Journal of Petrology

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Journal of Petrology Volume 42 Number 5 Pages 877-900 2001
© Oxford University Press 2001

Plume–Lithosphere Interactions in the Generation of the Basalts of the Kenya Rift, East Africa

R. MACDONALD^1,*, N. W. ROGERS², J. G. FITTON³, S. BLACK⁴ and M. SMITH⁵

¹ENVIRONMENTAL SCIENCE DIVISION, IENS, LANCASTER UNIVERSITY, LANCASTER LA1 4YQ, UK
²DEPARTMENT OF EARTH SCIENCES, THE OPEN UNIVERSITY, MILTON KEYNES MK7 6AA, UK
³DEPARTMENT OF GEOLOGY AND GEOPHYSICS, UNIVERSITY OF EDINBURGH, WEST MAINS ROAD, EDINBURGH EH9 3JW, UK
⁴PRIS, UNIVERSITY OF READING, WHITEKNIGHTS, READING RG6 2AB, UK
⁵BRITISH GEOLOGICAL SURVEY, MURCHISON HOUSE, WEST MAINS ROAD, EDINBURGH EH9 3LA, UK

Received January 1, 2000; Revised typescript accepted July 19, 2000

	ABSTRACT

TOP ABSTRACT INTRODUCTION GEOLOGICAL SUMMARY ANALYTICAL TECHNIQUES NOMENCLATURE AND DISTRIBUTION COMPOSITIONAL DIFFERENCES... BASALT COMPOSITION VS TIME:... COMPARISONS WITH OCEAN ISLAND... CRUSTAL CONTAMINATION MANTLE HETEROGENEITY MANTLE END-MEMBERS CONCLUSIONS APPENDIX : SAMPLE LOCALITY... REFERENCES

 
Major and trace element and Sr–Nd–Pb isotopic data for mafic volcanic rocks are used to assess the number of mantle plumes contributing to the Tertiary–Holocene magmatism of the Kenya Rift Valley, current estimates of which vary from none to three. Rocks ranging in composition from nephelinite to hypersthene-normative basalt have been sampled from three lithospheric zones: the Tanzanian craton, the craton margin reworked during the late Proterozoic, and the Mozambique mobile belt. The magmas are interpreted as the products of variable degrees of partial melting within the spinel–garnet peridotite transition zone. Trace element and isotopic compositions from all three zones are broadly similar to those of oceanic island basalts, but there is considerable compositional variation, which is related to a strong overprint from the lithosphere on plume-derived melts. Sr and Nd isotopic ratios provide the only clear distinction between magmatic rocks from the three lithospheric domains. Within each setting, mafic magmatism has tended to become less silica undersaturated with time, and at any one locality magmatism has migrated towards the centre of the rift. Magmas may have formed as a result of the infiltration of plume-derived melts into the base of the lithosphere. The extent of interaction of inferred plume melts with the lithosphere has not varied systematically in time or space. The plume component appears to be similar to the source of oceanic island basalts.

KEY WORDS: Kenya Rift Valley; mantle plumes; geochemistry; metasomatism

	INTRODUCTION

TOP ABSTRACT INTRODUCTION GEOLOGICAL SUMMARY ANALYTICAL TECHNIQUES NOMENCLATURE AND DISTRIBUTION COMPOSITIONAL DIFFERENCES... BASALT COMPOSITION VS TIME:... COMPARISONS WITH OCEAN ISLAND... CRUSTAL CONTAMINATION MANTLE HETEROGENEITY MANTLE END-MEMBERS CONCLUSIONS APPENDIX : SAMPLE LOCALITY... REFERENCES

 
Magmatism along the Kenyan and Ethiopian Rifts, the associated flood basalts of northern Ethiopia and the volcanism of the Afar Depression (collectively the East African Rift System) represent one of the largest, currently active, continental igneous provinces (Fig. 1). In the Kenya sector of the East African Rift System alone, an estimated 924 000 km³ of mafic magma, including underplated material, has been generated in the past 30 my (Latin et al., 1993). Such volumes are far too large to have been produced solely within (anhydrous) lithosphere, especially given the evidence for small melt fractions provided by rare earth element (REE) modelling; the melt zone must actively have been fed from upwelling asthenosphere (Latin et al., 1993). Furthermore, estimated extension rates across the Kenya rift are probably too small (ß-factors <1·7; Hendrie et al., 1994; Mechie et al., 1997) to explain the magma volumes and the protracted duration of magmatism (35 my) without the added effect of elevated mantle temperatures. The fact that, in both the northern and southern parts of the rift, volcanism pre-dated faulting by a few million years (Morley et al., 1992; Smith, 1994) also suggests that the initial magmatism was related to the presence of hot mantle material beneath the site of the present rift, such mantle having a potential temperature of perhaps 1500°C (Mechie et al., 1997).

View larger version (47K): [in this window] [in a new window]   Fig. 1. Location map of Cenozoic volcanic rocks (black) of the East African Rift System and of the Ethiopian and Kenyan plateaux or domes (dashed lines). The arrow shows the apparent migration as a result of plate motions of a mantle plume, currently centred on Lake Victoria, over the past 50 my. Modified from George et al. (1998, fig. 1).

 

Geophysical and geochemical evidence strongly points to the involvement of the Afar mantle plume in the generation of the Ethiopian flood basalts and the magmatism in Afar (Marty et al., 1996; Hoffman et al., 1997). There is, however, considerable uncertainty about the number of plumes influencing magmatism and tectonics beneath the southern Ethiopian, Kenyan and Tanzanian sectors of the rift system. Gravity and topographic anomalies reveal two broadly circular, dynamically supported plateaux, implying two distinct mantle plumes (Ebinger et al., 1989). However, a more recent suggestion (Ebinger & Sleep, 1998) involved impact of the Afar plume beneath southern Ethiopia at 45 Ma, followed by rapid lateral flow of plume material guided by the topography at the base of the lithosphere to feed magmatism within lithospheric thin-spots generated by Mesozoic and early Tertiary extension. Smith (1994) related magmatism in the main, Gregory, part of the Kenya rift to the rise of a plume focused beneath a weak zone marking the boundary between the Tanzanian craton and the adjacent Proterozoic mobile belt. He suggested that the mantle plume may lie beneath the craton or much further north, beneath Ethiopia. Birt et al. (1997) and Mechie et al. (1997) envisaged a mantle plume confined beneath the Tanzanian craton, from which hot material rises to shallower depths under the outer rims formed by the western and eastern arms of the East African Rift System.

George et al. (1998) favoured a two-plume model, proposing that the Ethiopian province was initially related to the thermal influence of the Kenyan, and subsequently the Afar, mantle plumes during northward drift of the African plate over the past 45 my. Rogers et al. (2000) used isotopic evidence to argue that a Kenya plume, whose origins are in the upper mantle, may be distinguished from the Afar plume, which has a deep mantle source. Burke (1996) favoured three plumes in the evolution of the East African Rift System. The first generated the basalts of southern Ethiopia between 45 and 35 Ma and perhaps also contributed to the Lokipiti volcanism at 35 Ma. The second, Afar, plume impinged on the base of the lithosphere and caused eruptions starting at 35 Ma. A third, Samburu, plume resulted in eruption of 30 000 km³ of basalt in north–central Kenya at 20 Ma.

In contrast, Bailey (1992) argued against the involvement of plumes at all, suggesting that Tertiary–Recent magmatism in the East African Rift System is related to lithospheric stresses from collisional tectonics between the African and Eurasian plates. Still others take the compromise view that rifting in Kenya is related to the rise of a plume (or plumes) under lithosphere already being extended in response to external forces. Zeyen et al. (1997) have argued that the Kenya rift area has been subjected to compressional far-field forces that arose partly from ridge push from the Indian Ocean and partly from the high topography of the Ethiopian Plateau above the Afar plume. Coblentz & Sandiford (1994), on the other hand, have suggested that a large part of the extensional forces in east Africa is a result of density variations within the lithosphere, rather than any far-field forces.

Attempts to use geochemical data for basalts to distinguish between the various plume (and non-plume) models are complicated by strong evidence of a lithospheric signature in the eruptive products (Norry et al., 1980; Davies & Macdonald, 1987; Rudnick et al., 1993; Class et al., 1994; Macdonald, 1994; Bell & Dawson, 1995; Paslick et al., 1995; Bell & Simonetti, 1996; Kalt et al., 1997; Rogers et al., 2000). Furthermore, the basement to the Kenya rift (implied here to be that part of the East African Rift System between Turkana and northern Tanzania) is complex, being divisible into three zones: the Archaean Tanzanian craton, the Panafrican Mozambique mobile belt and a zone of craton margin reworked during the Panafrican orogeny (Smith & Mosley, 1993). It might be expected that the nature of the lithosphere contribution to basalt composition would change at the boundaries between these zones.

In this paper, we use the first systematic survey of basalt geochemistry along the length of the Kenya rift to address three issues relating to the role of mantle plumes in rift evolution:

1. is rift magmatism related to the ascent of one or more mantle plumes? Did the plumes contribute materially to the magmatism or did they provide only the thermal energy for partial melting?
   
2. Has there been a significant lithospheric input to rift magmatism? Do the magmatic affinities of the eruptive rocks reflect variable inputs from the inferred plume and lithosphere components, and have the proportions changed either with time or with distance from any inferred plume axis?
   
3. Do the magma compositions reflect changes in the nature of the lithosphere, particularly the change from mobile belt to craton?
   

	GEOLOGICAL SUMMARY

TOP ABSTRACT INTRODUCTION GEOLOGICAL SUMMARY ANALYTICAL TECHNIQUES NOMENCLATURE AND DISTRIBUTION COMPOSITIONAL DIFFERENCES... BASALT COMPOSITION VS TIME:... COMPARISONS WITH OCEAN ISLAND... CRUSTAL CONTAMINATION MANTLE HETEROGENEITY MANTLE END-MEMBERS CONCLUSIONS APPENDIX : SAMPLE LOCALITY... REFERENCES

 
The earliest volcanic rocks in the Kenya rift are those of the Turkana region in the north, some 30–35 my old (Figs 1 and 2). Since then, the onset of magmatism has generally migrated southwards through central and southern Kenya, finally reaching northern Tanzania some 5–8 my ago. This represents a simplification of a more complicated picture, however. As the rift opened out into the rigid craton, it also developed and propagated north of Baringo (Fig. 2) to form the inner rift and Quaternary volcanic rocks. The overall southerly progression of the magmatism was accompanied by a change in the nature of the lithosphere.The Kenya rift straddles the margin of the Tanzanian craton (Fig. 2). This margin was reworked, overthrust and buried by rocks of the Mozambique mobile belt during a late Proterozoic collisional event. Integrating a wide range of geological and geophysical results, Smith & Mosley (1993) and Ebinger et al. (1997) showed that a major influence on the development of the rift has been its location above a lithological heterogeneity, with a contrast between the cold, thick, rigid Archaean lithosphere and thinner, anisotropic, warmer, mobile belt lithosphere. During the late Proterozoic, the mechanical anisotropy that existed between the two types of lithosphere was modified by a series of continental-scale NW–SE and north–south trending ductile–brittle shear zones that affected the craton margin and the mobile belt. Smith & Mosley (1993) argued that reactivation of these shear zones under varying stress field conditions has controlled the location and geometry of graben structures and the emplacement of magma since early Miocene times. Smith & Mosley (1993) recognized, therefore, three major lithospheric blocks in the structural framework of the rift (Fig. 2): the craton, reworked craton margin and the mobile belt.

