990) have reached Mozambique
gneissic basement at a depth of 530 m, after going through
Kerichwa Valley tuffs (80 m thick), Nairobi phonolites
(100 m thick), Athi tuffs (320 m thick) and Kapiti phonolites
(30 m thick). All these lavas and tuffs have densities less than
2.6r103 kg mx3. To the west of Nairobi, on the escarpment
bounding the rift valley, the south Kinangop tuffs form a layer
about 1.5 km thick. Within the rift, the positive anomaly is
60 km wide and has an amplitude of 30 mGal. The rift valley
area has been modelled as a graben whose fill varies in thickness
from 2.7 km in the west to 1.5 km in the east and whose density
is 2.4r103 kg mx3. Beneath the Suswa caldera, the gravity
high is due to an 18 km wide body at a depth of about 2.5 km
with a density of 2.90r103 kg mx3. A difference between
this model and models for profiles to the north (Menengai
and Naivasha) is that the Tanzania craton is present under
a significant portion of the western flank of the rift valley.
Thus, based on the work of Tesha et al. (1997), which showed
2.68r103 kg mx3 to be a good choice for the density of the
upper crust of the Tanzania craton, we used this lower value to
model the upper crust across the western flank.
Profile D–Dk (Figs 9 and 13) is an extension of the model
presented by Simiyu & Keller (1998) and crosses the southern
portion of the rift valley before it splays out in the area of Lake
Natron in northern Tanzania (Fig. 1). The location of this
profile away from the thick volcanic pile associated with the
Kenya dome makes it more favourable for analysing the details
of basement structures underlying the Kenya dome that are
otherwise obscured by these volcanics. The profile starts within
the Archaean Tanzanian craton east of Lake Victoria. At the
Oloololo escarpment (model coordinate 100 km), there is a
geological, structural and topographic boundary separating
surface exposures of the Archaean Tanzania craton and the
Neo-Proterozoic Mozambique belt (Fig. 13). East of the escarpment, the profile crosses the flat Mara plains, which are covered
by flood trachytes and phonolites of the Isuria Group. Further
east the profile crosses the Loita Hills and the major bounding
fault of the western part of the rift valley, the Nguruman
escarpment (Crossley 1979).
The Loita Hills and the Nguruman escarpment are composed of massive, highly fractured quartzite and gneisses of
Proterozoic age (Crossley 1979). These quartzites are thought
to extend west to the craton–mobile belt boundary beneath the
Mara plains (Smith 1994). The profile crosses the rift at coordinate 250 km, where the rift valley appears to have developed
within the mobile belt. East of the rift around Kajiado, the
profile crosses a band of marble surrounded by some discontinuous quartzites and gneisses (Shackleton 1986). Further
east of Kajiado, there are minor layers of volcanics from the
Chyulu Hills volcanic complex.
Unlike the three profiles to the north, there is only a small
short-wavelength gravity high in the rift valley area. This
anomaly is flanked by two minima with the western one having
the larger amplitude. Our modelling of these anomalies shows
that the rift fill is about 3.5 km thick against the Nguruman
escarpment to the west and only 2.0 km thick to the east. The
basin is divided into two parts by a horst structure that runs
along its axis, thus creating the small gravity high. The
crystalline upper crust on the western flank has been modelled
in two parts: an upper 4 km thick layer with a low density
of 2.60r103 kg mx3 (VP=5.6 km sx1; Simiyu & Keller 1998)
and a lower layer that is 11 km thick with a typical basement
density (2.68r103 kg mx3; VP=5.9 km sx1; Simiyu & Keller
1998). East of the rift, the upper crust has been modelled
with a 2.71r103 kg mx3 density based on velocity values
(VP=6.3 km sx1; Birt et al. 1997).