View larger version (77K): [in this window] [in a new window]   Fig. 2. Structural framework of the Kenya Rift Zone. (a) Map view. Thick lines are major faults. Much simplified from Smith & Mosley (1993, fig. 3a). Lines 1–3 refer to cross-rift traverses along which rocks ranging in age from Miocene to Recent were collected. (b) North–south profile of estimated depth to basement along line of rift between X and X' (in a). Dashed line is top of basement drawn by Smith & Mosley (1993, fig. 3b) on the basis of seismic profiles.

 

Deep rift structure
Crustal thickness varies along the axial part of the rift from 20 km beneath Turkana to 35 km beneath the Kenya dome in south–central Kenya to 38 km beneath northern Tanzania (Mechie et al., 1994, 1997; Last et al., 1996). The thinning has been accomplished by the northerly thinning of the lowermost crustal layer from 9 to 2 km. The layer has a P-wave velocity of 6·8 km/s, indicating that the lower crust may be a mix of high-grade metamorphic rocks, and underplating mafic and ultramafic material (Mooney & Christensen, 1994; Hay et al., 1995; Mechie et al., 1997).

Along the length of the rift, seismic refraction surveys have indicated P_n velocities of the uppermost mantle to be 7·5–7·7 km/s, which contrasts with values of 8·0–8·2 km/s beneath the rift shoulders and flanks. Teleseismic tomography shows a velocity reduction of 6–12% in an essentially vertical channel beneath the rift axis, extending to 160 km depth or more (Achauer et al., 1994; Slack et al., 1994). The velocity reduction has been explained by the presence of small amounts (3–6%) of partial melt (Green et al., 1991; Achauer et al., 1994) and/or an rise in potential temperature by 300 K beneath the rift (Sobolev et al., 1996). The presence of steep-sided low-velocity zones extending deep into the mantle is consistent with gravity models of low-density material reaching deep into the mantle (Ebinger et al., 1989; Hay et al., 1995; Birt et al., 1997; Simiyu & Keller, 1997).

Temperature distribution with depth beneath the rift is poorly constrained. As a guide, Mechie et al. (1997) used the following values for modelling velocity distribution. Under the northern and southern sections of the Kenya rift, they estimated Moho temperatures of 630°C (21 km depth) and 1000°C (35 km), respectively. These values correspond to surface heat flows of 105–120 mW/m², close to the average of 105 mW/m² measured from the rift by Morgan (1983). For greater depths, Mechie et al. (1997) estimated, for the northern rift, temperatures rising from 940°C at 45 km to 1100°C at 63 km; for the southern rift, the temperatures for the same depths are 1200°C and 1350°C, respectively.

Surface heat flow from the Tanzania craton is much lower (<50 mW/m²), approaching normal values for stable cratonic lithosphere (Nyblade et al., 1990, 1996). Lower lithospheric temperatures are supported by microseismic studies; focal depths are greater in northern Tanzania than in Kenya. Thus heat flow and seismic evidence are consistent with the idea of a southward propagating rift system, such that at least the upper parts of the craton have not been thermally modified by the action of any inferred plume.

	ANALYTICAL TECHNIQUES

TOP ABSTRACT INTRODUCTION GEOLOGICAL SUMMARY ANALYTICAL TECHNIQUES NOMENCLATURE AND DISTRIBUTION COMPOSITIONAL DIFFERENCES... BASALT COMPOSITION VS TIME:... COMPARISONS WITH OCEAN ISLAND... CRUSTAL CONTAMINATION MANTLE HETEROGENEITY MANTLE END-MEMBERS CONCLUSIONS APPENDIX : SAMPLE LOCALITY... REFERENCES

 
A total of 128 samples were analysed for major and trace elements by X-ray fluorescence (XRF) at the University of Edinburgh, using techniques outlined by Fitton et al. (1998). The samples came from several sources, including the collections of the British Geological Survey, the East African Geological Research Unit (University of Leicester), Lancaster University, the University of Oregon and Professor J. B. Dawson. Specimen details are given in the Appendix. Localities referred to in the text are shown in the Appendix (Fig. A1). We have also used the XRF datasets for Tanzanian eruptive rocks of Paslick et al. (1995; 27 samples) and Kenyan eruptive rocks of Rogers et al. (2000; 30 samples), because they were analysed in the Edinburgh laboratory. The total number of analyses is thus 185. A subset of 59 samples was also analysed by instrumental neutron activation analysis (INAA; trace elements) at the Open University (Potts et al., 1985). Representative analyses are given in Table 1; the complete dataset is available for downloading from the Journal of Petrology Web site, at http://www.petrology.oup.journals.org.

View larger version (19K): [in this window] [in a new window]   Fig. A1. Map, simplified from Fig. 2a, showing localities referred to in the text. Lor, Lorikipi or Loriu; Na, Namarunu; K, Kerio Valley; Ka, Kaparaina; Ng, Ngelesha Escarpment; LB, Lake Bogoria; Lois, Loisiomurto; Olg, Olokisalie (Olorgesailie).

 

View this table: [in this window] [in a new window]   Table 1: Compositional data for selected volcanic rocks of the Kenya Rift Valley

 

Sr, Nd and Pb isotopic data for a sub-set of these samples are taken from the literature (Davies & Macdonald, 1987; Macdonald et al., 1995; Paslick et al., 1995; Black et al., 1998; Rogers et al., 2000).

	NOMENCLATURE AND DISTRIBUTION

TOP ABSTRACT INTRODUCTION GEOLOGICAL SUMMARY ANALYTICAL TECHNIQUES NOMENCLATURE AND DISTRIBUTION COMPOSITIONAL DIFFERENCES... BASALT COMPOSITION VS TIME:... COMPARISONS WITH OCEAN ISLAND... CRUSTAL CONTAMINATION MANTLE HETEROGENEITY MANTLE END-MEMBERS CONCLUSIONS APPENDIX : SAMPLE LOCALITY... REFERENCES

 
For this study we decided to work on mafic igneous rocks from the axial regions and shoulders of the rift, thus ensuring, first, that we had the most complete geophysical information on deep structure, and, second, that we could confidently relate our samples to the basement zones established by Smith & Mosley (1993). This approach also allowed us to have ‘time-slices’ through geographically contiguous successions, in that earlier work has shown that there has been a tendency for volcanism in any given region of the rift to migrate with time towards the axial zone (Lippard & Truckle, 1978; Truckle, 1980). The main data gap is for the older (>20 Ma) successions of the Turkana rift (Bellieni et al., 1987; Morley et al., 1992; Ebinger et al., 1993). Most of the rocks from northern Kenya were selected to lie on or close to three cross-rift traverses, all within the mobile belt (Fig. 2). The geology of much of the area has been described by Dunkley et al. (1993). Traverse 1 runs from the Kerio valley in the west through Kaparaina to Baringo in the central rift to the Ngelesha escarpment. It incorporates successions ranging from Miocene (20–15 Ma) to Recent (<120 ka). Rocks of Traverse 2 (Ribkwo–Silali volcano–Laikipia) cover a similar age range (16 Ma–a few hundreds of years). Traverse 3 runs from Lorikipi–Loriu in the west through Namurunu on the rift floor to the Tirr Tirr plateau in the east, samples ranging from 16·8 ± 2 Ma to <10 ka. The range in ages of the rocks within each traverse allows us to examine compositional variations with time at the local, as well as regional, levels. All rocks in the dataset were chosen to be essentially unaltered, with an absence or scarcity of secondary carbonates or zeolitization and, at worst, minor serpentinization or oxidation of olivine.

All samples have mg-number 40 (calculated with Fe₂O₃ = 0·15 Fe₂O₃^*) and SiO₂ < 51 wt %. According to the LeBas et al. (1986) total alkali–silica (TAS) classification scheme, the rocks are foidites, picrobasalts, basanites, basalts and hawaiites. Here we use a simplified classification, recognizing four groups on the basis of CIPW normative characteristics: (1) nephelinites (ne + leu > 10%); (2) basanites (5 < ne <10%); (3) alkali olivine basalts (0% < ne < 5%); (4) hypersthene (hy)-normative basalts. The divisions are somewhat arbitrary, in that there is a complete compositional spectrum between the most and least silica-undersaturated rocks, and rocks of more than one type are normally intimately related in the field.

The number of samples in each lithospheric domain is shown in Table 2. An important difference is in the proportions of more and less silica-undersaturated rock types. For example, in the craton, nephelinites and basanites form 80% of our samples and alkali olivine basalts and hy-normative basalts 20%. The corresponding figures for the reworked craton margin are 33% and 67%, and for the mobile belt 21% and 74%. These proportions would strictly be applicable only if our sampling had been truly random. Nevertheless, they are consistent with previous observations of the distribution of magma types based on detailed field mapping (e.g. LeBas, 1987; Smith & Mosley, 1993; Smith, 1994). In the Kenya rift, olivine-poor nephelinitic and carbonatitic volcanism is largely restricted to the craton and reworked craton margins. In the mobile belt, olivine-rich nephelinitic volcanism is associated with large shield volcanoes and high-volume eruptions of alkali and transitional basalts. Smith (1994) has related the differences to thinner lithosphere beneath the mobile belt, which is more easily eroded and penetrated by rising magmas, partly because it had previously undergone significant stretching and warming during Cretaceous and Palaeogene rifting events.

View this table: [in this window] [in a new window]   Table 2: Average compositions in each tectonic setting

 

Mg-numbers for the mafic rocks vary from 40 to 80 but only 19 samples possibly represent primary magmas, i.e. mg-number >65. We appreciate that although the use of mafic rocks minimizes the effects of polybaric fractionation on incompatible trace element ratios, it does not totally preclude them. Nevertheless, the spread of compositions that we shall report here, e.g. in incompatible trace element ratios, is independent of mg-number, such that the trace element characteristics of the eruptive rocks dominantly reflect mantle source composition and mantle processes.

Compositional relationships between magma types
Within each setting, the transition from average nephelinite through basanite and alkali olivine basalt to hy-normative basalt is marked overall by decreases in Na₂O, K₂O and the incompatible trace elements (ITE), and increases in SiO₂ (Table 2). This relationship is normally considered to reflect different degrees of partial melting of compositionally similar mantle sources. Figure 3 is a Ce/Y vs Zr/Nb plot, where the continuous lines represent non-modal fractional melting curves for partial melting of fertile and depleted spinel peridotite and garnet peridotite (Hardarson & Fitton, 1991). Two features may be noted. First, within each setting, as exemplified by the rocks of the reworked craton margin in the inset in Fig. 3, the sequence of average compositions nephelinite hy-normative basalt appears to represent progressively shallower (increased component of spinel-facies peridotite) melting, rather than larger degrees of partial melting, although the spread within the dataset for each setting indicates variable melt fractions. Second, regardless of magma type, the craton rocks appear to have been generated at the highest pressures and the mobile belt rocks at the lowest, although there is substantial overlap between the magma types. The lower heat flow through the craton might be expected to result in smaller volumes of melt and smaller melt fractions. Whereas the former is clearly true, the latter does not seem to be so because the trace element ratios suggest similar melt fractions, albeit from greater depth. Therefore controls on melt fraction are not directly related to lithosphere thickness, although depth of melting is. However, this assumes that the characteristics of the melts are totally controlled by the melt regime, whereas they must also reflect the nature of the source region, e.g. whether it is enriched or not and the style of that enrichment.

View larger version (26K): [in this window] [in a new window]   Fig. 3. Ce/Y–Zr/Nb plot of mafic volcanic rocks. The continuous lines are non-modal fractional melting curves calculated by Hardarson & Fitton (1991) for four mantle compositions: GD, depleted garnet lherzolite; GP, primitive garnet lherzolite; SD, depleted spinel lherzolite; SP, primitive spinel lherzolite. Numbers on lines refer to percentages of melt. In the main diagram, all data are plotted, subdivided by structural setting. In the inset, the average compositions of each lithology in the reworked craton margin are shown. RCM, reworked craton margin; MB, mobile belt; BSN, basanite; AOB, alkali olivine basalt; HNB, hypersthene-normative basalt.

 

The major and trace element data indicate, therefore, that the mafic rocks were formed, within each setting, over a range of pressures, roughly spanning the garnet–spinel transition, and by variable, but apparently always small, degrees of melting.

	COMPOSITIONAL DIFFERENCES BETWEEN SETTINGS

TOP ABSTRACT INTRODUCTION GEOLOGICAL SUMMARY ANALYTICAL TECHNIQUES NOMENCLATURE AND DISTRIBUTION COMPOSITIONAL DIFFERENCES... BASALT COMPOSITION VS TIME:... COMPARISONS WITH OCEAN ISLAND... CRUSTAL CONTAMINATION MANTLE HETEROGENEITY MANTLE END-MEMBERS CONCLUSIONS APPENDIX : SAMPLE LOCALITY... REFERENCES

 
Major and trace element data
An aim of this project was to determine whether the composition of the mafic eruptive rocks of the rift is related to lithosphere type, in particular the primary distinction between mobile belt, reworked craton margin and craton. At this stage, we shall examine this question using average major and trace element compositions for each magma type (Table 2), selected major element plots (Fig. 4) and mantle-normalized plots for average lithologies in each setting (Fig. 5). Comparisons are slightly complicated by the different average mg-number of rocks from each setting. For example, highly magnesian, strongly undersaturated (leucite-normative) rocks are virtually restricted to the craton and reworked craton margin, such that the average nephelinites from these settings are higher in CaO, K₂O, TiO₂ and P₂O₅, and lower in SiO₂ and Na₂O than average nephelinite from the mobile belt. Allowing for such differences in mg-number, the major element compositions for each magma type are seen to be very similar (Table 2). Minor differences are that nephelinites and basanites from the craton tend to be more potassic than those from the other settings (Fig. 4a), whereas rocks from the mobile belt and reworked craton margin tend, at given MgO, to be more aluminous than those from the craton (Fig. 4b), possibly because they have equilibrated at slightly lower pressures or were derived from a more enriched or less depleted source.

View larger version (18K): [in this window] [in a new window]   Fig. 4. (a) Na2O/K2O–SiO2 plot for nephelinites and basanites. Apart from a subgroup of nephelinites with relatively low SiO2 (40 wt %) and low Na2O/K2O (2·5), restricted to the cratons, the rocks in each setting show extensive overlap. (b) MgO–Al2O3 plot, using all data, subdivided into lithospheric domains.

 

View larger version (19K): [in this window] [in a new window]   Fig. 5. Mantle-normalized element plots showing the average composition of each magma type in the craton, reworked craton margin and mobile belt. Normalizing values from Sun & McDonough (1989).

 

Incompatible trace element variations in each magma type are summarized in mantle-normalized diagrams in Fig. 5. Patterns for the nephelinites are remarkably similar in each setting, average abundances increasing in the order mobile belt through reworked craton margin to craton. The basanite and alkali olivine basalt patterns are overall similar to those in the nephelinites, although the order of enrichment of individual incompatible trace elements is more variable; for example, in the basanites, the order of enrichment in Th is craton > reworked craton margin > mobile belt; for P the order is reversed. The patterns for the hy-normative basalts are slightly different from those of the other types in that the K trough is much less marked, especially in the craton margin and mobile belt rocks.

Mantle-normalized REE plots are given in Fig. 6. Light REE (LREE) enrichment relative to heavy REE (HREE) is generally strongest in more undersaturated rocks, i.e. nephelinites and basanites > alkali olivine basalts > hy-normative basalts. This is consistent with a greater garnet control in undersaturated rocks and greater spinel control in hy-normative basalts. A La/Yb–Tb/Yb plot can be used to summarize differences in REE behaviour between settings (Fig. 7). There is very significant overlap between rocks from each setting but basalts from the mobile belt show a slight tendency to have less fractionated REE patterns, with lower values of both ratios than the craton and craton margin samples. The data define a broad linear trend between melting curves calculated for fractional melting of fertile garnet-bearing and garnet-free peridotite. The steep trend indicates an important role for garnet in the craton and craton margin basalts, i.e. they were generated at higher pressures, which is consistent with Ce/Yb–Zr/Nb relationships (Fig. 3). As with the other incompatible trace elements, there are no major differences in REE patterns for the same lithologies in the different settings. There would appear, therefore, to be no compelling evidence from the major and trace element data that lithospheric type has strongly influenced magma composition.

View larger version (29K): [in this window] [in a new window]   Fig. 6. Mantle-normalized REE plots of the average composition of each magma type in craton, reworked craton margin and mobile belt. Normalizing values from Sun & McDonough (1989).

 

View larger version (35K): [in this window] [in a new window]   Fig. 7. La/Yb–Tb/Yb plot of rocks from the craton, reworked craton margin and mobile belt, with the field of oceanic island basalts (OIB) for comparison. Continuous lines are for melting of fertile lherzolitic mantle, the contours representing the amount of modal garnet. The small, 0%, field represents melting in the spinel peridotite facies. Partition coefficients from Halliday et al. (1995). Data for the OIB field in this and subsequent diagrams are from a compilation by N. W. Rogers of published analyses.

 

Sr–Nd–Pb isotopic evidence
⁸⁷Sr/⁸⁶Sr in the whole dataset ranges from 0·7030 to 0·7057, ¹⁴³Nd/¹⁴⁴Nd from 0·5130 to 0·5124 and ²⁰⁶Pb/²⁰⁴Pb from <18 to >21. These ratios generally lie within the ranges defined by ocean island basalts (OIB) (Hofmann, 1997). In contrast to major and trace element compositions, the boundary between the craton (including its reworked margin) and the mobile belt is reflected in the Sr and Nd isotope ratios of the basalts (Fig. 8). Those basalts erupted through the craton have ¹⁴³Nd/¹⁴⁴Nd < 0·51275 and ⁸⁷Sr/⁸⁶Sr > 0·7035, whereas those from the mobile belt have ¹⁴³Nd/¹⁴⁴Nd > 0·51275 and ⁸⁷Sr/⁸⁶Sr ( 0·7035. This is the only clear evidence of a control exerted by the underlying lithosphere on magma compositions. The data for rocks from the craton and craton margin define two sub-parallel trends in which ⁸⁷Sr/⁸⁶Sr increases with little variation in ¹⁴³Nd/¹⁴⁴Nd. Such trends have been shown locally to be related to crustal contamination, e.g. at the Naivasha (Davies & Macdonald, 1987) and Silali (Macdonald et al., 1995) complexes.

View larger version (23K): [in this window] [in a new window]   Fig. 8. 87Sr/86Sr–143Nd/144Nd plot for all data, subdivided by structural setting. (Note the approximately horizontal data arrays at 143Nd/144Nd values of 0·5127 and 0·5126.)

 

Variations in Pb isotopic ratios in the whole dataset are parallel to, but generally displaced above, the Northern Hemisphere Reference Line (NHRL) (Fig. 9). Mobile belt rocks are restricted to the lower and middle parts of the range (²⁰⁶Pb/²⁰⁴Pb 18·5–19·5). Rocks from the reworked craton margin also show a rather restricted range and are displaced to higher ²⁰⁷Pb/²⁰⁴Pb than those from the mobile belt. Pb isotopic ratios in the craton rocks define a broadly linear array that crosses the NHRL. If the array has age significance, an age of 2–2·5 Ga is inferred, which is comparable with the 2·5–2·7 Ga age of the oldest rocks in the Tanzania craton (Bell & Dobson, 1980) and the age of 2·5–2·9 Ga for lithosphere stabilization, deduced from Re/Os studies of Tanzanian mantle xenoliths by Chesley et al. (1998).

View larger version (20K): [in this window] [in a new window]   Fig. 9. 207Pb/204Pb–206Pb/204Pb and 208Pb/204Pb–206Pb/204Pb plots, using all data, subdivided by structural setting. Northern Hemisphere Reference Line (NHRL) from Hart (1984).

 

The fact that Pb isotopes do not distinguish between mafic rocks erupted through the craton and mobile belt as effectively as Sr and Nd isotopes may be in part because Pb isotopes are more susceptible to crustal contamination than are Sr and Nd. However, there is no correlation between such indices of crustal contamination as Ce/Pb and ²⁰⁷Pb/²⁰⁴Pb, suggesting that the major control on Pb isotopes has been source composition. The complexity of isotopic relationships in the craton was noted by Paslick et al. (1995), who showed, for example, that the combination of HIMU Pb and EMI Sr and Nd means that the lavas did not acquire their isotopic compositions simply from an asthenospheric OIB source. Attempts to identify mantle components are even more complicated when the full dataset is considered. Excluding any crustal component, rocks from all three lithospheric domains would require three- or four-component mixes of the hypothetical end-members recognized in oceanic basalts—HIMU, EMI, EMII and DM (Fig. 10).

View larger version (18K): [in this window] [in a new window]   Fig. 10. 87Sr/86Sr–206Pb/204Pb plot, using all data, subdivided by structural setting. Values for HIMU, EMI, EMII and DM end-member compositions from Hart et al. (1992).