The lower crust has been modelled using the KRISP 94
seismic results (Birt et al. 1997) as constraints. From west to east,
there is little crustal thinning under the rift valley. However,
east of the rift, there is crustal thickening from y35 km
#
2001 RAS, GJI 147, 543–561
S. Kenya rift upper crustal structure
Bouguer Anomaly
(mGals)
D
D'
–50
Calculated
Observed
–100
Oloololo
Escarpment
–150
Tie
to
Axial Line
–200
–250
WEST
Tanzania Craton
Reworked Zone
Lake
Magadi
EAST
Rift Valley
Mozambique Belt
2.60
2.90
0
2.71
2.68
Depth (km)
20
555
2.85
2.86
2.85
2.88
40
60
3.26
3.26
3.15
80
100
Depth (km)
50
250
NE
Lake
2.4
2.5 Morijo
Magadi
150
0
2.60
2.65
Kajiado
2.71
2.68
2.73
2.74
15
200
450
2.60
5
10
350
250
2.5
2.75
300
350
Distance (km)
Figure 13. Integrated gravity model of Earth structure along profile D–Dk, which crosses the rift valley at Lake Magadi following KRISP 94 seismic
line G (Fig. 9). Numbers indicate density valuesr103 kg mx3. In order to emphasize crustal features, the density contrast in the mantle is very large in
order to limit the vertical extent of the model. A detailed upper crustal model (15 km depth) for the region outlined by the dashed box in the upper
diagram is shown in the lower diagram. NE: Nguruman Escarpment.
beneath the rift to about 40 km at the end of the profile. West
of the rift, the lower crust is modelled with a density of
2.85 r 103 kg mx3 (VP = 6.7 km sx1; Birt et al. 1997). The
lower crustal densities increase from 2.86 r 103 kg mx3
(VP = 6.9 km sx1; Birt et al. 1997) beneath the rift to
2.88 r 103 kg mx3 (VP = 7.2 km sx1; Novak et al. 1997)
beneath the eastern flank. Using mantle densities suggested
by Hay et al. (1995), the broad gravity low associated with
deep structure was modelled primarily by a mantle anomaly
(3.15r103 kg mx3) beneath the rift valley. However, west of
the rift valley, the combination of the reduced density of the
upper crust and slight crustal thickening, probably due to
Mozambique belt collision tectonics, also explains some of the
regional minimum.
The axial profile (M–Mk, Figs 9 and 14) was constructed
along the KRISP 85 and KRISP 90 seismic lines and extends
from Lake Baringo at 1uN southwards to Lake Magadi. The
significance of this profile is that it follows the axis of the rift
and the gravity axial high as closely as possible and that it ties
all of the E–W profiles together. The profile starts at Lake
Baringo within the Baringo–Bogoria sub-basin characterized
by recent volcanics and lake sediments (Mugisha et al. 1997).
South of Lake Bogoria at model coordinate 50 km (Fig. 14),
the profile crosses the proposed Metkei–Marmanet accommodation zone, which is also within the Nyangea–Athi–Ikutha
shear zone axis (Smith 1994). South of this accommodation zone,
the Nakuru–Naivasha sub-basin extends to Suswa. Seismic
#
2001 RAS, GJI 147, 543–561
data (Henry et al. 1990; Mechie et al. 1994) show this basin
to have the greatest thickness of rift fill (6 km) beneath the
Lake Naivasha area (Fig. 4). The profile crosses areas with
deep drill hole data between Lake Elmenteita and Suswa that
are underlain by a large thickness of tuffs, trachytes, rhyolites and
phonolites. South of Suswa, there is another proposed accommodation zone (Bosworth 1987), and then the profile passes
into the Magadi–Natron sub-basin. In this basin, the profile is
within the Aswa–Nandi–Loita shear zone, which is a highly
faulted and fractured zone of massive quartzitic basement rocks
covered by a thin layer of rift-related volcanics (Smith 1994).
The main features of the profile (Fig. 14) are the broad
negative Bouguer anomaly with a wavelength >350 km and
an amplitude of >50 mGal and a series of positive gravity
anomalies superimposed on it. These gravity highs (Fig. 8) are
aligned along the axis of the rift at the Menengai, Eburru,
Olkaria and Suswa volcanoes. Following Simiyu & Keller
(1997) and the seismic constraints, the broad gravity low was
modelled as resulting from two factors: (1) the mantle is less
dense at the apex of the Kenya dome because it is hotter
compared to the northern and southern sectors, and (2) a large
thickness (1–5 km) of low-density (2.3r103 kg mx3) tuffs and
lavas fill the graben associated with the rift valley. The gravity
highs were modelled as resulting from a number of volcanic
centres underlain by discrete mafic bodies with a density of
2.90r103 kg mx3. These bodies are interpreted to represent
upper-level mafic plutons (Omenda 1998).