 

	BASALT COMPOSITION VS TIME: SOME GENERALIZATIONS

TOP ABSTRACT INTRODUCTION GEOLOGICAL SUMMARY ANALYTICAL TECHNIQUES NOMENCLATURE AND DISTRIBUTION COMPOSITIONAL DIFFERENCES... BASALT COMPOSITION VS TIME:... COMPARISONS WITH OCEAN ISLAND... CRUSTAL CONTAMINATION MANTLE HETEROGENEITY MANTLE END-MEMBERS CONCLUSIONS APPENDIX : SAMPLE LOCALITY... REFERENCES

 
On the basis of spatial and temporal variations in basalt composition along the northern sector of the rift, Lippard & Truckle (1978) and Truckle (1980) suggested that (1) mafic lavas west and east of the rift are generally more silica undersaturated than are lavas within the rift; (2) most sectors of the within-rift sequences show a decrease of silica undersaturation of mafic lavas with time; (3) during the Miocene and Pliocene, mafic lavas became increasingly silica undersaturated along the rift axis towards the south; (4) during the Quaternary, mafic lavas have shown a trend of increasing silica undersaturation northwards and southwards away from the Naivasha area, the culmination of the so-called Kenya Dome.

These relationships do not always hold at the regional and local levels. Crossley & Knight (1981) found that there is no simple relationship between magma alkalinity and distribution in space and time in the western part of the southern rift in Kenya, whereas lavas ranging from Tertiary to Recent age in the Turkana rift seem to have been produced from similar sources by similar processes (Bloomer et al., 1989). Individual central volcanoes, such as Olokisalie (Olorgesaillie) in the southern rift, can contain lavas ranging from nephelinites to hy-normative basalts (Baker, 1987). In the part of the Kenya rift between Lake Bogoria and Turkana, Dunkley et al. (1993) reported that Quaternary basalts become increasingly undersaturated northwards and that, in centres that have erupted both types of basalts, the hy-normative varieties pre-date the ne-normative types.

Our new data largely corroborate the generalizations about the distribution of magma types in time, at least for the 20 my period represented by our dataset. At scales varying from that of the rift through regional (e.g. Traverse 2; Samburu volcanic rocks; southern rift) to that of individual centres (Olorgesaillie), magmatism tends on average to become less silica undersaturated with time (Fig. 11). This is caused by two effects; nephelinites disappear from the sequences and hy-normative basalts become more important volumetrically. For example, there is only one hy-normative basalt older than 10 Ma in our samples; in the mobile belt and reworked craton margin, all nephelinites are older than 11 Ma and 2·5 Ma, respectively. We have no samples of very recent nephelinitic magmatism from the mobile belt and reworked craton margin. Because at any one locality magmatism has tended to migrate towards the centre of the rift, it follows that the more silica-saturated rocks are concentrated towards the axis, although not exclusively.

View larger version (20K): [in this window] [in a new window]   Fig. 11. Range of eruptive ages of each magma type in each setting.

 

At first sight, the disappearance of nephelinites from the eruptive sequences might appear consistent with the idea of Thompson & Gibson (1994) that prolonged magmatism at a rift site removes most or all of the (usually potassic) lithospheric-source melt, allowing subsequent melts from convecting mantle to penetrate to the surface without significant interaction with lithosphere-derived melts. We shall show below, however, that at any site in the Kenya rift, the lithospheric signature in the mafic rocks does not decrease systematically with time. Instead, the time trends (Fig. 11) seem to indicate that soon after initiation of magmatism in each area, the melt column extended from the generation depths of nephelinites (100 km; Macdonald et al., 1994) to those of the hy-normative basalts (<50 km). With time, the column tended to move upwards, such that nephelinitic magmatism ceased and alkali olivine basalts and hy-normative basalts became increasingly important volumetrically.

	COMPARISONS WITH OCEAN ISLAND BASALTS

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Deviations of the Kenya data from OIB are clearly seen in plots of elements of different compatibilities, normalized to Y to minimize the effects of fractional crystallization. Thus in the Nb/Y–Zr/Y plot (Fig. 12), several rocks from all three tectonic settings show elevated Nb/Y at a given Zr/Y relative to OIB. Nevertheless, the large degree of overlap between the Kenya and OIB fields suggests some common source and/or process. None of the chemical features is ubiquitous. Zr depletion relative to Ti, noted previously in Kenyan basalts by Davies & Macdonald (1987) and Macdonald (1994), is mainly developed in rocks of the mobile belt. Phosphorus enrichment relative to Nd is essentially restricted to the mobile belt. Rb depletion relative to Ba and K, and Nb enrichment relative to K and La are common, but not universal, characteristics. Most, but not all, of the rocks show Ba enrichment relative to mean OIB. A measure of the complexity is given by Ce/Pb and P/Nd ratios, which are relatively constant in OIB (25 ± 5 and 74 ± 13, respectively; Sun & McDonough, 1989) but show large ranges in the rift rocks (5–100 and 20–155). The great diversity in the rift eruptive rocks indicates that they did not acquire their trace element features simply from an asthenospheric OIB-like source (see Macdonald, 1994). We noted above that, with the exception of rocks of the mobile belt, isotopic evidence also presents a more complicated picture than a simple OIB-like source for the rift magmas, a point made earlier for Tanzanian rocks by Paslick et al. (1995) and for Kenya rift carbonatites by Kalt et al. (1997). It would appear that the mafic rocks contain a very heterogeneous, lithospheric component. We now assess the contribution of crustal contamination to the lithospheric component.

View larger version (38K): [in this window] [in a new window]   Fig. 12. Zr/Y–Nb/Y plot to show the Zr depletion relative to Nb of many Kenyan eruptive rocks compared with OIB.

 

	CRUSTAL CONTAMINATION

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Various lines of major and trace element evidence suggest that crustal contamination has not contributed significantly to the compositional spread in the mafic rocks of the Kenya rift: (1) the majority of specimens (81%) are nepheline-normative, precluding an extensive crustal fractionation history, a point already made for Tanzanian rocks by Paslick et al. (1995); (2) 58% of our samples contain higher abundances of Ba than the average upper continental crust (500 ppm; Taylor & McLennan, 1985); Ba concentrations in the Kenya rocks have been buffered by high concentrations in the melt derived from the mantle sources and any evidence of crustal contamination is obscured (see Paslick et al., 1995); (3) the spread of data in a La/Yb–Sm/Eu plot (Fig. 13) precludes significant contamination of OIB-type magmas with rocks comparable with either average upper or lower continental crust.

View larger version (28K): [in this window] [in a new window]   Fig. 13. Sm/Eu–La/Yb plot for rocks of all three structural settings. The distribution is not consistent with significant amounts of mixing of OIB with either average upper (UCC) or lower (LCC) continental crust (Taylor & McLennan, 1985).

 

Within the rocks of the craton and craton margin, Sr and Nd isotopic data define two flat-lying trends in which ⁸⁷Sr/⁸⁶Sr increases with little variation in ¹⁴³Nd/¹⁴⁴Nd (Fig. 8). In basalts from the Naivasha area in south–central Kenya, this increase in ⁸⁷Sr/⁸⁶Sr is associated with an increase in ²⁰⁷Pb/²⁰⁶Pb, a decrease in ²⁰⁶Pb/²⁰⁴Pb and little change in La/Y or Nb/Y, and has been related by Davies & Macdonald (1987) to crustal contamination. Paslick et al. (1995) have also used Pb isotopic data to show that minor crustal contamination may have affected some rocks from Tanzania, such as the hy-normative basalt MD93-7, which has the lowest ²⁰⁶Pb/²⁰⁴Pb ratio (17·632) in our dataset. Samples with high ⁸⁷Sr/⁸⁶Sr are not generally associated with strong indicators of crustal interaction. On the contrary, they appear to have the characteristics of small melt fractions from the mantle, in particular low SiO₂ contents and Rb/Sr ratios and high Sr contents, and, within the craton samples, high Nb/Y and La/Y. The samples from the reworked craton margin show elevated ⁸⁷Sr/⁸⁶Sr with little change in La/Y or Nb/Y and these may well have been influenced by crustal interaction.

In contrast, basalts erupted through the mobile belt show no systematic correlation between isotopic variation and indices of crustal contamination, such as Rb/Sr or Ce/Pb. This apparent lack of crustal interaction may be related to the relative thinness of the crust. Seismic profiles along the rift have revealed that the crust thins from 35 km close to the craton margin to 20 km beneath the Turkana Depression. The attenuated crust is probably related to the numerous extensional events that have affected northern Kenya during the Mesozoic and early Tertiary (Mechie et al., 1997).

We conclude that although some compositional spread in the dataset is undoubtedly related to minor crustal contamination, the effects tend to be swamped by mantle-derived variability.

	MANTLE HETEROGENEITY

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Thus far, we have discussed major and trace element data partly in terms of average compositions. This approach disguises considerable compositional variation within groups, particularly in the incompatible trace elements, taken to include K, Ti and P. Thus, incompatible trace element abundances commonly have a range, within each lithological type, of a factor of four at given mg-number. Class & Goldstein (1997) also reported fractionation-corrected, incompatible trace element concentrations with a range of a factor of four in alkali basalts from La Grille, Grande Comore. In the Kenya case, the large ranges result in large overlaps between magma types. We now point out some important features of the variability in incompatible trace element abundances, first with reference to P. Whereas the great majority (>160) of rocks show a fair TiO₂–P₂O₅ correlation (and TiO₂/P₂O₅ ratio >3), around 20 have lower ratios. On a P/Y–Zr/Y plot (Fig. 14), these show a trend of steeply increasing P/Y with little change in Zr/Y. The high-P rocks are restricted to the reworked craton margin and mobile belt. Although they are particularly common along Traverse 2 (in successions ranging in age from Miocene to Recent), they also occur as Miocene lavas at Loisiumurto in SW Kenya, in Pleistocene basalts at Thiba (central Kenya) and in young basalts at the Barrier. Thus the anomalously high-P rocks are found throughout Kenya and in rocks covering the age span of our dataset.

View larger version (44K): [in this window] [in a new window]   Fig. 14. P/Y–Zr/Y plot to show the sub-group of Kenyan mafic rocks with relative P enrichment. Field of OIB also shown.

 

Relationships between many pairs of incompatible trace elements show similar features to P, namely a majority of specimens showing a strong positive correlation and a smaller number of samples adding wide scatter. The anomalous features do not characterize all rocks in a given volcanic succession; for example, the volcano Silali contains high-P and normal-P basalts of essentially the same age. The ‘outriders’ for one element are not usually those for another, i.e. there is substantial decoupling of the incompatible trace elements, reflected, inter alia, in the ranges of ratios and in different mantle-normalized patterns, even for rocks of the same magma type from the same successions (Fig. 15).

View larger version (17K): [in this window] [in a new window]   Fig. 15. Mantle-normalized plots for Pliocene alkali olivine basalts of Traverse 2. Normalizing values from Sun & McDonough (1989). , , •, rocks from the Nasorut Basalt Formation (specimen numbers 12/1012, 12/384 and 12/980); , from the Lokweilabit Basalt (3/190); , from the Ribkwo Basalt (5/104); , from the Kaparaina Basalt (3/490).