556
S. M. Simiyu and G. R. Keller
M'
M
Calculated
Observed
Equator regional
profile
–100
(mGals)
Bouguer Anomaly
0
C-C'
B-B'
A-A'
D-D'
–200
SOUTH
NORTH
0
Menengai
Lake Baringo
0
Depth (km)
200
100
2.70
300
Olkaria
400
Lake Magadi
2.90
Suswa
2.74
2.70
2.70
2.85
20
40
3.05
60
3.15
3.15
3.12
80
100
0
200
Distance (km)
100
300
400
Figure 14. Integrated gravity model of Earth structure along the axis of the rift valley (M–Mk) from Lake Baringo to Lake Natron in northern
Tanzania (Fig. 9). Numbers indicate density valuesr103 kg mx3. In order to emphasize crustal features, the density contrast in the mantle is very
large to limit the vertical extent of the model. Ties to the E–W profiles are shown.
pretation of geological, seismic refraction, gravity and drill hole
data. Key points addressed by our analysis are discussed below.
DISCUSSION
The results of this study show that the southern Kenya rift has
a crustal structure that is equally as complex as its tectonic
history. To synthesize our results and illustrate the main
features we have constructed three schematic geological crosssections (Figs 15–17), and one interpreted map of the major units
that crop out or subcrop beneath the volcanic cover (Fig. 18)
was constructed. These diagrams represent our integrated inter-
Elev. (km)
WEST
4
2
Rift basin geometry
Because of the deep drill holes associated with geothermal
exploration, the best-constrained cross-section can be drawn
through the Olkaria region (Fig. 15). In this area, Bosworth
et al. (1986) and Bosworth (1987) attributed the sinuous course
EAST
Rift Valley
DH
Mau
Escarpment
Olkaria
Kinangop–Kijabe plateau
OW305 OW709 OW30
OW704
OW–304
DH
DH
DH
DH
D
Depth (km)
0
SL
2
4
Upper crust
Mozambique Belt
6
8
10
12
14
0
10
20
30
40
50
60
70
80
90
100
Distance (km)
Pleistocene–Recent
pyroclastics and sediments
Pliocene Mau and
Kinangop tuffs
Precambrian metamorphic
basement (upper crust)
Plateau trachytes
intercalated with tuffs
Miocene lavas
(thickness estimated)
Intrusive body
OW–304
Deep geothermal drillholes
DH
Shallow drillhole
Figure 15. An interpreted geological cross-section across the rift valley in the Lake Naivasha–Mt Longonot area.
#
2001 RAS, GJI 147, 543–561
Elev. (km)
S. Kenya rift upper crustal structure
WEST
3
EAST
Nguruman Rift Valley
Escarpment DH
L. Magadi
DH
557
DH
Kajiado
DH
DH
(SL) 0
Depth (km)
3
Zone of tectonic reworking
Mozambique belt
6
9
12
Tanzanian Craton?
15
120
160
200
240
280
320
360
Distance (km)
Recent volcaniclastics,
and sediments
Crystalline upper crust
(Precambrian basement)
Pliocene / Miocene
volcanics and sediments
Pleistocene–Recent
pyroclastics
Mozambique belt
metasediments
Archean crust
Highly fractured quartzitic
Mozambique belt rocks
Plateau trachytes,
DH Drillhole
Figure 16. An interpreted geological cross-section across the rift valley in the Lake Magadi area (modified from Simiyu & Keller 1998).
crossing the rift valley north of Suswa volcano (Fig. 5). The
upper surface of the Pliocene Mau and Kinangop volcanics is
exposed along the western boundary of the rift valley, where it
is offset by more than 1000 m by the Mau fault zone (Roessner
& Strecker 1997). The depth of the Pliocene volcanics beneath
the inner graben in drill holes in the Olkaria geothermal field
shows pronounced subsidence of this unit to a depth of 1500 m.