 

The variability in incompatible trace element ratios is present in all four compositional groups and in rocks with mg-number ranging from 40 to 80. Some of the relative enrichments and depletions must, therefore, be mantle derived and we can envisage one or more of the mantle sources as being strongly heterogeneous on a local scale. As it is unlikely that convecting mantle can show heterogeneity on such a scale, we infer that this highly variable component must be lithospheric.

	MANTLE END-MEMBERS

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On the basis of the trace element and isotopic characteristics and the strong heterogeneity implied by incompatible trace element variability, it would appear, therefore, that we must consider at least a two source component model for the rift magmas, one of sublithospheric (plume?) origin, as is also required by magma volumes (Latin et al., 1993), and the other originating within the lithosphere.

The lithospheric component
We consider first the possibility that the variable composition of the lithospheric mantle is a result of metasomatism by either subduction-related fluids or carbonatitic melts.

On the basis of xenolith studies in the Tanzanian craton, Rehkamper et al. (1997) identified a metasomatic component derived from fluids released from a subducting slab. Such a component would have the advantage, in the broader rift context, of explaining the Ba enrichment in many rocks. We test it here using Ce/Pb–La/Nb relationships. Subduction-related fluids carry Pb and La in preference to Ce and Nb, respectively (Keppler, 1996; You et al., 1996). Thus arc rocks tend to have higher La/Nb and lower Ce/Pb than OIB. Using these criteria, Fitton (1995) has interpreted older (>5 Ma) lavas of the Basin and Range Province in the western USA as having a subduction component whereas younger (<5 Ma) lavas were generated within OIB-type asthenosphere (Fig. 16). In the same plot, the great majority of Kenyan rocks have OIB-like ratios of these elements; although some have higher La/Nb, they do not have the lower Ce/Pb indicative of the influence of subduction fluids.

View larger version (28K): [in this window] [in a new window]   Fig. 16. Ce/Pb–La/Nb plot of Kenya mafic rocks. For comparison, two fields for rocks from the Basin and Range Province are also shown. The older (>5 Ma) rocks are believed to show a subduction zone signature, the younger rocks to be of asthenospheric origin (Fitton, 1995).

 

Rudnick et al. (1992, 1993) argued that, at least locally, the source regions of Tanzanian xenoliths have been modified by reaction of refractory peridotite residual mantle with carbonatite melts. The criteria proposed by Rudnick et al. (1993) for recognition of the signature of carbonatite metasomatism include high La/Yb, very high Zr/Hf and very low Ti/Eu. In rocks of the Kenya rift, Zr/Hf ratios are close to constant at 44 ± 2, slightly higher than the average OIB value (36; Sun & McDonough, 1989). Ti/Eu ratios are also distinctly OIB-like, the Kenyan average (5941) being close to the OIB average of 5733. Accepting the Rudnick et al. criteria as viable, we see little evidence of carbonatite metasomatism of the sources of the primary magmas of the Kenya rift. A notable feature of the Kenya rift rocks is the relative K depletion, seen as K troughs in Figs 5 and 15. There is an overall covariation between the degree of K depletion and the degree of silica undersaturation; in each setting, for example, the average K/Nb ratio increases in the sequence nephelinite basanite alkali olivine basalt hy-normative basalt. Such behaviour has frequently been recorded from alkali basalt sequences, in Kenya (Davies & Macdonald, 1987; Class et al., 1994) and elsewhere (Hoernle & Schmincke, 1993; Haase & Devey, 1994; Halliday et al., 1995; Späth et al., 1996; Class & Goldstein, 1997). The average OIB also shows relative K depletion (Fitton et al., 1991), such that the mechanism that generates it must be generally applicable.

Halliday et al. (1995) noted that the abundances of incompatible trace elements in OIB from the central Atlantic included relatively enriched compositions in areas of older oceanic lithosphere. A feature of these magmas was relative K depletion. They modelled the magma genesis in two stages. First, small degree melts were generated in equilibrium with minor amphibole (2%), sulphide (0·2%) and phlogopite (0·2%), as well as ol + opx + cpx + gt. These melts metasomatized the uppermost mantle, which was then remelted to produce the enriched magmas. Geochemical features of the second phase melts included depletions in K, Pb and Ti, and enrichment in Zr. Such a model is not directly applicable to the Kenya case, where many rocks show Zr and Ti depletion and Pb enrichment, as well as K depletion.

Class & Goldstein (1997) also appealed to residual amphibole as a means of generating K depletion in OIB type magmas from Oahu, Hawaii and Grande Comore. Modelled compositions of small-degree batch melts of amphibole harzburgites show relative depletion in K. Addition of phlogopite to the residual mineralogy further results in Rb depletion (relative to Ba) and Zr depletion (relative to Sm). The Nb and Ta enrichment characteristic of the Kenya mafic rocks is not, however, easily explained by the presence of either phlogopite or amphibole, both of which have relatively high partition coefficients for those elements (Chazot et al., 1996; Class & Goldstein, 1997). Class & Goldstein (1997) proposed that the hydrous phases resulted from metasomatism of the source regions by volatile-rich melts or fluids before partial melting. An important difference between their model and that of Halliday et al. (1995) is that they see the metasomatized mantle as being lithospheric rather than asthenospheric, because the upper temperature stability limits of amphibole and phlogopite are much lower than temperatures reasonably inferred for convecting upper mantle (Mengel & Green, 1989). Class & Goldstein (1997) envisaged that the presence of the hydrous phases lowers solidus temperatures sufficiently that partial melting results from conductive heating of the lithosphere by upwelling thermal plumes. Formation of the Oahu and Grande Comore basalts is then the result of mixing of plume melts and small-degree melts of metasomatized lithosphere.

Broadly comparable conclusions were reached by Bedini et al. (1997) from their study of spinel peridotite xenoliths from the Sidamo region in the southeastern Ethiopian rift. They described a suite of apatite-bearing porphyroclastic peridotites strongly enriched in large ion lithophile elements (LILE; Ba, Th, U, Sr and LREE), with negative anomalies of the high field strength elements (HFSE; Nb, Ta, Zr, Hf and Ti). Bedini et al. (1997) ascribed the formation of the xenoliths to the infiltration of plume-derived magmas into the lithosphere, the process probably being related to thermo-mechanical erosion of the lithosphere mantle. Infiltration generated a high-porosity domain of regional extent at the transition between adiabatic and conductive mantle. The high porosity is thought to reflect melt accumulation at the base of the lithosphere, because of its impermeability to basaltic melts, combined with melt–rock reactions at increasing melt mass. This domain represents a potential source for extension-related basalts with a lithospheric geochemical signature.

A model of plume–lithosphere interaction is attractive for the Kenya rift. It explains some of the geochemical features of the mafic rocks, such as enrichment in Ba and depletion in Zr, although it predicts Nb and Ta depletion rather than the enrichment shown by the Kenya rocks. Furthermore, by assuming that the degree of infiltration by plume melts and of melt–rock reactions were highly variable over short distances, it allows for considerable variation in incompatible trace element concentrations in closely associated rocks. Importantly, locating the interaction zone close to the base of the lithosphere would decrease the involvement of the overlying lithosphere, resulting in basalts with very similar compositional characteristics being erupted through a range of basement types.

Identification of mantle plume component
The lithospheric mantle has exerted a strong control on the trace element and isotope geochemistry of the mafic rocks of the Kenya rift, making identification of the mantle plume component required by the erupted volumes very difficult. However, in the Sr–Nd isotope plot (Fig. 8), the craton and mobile belt trends have a common end-member at ¹⁴³Nd/¹⁴⁴Nd 0·51275 and ⁸⁷Sr/⁸⁶Sr 0·7035, as noted by Rogers et al. (2000). This composition lies within the range of OIB (Hofmann, 1997). Although correlations between Pb and Nd and Sr isotopes are complex, ²⁰⁶Pb/²⁰⁴Pb ranges from 18·7 to 19·8 in alkali olivine basalts and hy-normative basalts of the mobile belt, with ¹⁴³Nd/¹⁴⁴Nd 0·5127–0·51275 and ⁸⁷Sr/⁸⁶Sr 0·7035. This Pb value is also shown by many OIB (Hofmann, 1997). Furthermore, in mantle-normalized plots (Fig. 5), the majority of rift patterns are broadly similar to OIB, although varying in detail. It seems probable, therefore, that any plume component had OIB-like characteristics.

The majority of basalts of the Afar region, derived from the Afar plume, have ¹⁴³Nd/¹⁴⁴Nd around 0·5129, ⁸⁷Sr/⁸⁶Sr close to 0·7035 and ²⁰⁶Pb/²⁰⁴Pb of 19·5 (Rogers et al., 2000). They thus differ from the Kenyan mafic eruptive rocks in having generally higher ¹⁴³Nd/¹⁴⁴Nd for a given ⁸⁷Sr/⁸⁶Sr. The Afar plume shows elevated ³He/⁴He signatures (up to 18 R/R_a; R/R_a is the ratio of ³He/⁴He in the sample to that in the atmosphere). By contrast, the Kenyan mantle may have an R/R_a value of eight, which is the maximum R/R_a ratio found in Kenyan groundwaters (Darling et al., 1995; Marty et al., 1996; Scarsi & Craig, 1996). Rogers et al. (2000) have inferred the existence of a mantle plume below Kenya that is a different entity from that beneath Afar, implying that at least two sub-lithospheric upper-mantle source regions contribute to the mafic rocks of the EARS within a length scale of 2000 km.

An important issue for magmatism of the Kenya rift is whether the influence of any inferred mantle plume(s) has increased with time. To test this, we assume that an increasing plume component of OIB-type would be seen as a trend of Ce/Pb and P/Nd ratios towards their (roughly constant) OIB values of 25 ± 5 and 74 ± 13 (Sun & McDonough, 1989). We concentrate on rocks of the mobile belt, where the isotopic features of the basalts are most OIB-like and where geophysical evidence suggests that the asthenosphere has penetrated to high level. Figure 17 presents data for rocks of Miocene, Pliocene and Pleistocene–Holocene age; there has clearly been no significant change with time towards an increased mantle plume, or any other, component. If the southwards initiation of magmatism has been caused by plume migration or by the passage of the African plate over a plume, then the chemical signature of the plume has been heavily overprinted by interaction with lithospheric mantle.

View larger version (22K): [in this window] [in a new window]   Fig. 17. Ce/Pb–P/Nd plot for alkali olivine basalts of differing ages in the mobile belt. There has apparently been no trend with time towards an OIB-type composition, shown by a star (Ce/Pb = 25 ± 5, P/Nd = 74 ± 13).

 

	CONCLUSIONS

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1. Mafic eruptive rocks, ranging in composition from nephelinites to hy-normative basalts and in age from 16–20 Ma to Recent, have been analysed from near-axial portions of the Kenya rift underlain (from south to north) by craton, reworked craton margin and mobile belt.
   
2. The only clear distinction geochemically between rocks from the craton (and its reworked margin) and the mobile belt are Sr and Nd isotopic ratios. Otherwise, their compositions do not seem to reflect the nature of the basement through which they were emplaced.
   