The overall implication of this result is that the southern and
central Kenya rift valley contains a series of asymmetric basins
that are all bounded to the west by master faults. In addition,
intrarift horst blocks occur near the axis of the rift valley in
several locations. To the north of Menengai, the Kerio basin
Elev. (km)
of the rift valley (Fig. 1) to the en echelon arrangement of
faults defining a series of asymmetric half-grabens (basins)
that change polarity at accommodation zones. In their model,
extension is accommodated along major listric detachments of
lithospheric extent that alternate in polarity along the rift. They
suggested that the major detachment fault in the Nakuru–
Naivasha sub-basin is the Sattima fault (Fig. 1), which lies along
the eastern margin of the rift valley, the Aberdare Range. Our
results show the graben to be deepest in front of the western
margin of the rift valley at the Mau escarpment (Fig. 15). This
rift basin geometry is also clearly shown on the cross-section of
the Suswa area (Fig. 16) and the seismic model for the profile
NORTH
3 L. Baringo
Menengai
Eburru
DH
DH
Olkaria
SOUTH
L. Magadi
Suswa
DH
DH
DH
(SL) 0
Depth (km)
3
6
9
12
15
0
40
80
120
160
240
200
280
Distance (km)
Pleistocene–Recent
pyroclastics and sediments
Miocene lavas
(thickness estimated)
Intrusive bodies in
the upper crust
Plateau trachytes
intercalated with tuffs
Highly fractured quartzitic
Precambrian basement
Mid–lower crust
Pliocene Mau and
Kinangop tuffs
Metamorphic granitoid
Precambrian basement
DH
Figure 17. An interpreted geological cross-section along the rift axis from Lake Baringo to Lake Magadi.
#
2001 RAS, GJI 147, 543–561
Drillhole
558
S. M. Simiyu and G. R. Keller
o
o
34 E
Mt Elgon
38 E
?
NA
L
?
?
I
AN
ELDORET
L. Baringo
L. Bogoria
?
?
KISUMU
Aberdare
Range
o
Mt Kenya
Menengai
L. Victoria
0
?
L. Naivasha
?
?
Suswa
NA
I
?
NAIROBI
L. Magadi
o
yu
ng
L
e
Archean
Tanzania craton
? Unsure of the
buried geology
Off—axismajor volcanoes
lu
Ra
AN
L. Natron
Volcanic ranges
2S
Ch
Tanzania
Craton
Kilimanjaro
Zone with thrusts, reworked
craton and mobile belt rocks
Mobile belt gneisses and
other metasediments
Rift axis caldera volcanoes
Outline of the apex
the Kenya dome
NIA & ANL - Nyangea—Athi-Ikutha &Aswa—Nandi—Loita shear zones
Figure 18. Interpreted map of the major units that crop out or
subcrop beneath the volcanic cover in the southern rift valley region.
Boundaries are based on gravity anomaly characteristics and geophysical modelling together with drilling and geological information
(Crossley 1979; Shackleton 1986; Smith & Mosley 1993; Maguire et al.
1994; Smith 1994; Stern 1994).
west of Lake Baringo (Fig. 1) is bounded by a master fault on
the west and by a horst block on the east that is, in turn,
bounded on the east by the relatively shallow Baringo basin
(Swain et al. 1981; Maguire et al. 1994). Thus, there is a lack
of polarity reversals in rift basins for a distance of about
300 km along the rift valley. This observation contrasts with the
western branch of the East African rift system, where polarity
reversals are well documented (Rosendahl 1987).
The axial gravity maxima and magmatic modification of
the crust
Our new data and improved density analysis produce a
particularly clear picture of the axial gravity high (Fig. 8). As
summarized by Swain et al. (1994), early workers interpreted
the axial gravity high in the rift valley to be the result of a
crustal-scale dyke. The results of KRISP have shown that no
continuous feature of this magnitude exists. However, this is
not to say that magmatic modification of the crust is negligible.
Obviously, the volume and nature of volcanic activity in the
region require considerable modification of the crust. However,
the nature and extent of this modification is an important
issue. As discussed by Keller et al. (1994b), Hay et al. (1995)
and Thybo et al. (2000), a significant amount of underplating /
intrusion near the base of the crust (Moho) has occurred under
the region between Lake Baringo and Lake Magadi (Fig. 14).
However, this material extends across the entire rift valley and
a portion of the flanks and is the cause of only long-wavelength
gravity anomalies (Figs 10–13). Swain (1992) showed that the
early KRISP results suggested some densification (distributed
intrusions) of the upper crust in the vicinity of the rift valley
axis, and as we have shown, subsequent KRISP data are
consistent with this interpretation.
However, a small level of densification is not enough to
explain the axial gravity high, and our analysis and those of
Swain (1992), Maguire et al. (1994) and Simiyu & Keller (1998)
show that intrarift structures (horst blocks) play a significant
role in creating the axial gravity high. Details revealed by the
new gravity data presented here show that the magnitude of the
axial high changes significantly along its extent. Our seismic
and gravity modelling show that a series of local gravity highs
are coincident with the major volcanic centres (Figs 8 and 14).