3. Although an OIB-type component is apparently present in the eruptive rocks of the Kenya rift, it has been overprinted by heterogeneous lithospheric mantle. The extent of lithosphere interaction does not seem to have varied systematically in space or time.
   
4. Any effects of crustal contamination have been swamped by the mantle lithosphere component.
   
5. It is difficult to characterize chemically a plume component in the rift rocks with any great confidence, although Nd and He isotopic systematics suggests that this component is distinct from the Afar plume (Rogers et al., 2000). It is still more difficult to determine whether the rocks have been affected by more than one plume below Kenya.
   
6. The geochemical characteristics of the rift mafic eruptive rocks may be a result of infiltration of plume-derived melts into the base of the lithosphere, in a process related to the thermo-mechanical erosion of the lithosphere.
   

	APPENDIX : SAMPLE LOCALITY LIST

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Volcanic centre and/or formation name are given, along with an age determination for individual samples or the age range of the relevant formation, where known. Sample locations, where known in detail, are given either as six- or eight-figure grid references or as latitude and longitude. Otherwise, reference is given to the source literature. Sample details for additional samples have been given by Paslick et al. (1995) and Rogers et al. (2000). Figure A1 shows sample localities.

Mobile belt
Traverse 3
11/100, 11/110: Kopsorogel nephelinite (Miocene; 16·8 ± 2 Ma; Truckle, 1977)

11/213, 11/223, 11/115: Lokwanmur Basalt (Miocene; 11/213, 11·5 ± 2 Ma; 11/223, 15 ± 0·4 Ma; Truckle, 1977)

11/202: Lomujal Basalt (16·1 ± 0·5 Ma; Truckle, 1977)

11/187: Nathelat Basalt (Pliocene; 4·5 ± 0·4 Ma; Truckle, 1977)

11/511: Lorikipi Basalt (Pliocene; 2·21 ± 0·2 and 1·87 ± 0·3 Ma, respectively; Truckle, 1977)

11/188: Napetian Basalt (Pliocene; 5·9 ± 2 Ma; Truckle, 1977)

KB202: young basalt, Namarunu volcano (BN 21932144; 500 ka–Recent; Dunkley et al., 1993)

KB240, KB187: young basalts, Barrier volcano (BN 23322648, BN 22662522; <10 ka; Dunkley et al., 1993)

Traverse 2
3/751: Kapcherat Basalt (213362; 16–13 Ma; McClenaghan, 1971)

5/43, 5/330, 3/559, 3/718, 3/705: Tirioko Basalt (718485, 311499, 264321, 371340, 281375; Miocene; 6·9 Ma; McClenaghan, 1971; Webb, 1971)

12/1044, 12/1068, 12/1069, 12/1087, 12/1090, 12/536, 12/702: Chembalo Basalt (Miocene; 16–14 Ma; Golden, 1979)

12/790, 12/777, 12/790, 12/800: Lotodo Basalt (Miocene; 14 Ma; Golden, 1979)

5/104: Ribkwo Basalt (728462; Pliocene; 4·9–3·7 Ma; Webb, 1971)

3/490, 3/461: Kaparaina Basalt (212199, 231180; Pliocene; 8–4 Ma; McClenaghan, 1971)

3/33, 3/57, 3/190: Lokweilabit Basalt (323105, 328112, 311115; Pliocene; 4–3 Ma; McClenaghan, 1971)

12/980, 12/984, 12/384, 12/241, 12/1012, 12/814, 12/829, 12/1000: Nasorut Basalt (Pliocene; 9·41 ± 0·3 Ma; Golden, 1979)

10/694, 10/322: unnamed basaltic volcanic rocks (878372, 871793; Pleistocene; Carney, 1972)

KB145: Upper Basaltic Lavas, Emuruangogolak volcano (BM 20961699; >38 ka?–250 ± 100 years; Dunkley et al., 1993)

5/48: Murgisian basalt (762509; Pleistocene; <3·4 Ma; Webb, 1971)

61480, 61487: Flank fissure basalts, Silali volcano (120–63 ka; Macdonald et al., 1995; Smith et al., 1995)

56187, 56211, KB128: Katenmening Basalts, Silali volcano (11°2·0'S, 36°12·7'E; AM 18801230, AM 18681355; 63 ka; Macdonald et al., 1995; Smith et al., 1995)

61544: Post-caldera basalts, Silali volcano (1°10·4' S, 36°12·9'E; 7 ka; Macdonald et al., 1995; Smith et al., 1995)

Traverse 1
1/409, 1/609 (v), 1/609 (vi), 1/515, 1/567a: Saimo Basalt Formation (221665, 168762, 168762, 170670, 169765; Miocene; 20–15 Ma?; Martyn, 1969)

1/761: Eron Basalt (015562; Miocene; 8 Ma; Martyn, 1969)

1/858: Keiturin Basalt (Miocene; Martyn, 1969)

2/46, 2/16, 2/173, 2/156B: (Pliocene basalts; 8–4 Ma; Chapman, 1971)

13/706, 13/333: Kwaibus Basalt (685311, 289362; 0·95–0·89 Ma; Griffiths, 1977)

Reworked craton margin
410-1, 419-1: olivine melanephelinites, Loisiumurto volcanic rocks (1°06'S, 35°55'E; Baker et al., 1978; Crossley, 1979)

254-3 (augitite): Lisudwa volcanic rocks (1°58'S, 36°02'E; Baker et al., 1978; Crossley, 1979)

244-2: olivine melanephelinite, centre west of Lisudwa (1°58'S, 36°00'E; Crossley, 1979)

W298, W299: Simba basalts, Chyulu Hills (2°9·2'S, 37°36·7'E; 2°9·0'S, 37°34·3'E)

W642: Shaitani basalts, Chyulu Hills (2°57'S, 38°6'E)

KLR2, KLR17, KLR32: basalts and ferrobasalts, Ol Tepesi volcanic rocks (1°43'S, 36°42'E; Baker et al., 1977, 1978)

W173, W174: Singaraini basalts (BH 284882; Baker et al., 1978)

975-18, 976-6, 1009-4: Kordjya basalts (1°24'S, 36°10'E; 1°24'S, 36°10'E; 1°24'S, 36°10'E; Baker et al., 1978)

W181: Ol Keju Nero basalts (BJ 282295; Baker et al., 1978)

W185: basalt, Ol Esayeti volcanic rocks (BJ 315314; 6·7–5·6 Ma; Baker et al., 1978)

W86A, W153, W164 melanephelinites, Ngong volcanic series (BJ 371353, BJ 406390, BJ 430394; Baker et al., 1988)

W150; olivine melanephelinite, Ol Esakut (BJ 324310; Baker et al., 1978)

W151; olivine melabasalt, Ol Esayeti volcanic rocks (BJ 325307; Baker et al., 1978)

KLR289, KLR306, KLR151, KLR259, KLR311: alkali basalts, Group 1, Olokisalie (Olorgesailie) volcano (1°40'S, 36°30'E; Henage, 1977; Baker et al., 1988)

KLR134: Group 2, Olokisalie (Olorgesailie) volcano (1°40'S, 36°30'E; Henage, 1977; Baker et al., 1988)

KLR129, KLR196, KLR193: Group 4, Olokisalie (Olorgesailie) volcano (1°40'S, 36°30'E; Henage, 1977; Baker et al., 1988)

BB85-250: Akira Basalt Formation, Naivasha (AJ 951952; Clarke et al., 1990)

ND25, ND30, ND64, ND218: alkali basalts, Early Basaltic Series, Naivasha (Olkaria) Complex (AK 904169, AK 886176, AK 934163, AK 986127; Late Pleistocene; Davies & Macdonald, 1987)

ND124b, ND 226a: alkali basalts, Late Basaltic Series, Naivasha (Olkaria) Complex (AK 952165, AK 950165; Late Pleistocene; Davies & Macdonald, 1987)

EL1, EL3: alkali basalts, Elmenteita Volcanic Group (AK 915467, AJ 905467; Clarke et al., 1990)

W214: basaltic lava, east of Lake Naivasha (BK 200188; Baker et al., 1988, table 2)

SB24: Older Elmenteita basalt, near Gilgil (BK 039403; <0·45 Ma; Clarke et al., 1990)

W220, W205: Lower Simbara basalts, Aberdare Range (BK 822510, BK 972216; Upper Miocene; Baker et al., 1988)

W217: Upper Simbara basalts, Aberdare Range (BK 729260; Upper Miocene; Baker et al., 1988)

W223, W222, W224: Laikipian basalts, Aberdare Range (BK 532914, BK 751611, BK 525951)

W201, W202, W219: Thiba basalts, east of Aberdare Range (CK 075190, CK 075190, CK 978410; Lower Pleistocene)

THEL1: Theloi basalts, Eldama Ravine (0°05'N, 35°41'E; Upper Miocene?)

GOIT2: basalt, Goitumet volcano (329096; Upper Pleistocene)

W673, 43/1/S/38: Turasha basalts (BL 119463, BK 115460; Lower Pliocene?)

W329: Kasurein basalts (ZR 320642; Upper Pliocene)

W238: Kwaibus basalts, SSW of Lake Baringo (ZR 270431; Lower Pleistocene)

Craton
BD105: olivine melilite nephelinite, Loolmurwak volcano, Tanzania (2°47'S, 35°59'E; Dawson et al., 1985)

BD4171: olivine melilitite, Armykon Hill, Tanzania (2°36'S, 35°55'E)

BD4200: olivine melilitite, Labait Crater, Lake Hill, Tanzania (4°34'S, 35°26'E; Dawson et al., 1997)

932-3: alkali olivine basalt, Sambu Volcanic series, SW Kenya (2°02'S, 35°59'E; Baker et al., 1978; Crossley, 1979

	ACKNOWLEDGEMENTS

 
We thank Barry Dawson, Alex McBirney, Diana Sutherland and Laurie Williams for generously supplying specimens, Andrew Lloyd (Open University) for the diagrams, and Joerg Keller, Philip Leat, Anton le Roex and Marge Wilson for very helpful reviews. M.S. publishes with the permission of the Director, British Geological Survey. R.M. acknowledges support from NERC for work in Kenya over many years.

	FOOTNOTES

 
^*Corresponding author. E-mail: r.macdonald{at}lancaster.ac.uk

Extended dataset can be found at: http://www.petrology.oup.journals.org

	REFERENCES

TOP ABSTRACT INTRODUCTION GEOLOGICAL SUMMARY ANALYTICAL TECHNIQUES NOMENCLATURE AND DISTRIBUTION COMPOSITIONAL DIFFERENCES... BASALT COMPOSITION VS TIME:... COMPARISONS WITH OCEAN ISLAND... CRUSTAL CONTAMINATION MANTLE HETEROGENEITY MANTLE END-MEMBERS CONCLUSIONS APPENDIX : SAMPLE LOCALITY... REFERENCES

 
Achauer, U. & KRISP Teleseismic Working Group (1994). New ideas on the Kenya rift based on the inversion of the combined dataset of the 1985 and 1989/90 seismic tomography experiments. Tectonophysics 236, 305–329.