The individual anomalies associated with each centre sometimes have half-wavelengths greater than their distance apart so
that the anomalies merge. Away from these centres, the axial
high can be explained by the combination of an intrarift horst
and modest magmatic modification of the upper crust. Thus,
the previous perception of a simple continuous axial positive
gravity anomaly was the result of sparse data and the alignment
of closely spaced volcanic centres across which gravity profiles
were constructed.
Apart from the intrusions at volcanic centres, the whole
crustal section of the area from Elmenteita to Suswa shows
evidence for increased density that may result from magmatic
injection by dyke swarms as evidenced by drill hole data. This
has had the effect of raising the upper crustal density by about
1–2 per cent (from 2.70 to 2.74r103 kg mx3, Fig. 14), thus
effectively counteracting the effect of the low-density rift fill.
This interpretation is consistent with a southward increase in
gravity values from Elmenteita to Suswa even though the area
is covered by 6 km of relatively low-density volcanic deposits
dominated by ash-flow tuffs, air-fall tuffs and minor phonolites
with trachytes.
Crustal genesis and the role of older structures
Analysis of the gravity maps and modelled seismic and gravity
profiles shows different signatures that have led us to classify
the crust in the region into broad zones: the Archaean craton,
the mobile belt with reworked craton and the Proterozoic
mobile belt (Fig. 18). Regional relationships (Smith 1994) and
the seismic and gravity models (Fig. 16 and Birt et al. 1997)
show that the mobile belt /reworked craton layer is about 4 km
thick and ends near the western margin of the rift valley.
Stern (1994) proposed an alternative Precambrian collision
model that suggests that more than one orogenic event involving
smaller land masses occurred in the southern Kenya area. This
suggestion implies that there is not a single suture contact
zone between the Archaean craton and the mobile belt but
rather a wide zone of imbricate thrusting and suturing. Gravity
profiles constructed during this study do not clearly show the
characteristic suture signature at the surface contact between
the Archaean craton and the mobile belt that is seen in
northern Tanzania (Tesha et al. 1997) and north of the equator
(Nyblade & Pollack 1992). This may imply that the reworked
and buried craton of Smith & Mosley (1993) is actually part of
the small land masses that were involved in the multicollision
events of Stern (1994).
#
2001 RAS, GJI 147, 543–561
S. Kenya rift upper crustal structure
The Kenya rift occurs in an area between contrasting
lithosphere types (craton and mobile belt) that are expected to
have different mechanical strengths. In the event of applied
stresses, failure is likely to occur at such a boundary. Another
source of weakness is the likelihood that during Mozambique
belt collisional tectonics, there was crustal thickening near
the suture as evidenced in gravity studies (Nyblade & Pollack
1992; Tesha et al. 1997). Smith & Mosley (1993) and Hetzel
& Strecker (1994) have suggested that the overall rift trend
closely follows large- and small-scale basement structures. This
study provides further evidence that localization of rifting was
controlled by the existing contact between the Archaean craton
and the Proterozoic mobile belt.
The Kenya dome gravity low
The Kenya dome is located at the confluence of structures that
resulted from a variety of tectonic events. The area has been
affected by (1) thrusting related to one or more collisional events
during the formation of the Mozambique belt, (2) deformation
associated with its location between the Nyangea–Athi–Ikutha
shear zone to the north and the Aswa–Nandi–Loita shear zone
to the south (Fig. 18), which are major NW–SE-trending zones
of ductile shear and thrusting in the underlying Precambrian
basement, and (3) sub-Miocene uplift and doming followed by
magmatism and N–S-trending rift faulting. All these events
must have some influence on the gravity anomalies observed
today.
The Bouguer gravity map, bandpass filtered maps and the
modelled gravity and seismic profiles show that the gravity
low associated with the Kenya dome is the result of several
features. At the apex of the Kenya dome, Cenozoic low-density
tuffaceous volcanic rocks reach thicknesses of 2 km on the rift
flanks (Baker et al. 1988) and 6 km within the rift valley (Henry
et al. 1990). Thus some of the relief of the dome is in fact due
to volcanic construction. The bandpass filtered map (Fig. 8)
shows that part of the gravity low is contributed by low-density
material on the Mau escarpment and in the Aberdare Range.
This low-density material is associated with a number of small
volcanic centres, but since they are closely spaced, their anomalies
merge and appear as a single long-wavelength anomaly. An
example of this phenomenon can be seen on the eastern rift
flank where the Aberdare Range volcanic centres form a series
of closely linked silicic volcanoes (Smith 1994). The overall
effect of these is to create a long-wavelength gravity low that
merges with the adjacent low due to the rift fill.
Crustal thickening by about 5 km due partly to underplating /intrusion near the Moho (Keller et al. 1994b; Hay et al.
1995) and pre-existing thickened crust caused by Precambrian
collisional tectonics (Nyblade & Pollack 1992; Tesha et al.
1997) also contribute to the gravity low. On a regional basis,
crustal thickening provides Airy-type (Banks & Swain 1978)
compensation for the dome.
In the area west of Lake Magadi, the uppermost Precambrian
crust is a relatively low-density layer that was probably
structurally emplaced, adding to the crust from above. The
extent of this layer shown in Fig. 18 is based on the persistence
of a characteristic gravity minimum that also contributes to the
Kenya dome gravity low.
The low-density mantle beneath the apex of the dome is the
major contributor to the long-wavelength component of the
gravity low. Simiyu & Keller (1997) showed that the uppermost
#
2001 RAS, GJI 147, 543–561
559
mantle density varies along the rift by 0.03r103 kg mx3,
decreasing from the Lake Baringo region to a minimum at the
apex of the Kenya dome (Lake Naivasha). South of Naivasha,
the density increases towards Lake Magadi (Fig. 14). The
implication of this density variation is that the mantle is hotter
beneath the apex of the Kenya dome and provides part of the
isostatic compensation for the Kenya dome via a Pratt-type
mechanism (Banks & Swain 1978).
Crustal symmetry with respect to the Kenya dome
The KRISP 90 axial profile showed a dramatic increase in
crustal thickness from the north (Lake Turkana) to the Kenya
dome (Mechie et al. 1994) that is the result of the thickening
of a high-velocity layer at the base of the crust. This layer
was interpreted to be due to magmatic underplating /intrusion
at the base of the crust (Keller et al. 1994b; Hay et al. 1995). If
the crustal thickening at the apex of the dome is only due to
underplating as a result of recent magmatic activity, then the
crust should also thin southwards away from the dome.
However, this study and the analysis of Birt et al. (1997) and
Simiyu & Keller (1997) show that the crustal thickness does not
decrease greatly to the south away from the apex of the dome
and that the low-velocity /density mantle does not reach the
base of the crust as happens to the north. However, a highvelocity layer at the base of the crust extends for more than
200 km outside the rift in the Lake Magadi area (Birt et al.
1997). This suggests that the relatively thick crust beneath the
Kenya dome may be only partly due to underplating of highvelocity material. We believe that the most likely explanation
for areas of thick crust not due to underplating is pre-existing
variations in crustal thickness resulting from the formation of
the Mozambique belt. These pre-existing zones of thickened
crust may have created an ideal weak zone for rift propagation.
CONCLUSIONS
Our integrated interpretation of the crustal structure of the
southern Kenya rift shows that the graben master faults are
all located along the western flank of the rift valley, and there
are no half-graben polarity reversals. There is no symmetry
in crustal thickness variations with respect to the apex of the
Kenya dome, and the existing crustal thickness is in part due to
pre-rift crustal type and thickness in addition to underplating /
intrusion and extension. The pre-existing lithospheric contrast
between the Archaean and Proterozoic basement terranes
played a significant role in the location and structural geometry
of the rift. Low upper crustal velocities and densities observed
under the rift flank west of Magadi probably represent material
added to the top of the Archaean Tanzania craton by tectonic
transport. Magmatic modification of the upper crust is modest
except near the major Quaternary volcanic centres. An intrarift
horst block that extends along at least most of the rift axis and
this series of volcanic centres create gravity maxima that
together produce the axial gravity high.
ACKNOWLEDGMENTS
We would like to acknowledge the help and input provided by
our colleagues in the KRISP Working group. We also wish to
thank Dr Jeff Karson of the Division of Earth and Ocean
Science for his constructive review of the original manuscript as
560
S. M. Simiyu and G. R. Keller
well as the reviewers of this paper, who provided many helpful
and constructive comments. The Kenya Electricity Generating
Company helped support the gravity surveying. The seismic
data analysed in this study were largely gathered using
equipment provided by the Incorporated Research Institutions
for Seismology. Major funding was provided by NSF grant
EAR-9316868. Additional support was provided by Conoco
Inc and NASA through its support of the Pan-American
Center for Earth and Environmental Studies at the University
of Texas at El Paso.
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