Bailey, D. K. (1992). Episodic alkaline igneous activity across Africa: implications for the causes of continental break-up. In: Storey, B. C., Alabaster, T. & Pankhurst, R. J. (eds) Magmatism and the Causes of Continental Break-up. Geological Society, London, Special Publications 68, 91–98.

Baker, B. H. (1987). Outline of the petrology of the Kenya rift alkaline province. In: Fitton, J. G. & Upton, B. G. J. (eds) Alkaline Igneous Rocks. Geological Society, London, Special Publications 30, 293–311.

Baker, B. H., Goles, G. G., Leeman, W. P. & Lindstrom, M. M. (1977). Geochemistry and petrogenesis of a basalt–benmoreite–trachyte suite from the southern part of the Gregory Rift, Kenya. Contributions to Mineralogy and Petrology 64, 303–332.[Web of Science]

Baker, B. H., Crossley, R. & Goles, G. G. (1978). Tectonic and magmatic evolution of the southern part of the Kenya rift valley. In: Neumann, E.-R. & Ramberg, I.B. (eds) Petrology and Geochemistry of Continental Rifts. Dordrecht: Reidel, pp. 29–50.

Baker, B. H., Mitchell, J. G. & Williams, L. A. J. (1988). Stratigraphy, geochronology and volcano-tectonic evolution of the Kedong–Naivasha–Kinangop region, Gregory Rift valley, Kenya. Journal of the Geological Society, London 145, 107–116.[Abstract/Free Full Text]

Bedini, R. M., Bodinier, J.-L., Dautria, J.-M. & Morten, L. (1997). Evolution of LILE-enriched small melt fractions in the lithospheric mantle: a case study from the East African Rift. Earth and Planetary Science Letters 153, 67–83.

Bell, K. & Dawson, J. B. (1995). Nd and Sr isotope systematics of the active carbonatite volcano, Oldoinyo Lengai. In: Bell, K. & Keller, J. (eds) Carbonatite Volcanism. Berlin: Springer, pp. 100–112.

Bell, K. & Dobson, M.H. (1980). The geochronology of the Tanzanian shield. Journal of Geology 89, 109–128.

Bell, K. & Simonetti, A. (1996). Nd, Pb and Sr isotope systematics from Oldoinyo Lengai carbonatite volcano (Tanzania): carbonatite magmatism and lithosphere–plume interaction. Journal of Conference Abstracts 1, 56.

Bellieni, G., Justin Visentin, E., Piccirillo, E. M. & Zanettin, B. (1987). Volcanic cycles and magmatic evolution in northern Turkana (Kenya). Tectonophysics 143, 161–168.[Web of Science]

Birt, C., Thybo, H., Maguire, P. K. H., Khan, M. A., Keller, G. R. & Patel, J. (1997). The influence of pre-existing structures on the evolution of the southern Kenya rift valley—evidence from seismic and gravity studies. Tectonophysics 278, 211–242.

Black, S., Macdonald, R., Barreiro, B. A., Dunkley, P. N. & Smith, M. (1998). Open system alkaline magmatism in northern Kenya: evidence from U-series disequilibria and radiogenic isotopes. Contributions to Mineralogy and Petrology 131, 364–378.

Bloomer, S. H., Curtis, P. C. & Karson, J. A. (1989). Geochemical variation of Quaternary basaltic volcanics in the Turkana Rift, northern Kenya. Journal of African Earth Sciences 8, 511–532.[Web of Science]

Burke, K. (1996). The African Plate. South African Journal of Geology 99, 341–409.[Web of Science]

Carney, J. N. (1972). The geology of the area to the east of Lake Baringo, Rift Valley Province, Kenya. Ph.D. thesis, University of London.

Chapman, G. R. (1971). The geological evolution of the northern Kamasia Hills, Baringo District, Kenya. Ph.D. thesis, University of London.

Chazot, G., Menzies, M. A. & Harte, B. (1996). Determination of partition coefficients between apatite, clinopyroxene, amphibole, and melt in natural spinel lherzolites from Yemen: implications for wet melting of the lithospheric mantle. Geochimica et Cosmochimica Acta 60, 423–437.[Web of Science]

Chesley, J. T., Rudnick, R. L. & Lee, C.-T. (1998). Longevity of cratonic mantle beneath an active rift: Re–Os evidence from xenoliths from the Tanzanian East African Rift. Extended Abstracts, 7th International Kimberlite Conference. Cape Town: University of Cape Town, pp. 149–151.

Clarke, M. C. G., Woodhall, D. G., Allen, D. & Darling, G. (1990). Geological, volcanological and hydrogeological controls on the occurrence of geothermal activity in the area surrounding Lake Naivasha, Kenya. Report. Ministry of Energy of Kenya, Nairobi, 138 pp.

Class, C. & Goldstein, S. L. (1997). Plume–lithosphere interactions in the ocean basins: constraints from the source mineralogy. Earth and Planetary Science Letters 150, 245–260.[Web of Science]

Class, C., Altherr, R., Volker, F., Eberz, G. & McCulloch, M. T. (1994). Geochemistry of Pliocene to Quaternary alkali basalts from the Huri Hills, northern Kenya. Chemical Geology 113, 1–22.

Coblentz, D. D. & Sandiford, M. (1994). Tectonic stresses in the African plate: constraints on the ambient lithospheric stress state. Geology 22, 831–834.[Abstract/Free Full Text]

Crossley, R. (1979). The Cenozoic stratigraphy and structure of the western part of the Rift Valley in southern Kenya. Journal of the Geological Society, London 136, 393–405.[Abstract/Free Full Text]

Crossley, R. & Knight, R. M. (1981). Volcanism in the western part of the rift valley in southern Kenya. Bulletin Volcanologique 44, 117–128.

Darling, W. G., Greisshaber, E., Andrews, J. N., Armannsson, H. & O’Nions, R. K. (1995). The origin of hydrothermal and other gases in the Kenya Rift Valley. Geochimica et Cosmochimica Acta 59, 2501–2512.

Davies, G. R. & Macdonald, R. (1987). Crustal influences in the petrogenesis of the Naivasha basalt–comendite complex: combined trace element and Sr–Nd–Pb isotope constraints. Journal of Petrology 28, 1009–1031.[Abstract/Free Full Text]

Dawson, J. B., Smith, J. V. & Jones, A. P. (1985). A comparative study of bulk rock and mineral chemistry of olivine melilitites and associated rocks from East and South Africa. Neues Jahrbuch für Mineralogie, Abhandlungen 152, 143–175.[Web of Science]

Dawson, J. B., James, D., Paslick, C. & Halliday, A. N. (1997). Ultrabasic potassic low-volume magmatism and continental rifting in north–central Tanzania: association with enhanced heat flow. Russian Geology and Geophysics 38, 69–81.

Dunkley, P. N., Smith, M., Allen, D. J. & Darling, W. G. (1993). The geothermal activity and geology of the northern sector of the Kenya Rift Valley. Research Report SC/93/1. Keyworth: British Geological Survey.

Ebinger, C. J. & Sleep, N. H. (1998). Cenozoic magmatism throughout east Africa resulting from impact of a single plume. Nature 395, 788–791.

Ebinger, C., Bechtel, T., Forsyth, D. & Bowin, C. (1989). Effective elastic plate thickness beneath the east African and Afar plateaus and dynamic compensation for the uplifts. Journal of Geophysical Research 94, 2883–2901.

Ebinger, C. J., Rex, D. C., Yemane, T. & Kelley, S. (1993). Volcanism and extension between the Main Ethiopian and Gregory rifts. In: Thorweihe, U. & Schandelmeier, H. (eds) Geoscientific Research in Northeast Africa. Rotterdam: Balkema, pp. 301–304.

Ebinger, C., Djomani, Y. P., Mbede, E., Foster, A. & Dawson, J. B. (1997). Rifting Archaean lithosphere: the Eyasi–Manyara–Natron rifts, East Africa. Journal of the Geological Society, London 154, 947–960.[Abstract/Free Full Text]

Fitton, J. G. (1995). Coupled molybdenum and niobium depletion in continental basalts. Earth and Planetary Science Letters 136, 715–721.

Fitton, J. G., James, D. & Leeman, W. P. (1991). Basic magmatism associated with late Cenozoic extension in the western United States: compositional variation in space and time. Journal of Geophysical Research 96, 13696–13711.

Fitton, J. G., Saunders, A. D., Larsen, L. M., Hardarson, B. S. & Norry, M. J. (1998). Volcanic rocks from the southeast Greenland margin at 63°N: composition, petrogenesis and mantle sources. In: Saunders, A. D., Larsen, H. C. & Wise, S. H. (eds) Proceedings of the Ocean Drilling Program Scientific Results, 152. College Station, TX: Ocean Drilling Program, pp. 331–350.

George, R. M. M., Rogers, N. W. & Kelley, S. (1998). Earliest magmatism in Ethiopia: evidence for two mantle plumes in one flood basalt province. Geology 26, 923–926.[Abstract/Free Full Text]

Golden, M. (1979). The geology of the area east of Silale, Rift Valley Province, Kenya. Ph.D. thesis, University of London.

Green, W. V., Achauer, U. & Meyer, R. P. (1991). A three-dimensional seismic image of the crust and upper mantle beneath the Kenya rift. Nature 354, 199–203.

Griffiths, P. S. (1977). The geology of the area around Lake Hannington and the Perkerra River, Rift Valley Province, Kenya. Ph.D. thesis, University of London.

Haase, K. M. & Devey, C. W. (1994). The petrology and geochemistry of Vesteris Seamount, Greenland basin—an intraplate alkaline volcano of non-plume origin. Journal of Petrology 35, 295–328.[Abstract/Free Full Text]

Halliday, A. N., Lee, D.-C., Tommasini, S., Davies, G. R., Paslick, C. R., Fitton, J. G. & James, D. E. (1995). Incompatible trace elements in OIB and MORB and source enrichment in the sub-oceanic mantle. Earth and Planetary Science Letters 133, 379–395.

Hardarson, B. S. & Fitton, J. G. (1991). Increased mantle melting beneath Snaefellsjokull volcano during Late Pleistocene glaciation. Nature 353, 62–64.

Hart, S. R. (1984). A large-scale isotope anomaly in the Southern Hemisphere mantle. Nature 509, 753–757.

Hart, S. R., Hauri, E. H., Oschmann, L. A. & Whitehead, J. A. (1992). Mantle plumes and entrainment: isotopic evidence. Science 256, 517–520.[Abstract/Free Full Text]

Hay, D. E., Wendlandt, R. F. & Keller, G. R. (1995). The origin of Kenya rift plateau-type flood phonolites: integrated petrologic and geophysical constraints on the evolution of the crust and upper mantle beneath the Kenya Rift. Journal of Geophysical Research 100, 10549–10557.[Web of Science]

Henage, L. F. (1977). The geology and compositional evolution of Mt. Olokisalie, Kenya. Ph.D. thesis, University of Oregon, Eugene.

Hendrie, D. B., Kusznir, N. J., Morley, C. K. & Ebinger, C. J. (1994). Cenozoic extension in northern Kenya: a quantitative model of rift basin development in the Turkana region. Tectonophysics 236, 409–438.

Hoernle, K. & Schmincke, H.-U. (1993). The petrology of the tholeiites through melilite nephelinites on Gran Canaria, Canary Islands: crystal fractionation, accumulation, and depths of melting. Journal of Petrology 34, 573–597.[Abstract/Free Full Text]

Hofmann, A. (1997). Mantle geochemistry: the message from oceanic volcanism. Nature 385, 219–229.

Hoffman, C., Courtillot, V., Féraud, G., Rochette, P., Yirgu, G., Ketefo, E. & Pik, R. (1997). Timing of the Ethiopian flood basalt event and implications for plume birth and global change. Nature 389, 838–841.

Kalt, A., Hegner, E. & Satir, M. (1997). Nd, Sr, and Pb isotopic evidence for diverse lithospheric mantle sources of East African Rift carbonatites. Tectonophysics 278, 31–45.

Keppler, H. (1996). Constraints from partitioning experiments on the composition of subduction-zone fluids. Nature 380, 237–240.

Last, R., Nyblade, A. A., Langston, C. A. & Owens, T. J. (1996). Lithospheric structure beneath the east African plateau from receiver functions and Rayleigh wave dispersion (abstract). EOS Transactions, American Geophysical Union 77, F463.

Latin, D., Norry, M. J. & Tarzey, R. J. E. (1993). Magmatism in the Gregory Rift, East Africa: evidence for melt generation by a plume. Journal of Petrology 34, 1007–1027.[Abstract/Free Full Text]

LeBas, M. J. (1987). Nephelinites and carbonatites. In: Fitton, J. G. & Upton, B. G. J. (eds) Alkaline Igneous Rocks. Geological Society, London, Special Publications 30, 53–83.

LeBas, M. J., Le Matre, R. W., Streckeisen, A. & Zanettin, B. (1986). A chemical classification of volcanic rocks based on the total alkali–silica diagram. Journal of Petrology 27, 745–750.[Abstract/Free Full Text]

Lippard, S. J. & Truckle, P. H. (1978). Spatial and temporal variations in basalt geochemistry in the northern Kenya rift. In: Neumann, E. R. & Ramberg, I. B. (eds) Petrology and Geochemistry of Continental Rifts. Dordrecht: Reidel, pp. 123–132.

Macdonald, R. (1994). Petrological evidence regarding the evolution of the Kenya Rift Valley. Tectonophysics 236, 373–390.

Macdonald, R., Williams, L. A. J. & Gass, I. G. (1994). Tectonomagmatic evolution of the Kenya rift valley: some geological perspectives. Journal of the Geological Society, London 151, 879–888.[Abstract/Free Full Text]

Macdonald, R., Davies, G. R., Upton, B. G. J., Dunkley, P. N., Smith, M. & Leat, P. T. (1995). Petrogenesis of Silali volcano, Gregory Rift, Kenya. Journal of the Geological Society, London 152, 703–720.[Abstract/Free Full Text]

Marty, B., Pik, R. & Gezahegn, Y. (1996). Helium isotopic variations in Ethiopian plume lavas: nature of magmatic sources and limit on lower mantle contribution. Earth and Planetary Science Letters 144, 223–237.

Martyn, J. E. (1969). The geological history of the country between Lake Baringo and the Kerio River, Baringo District, Kenya. Ph.D. thesis, University of London.

McClenaghan, M. P. (1971). The geology of the Ribkwo area, Baringo District, Kenya. Ph.D. thesis, University of London.

Mechie, J., Keller, G. R., Prodehl, C., Gaciri, S. J., Braile, L. W., Mooney, W. D., Gajewski, D. & Sandmeier, K.-J. (1994). Crustal structure beneath the Kenya Rift from axial profile data. Tectonophysics 236, 179–200.

Mechie, J., Keller, G. R., Prodehl, C., Khan, M. A. & Gaciri, S. J. (1997). A model for the structure, composition and evolution of the Kenya rift. Tectonophysics 278, 95–119.

Mengel, K. & Green, D. H. (1989). Stability of amphibole and phlogopite in metasomatised peridotite under water-saturated and water-undersaturated conditions. In: Ross, J. (ed.) Fourth International Kimberlite Conference. Australian Journal of Earth Sciences Special Publication 14, 571–581.

Mooney, W. D. & Christensen, N. I. (1994). Composition of the crust beneath the Kenya rift. Tectonophysics 236, 391–408.[Web of Science]

Morgan, P. (1983). Constraints on rift thermal processes from heat flow and uplift. Tectonophysics 94, 277–298.

Morley, C. K., Westcott, W. A., Stone, D. M., Harper, R. M., Wigger, S. T. & Karanja, F. M. (1992). Tectonic evolution of the northern Kenya Rift. Journal of the Geological Society, London 149, 333–348.[Abstract/Free Full Text]

Norry, M. J., Truckle, P. H., Lippard, S. J., Hawkesworth, C. J., Weaver, S. D. & Marriner, G. F. (1980). Isotopic and trace element evidence from lavas, bearing on mantle heterogeneity beneath Kenya. Philosophical Transactions of the Royal Society, London, Series A 297, 259–271.

Nyblade, A. A., Pollack, H. N., Jones, D. L., Podmore, F. & Mushayandebvu, M., (1990). Terrestrial heat flow in east and southern Africa. Journal of Geophysical Research 95, 17371–17384.

Nyblade, A. A., Birt, C., Langston, C. A., Owens, T. J. & Last, R. J. (1996). Seismic experiment reveals rifting of craton in Tanzania. EOS Transactions, American Geophysical Union 77, 517–521.

Paslick, C., Halliday, A., James, D. & Dawson, J. B. (1995). Enrichment of the continental lithosphere by OIB melts: isotopic evidence from the volcanic province of northern Tanzania. Earth and Planetary Science Letters 130, 109–126.

Potts, P. J., Williams-Thorpe, O., Issacs, M. C. & Wright, D. W. (1985). High precision instrumental neutron activation analysis of geological samples employing simultaneous counting with both planar and coaxial detectors. Chemical Geology 48, 145–155.

Rehkamper, M., Halliday, A. N., Barfod, D., Fitton, J. G. & Dawson, B. D. (1997). Platinum-group element abundance patterns in different mantle environments. Science 278, 1595–1598.[Abstract/Free Full Text]

Rogers, N. W., Macdonald, R., Fitton, J. G., George, R. W. W., Smith, M. & Barreiro, B. A. (2000). Two mantle plumes beneath the East African rift system; Sr, Nd and Pb isotope evidence from Kenya Rift basalts. Earth and Planetary Science Letters 176, 387–400.

Rudnick, R. L., McDonough, W. F. & Orpin, A. (1992). Northern Tanzanian peridotite xenoliths: a comparison with Kaapvaal peridotites and inferences on metasomatic interactions. In: Meyer, H. O. A. & Leonardos, O. (eds) Kimberlites, Related Rocks and Mantle Xenoliths, Volume 1. Proceedings of the 5th International Kimberlite Conference. Rio de Janeiro: Companhia de Pesquisa de Recursos Minerais, pp. 336–353.

Rudnick, R. L., McDonough, W. F. & Chappell, B. W. (1993). Carbonatite metasomatism in the northern Tanzanian mantle: petrographic and geochemical characteristics. Earth and Planetary Science Letters 114, 463–475.[Web of Science]

Scarsi, P & Craig, H. (1996). Helium isotope ratios in Ethiopian Rift basalts. Earth and Planetary Science Letters 144, 505–516.

Simiyu, S. & Keller, G. R. (1997). An integrated analysis of lithospheric structure across the East African plateau based on gravity anomalies and recent seismic studies. Tectonophysics 278, 291–313.

Slack, P. D., Davis, P. M. & KRISP Teleseismic Working Group (1994). Attenuation and velocity of P-waves in the mantle beneath the East African Rift, Kenya. Tectonophysics 236, 331–358.

Smith, M. (1994). Stratigraphic and structural constraints on mechanisms of active rifting in the Gregory Rift, Kenya. Tectonophysics 236, 3–22.

Smith, M. & Mosley, P. (1993). Crustal heterogeneity and basement influence on the development of the Kenya rift, East Africa. Tectonics 12, 591–606.[Web of Science]

Smith, M., Dunkley, P. N., Deino, A., Williams, L. A. J. & McCall, G. J. H. (1995). Geochronology, stratigraphy and structural evolution of Silali volcano, Gregory Rift, Kenya. Journal of the Geological Society, London 152, 297–310.[Abstract/Free Full Text]

Sobolev, S. V., Zeyen, H., Stoll, G., Werling, F., Altherr, R. & Fuchs, K. (1996). Upper mantle temperatures from teleseismic tomography of French Massif Central including effects of composition, mineral reactions, anharmonicity, anelasticity and partial melt. Earth and Planetary Science Letters 139, 147–163.

Späth, A., Le Roex, A. P. & Duncan, R. A. (1996). The geochemistry of lavas from the Comores Archipelago, western Indian Ocean: petrogenesis and mantle source region characteristics. Journal of Petrology 37, 961–991.[Abstract/Free Full Text]

Sun, S.-S. & McDonough, W. F. (1989). Chemical and isotopic systematics of oceanic basalts: implications for mantle composition and processes. In: Saunders, A. D. & Norry, M. J. (eds) Magmatism in the Ocean Basins. Geological Society, London, Special Publications 42, 313–345.

Taylor, S. R. & McLennan, S. M. (1985). The Continental Crust: its Composition and Evolution. Oxford: Blackwell, 312 pp.

Thompson, R. N. & Gibson, S. A. (1994). Magmatic expression of lithospheric thinning across continental rifts. Tectonophysics 233, 41–68.

Truckle, P. H. (1977). The geology of the area to the south of Lokori, South Turkana, Kenya. Ph.D. thesis, University of London.

Truckle, P. H. (1980). Variation of basic lava petrology of the Kenya Rift Valley. In: Carelli, A. (president) Geodynamic Evolution of the Afro-Arabian Rift System. Rome: Accademia Nazionale dei Lincei, pp. 133–142.

Webb, P. K. (1971). The geology of the Tiati Hills, Rift Valley Province, Kenya. Ph.D. thesis, University of London.

You, C.-F., Castillo, P. R., Gieskes, J. M., Chan, L. H. & Spivack, A. J. (1996). Trace element behaviour in hydrothermal experiments: implications for fluid processes at shallow depths in subduction zones. Earth and Planetary Science Letters 140, 41–52.[Web of Science]

Zeyen, H., Volker, F., Wehrle, V., Fuchs, K., Sobolev, S. V. & Altherr, R. (1997). Styles of continental rifting: crust–mantle detachment and mantle plumes. Tectonophysics 278, 329–352.

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