In her RAS Presidential Addressfor 2005, Kathryn A Whalerdiscusses how mapping theelectrical resistivity of thesubsurface can be used toinvestigate continental rifting,with examples from Africa.

The magnetotelluric (MT) methodinvolves recording time series of the mag-netic and electric (telluric) fields in

orthogonal, horizontal directions often, for con-venience, magnetic north–south and east–west,for analysis in the frequency domain. The stud-ies reported here are based primarily on datacollected in the audio-frequency range of a fewhundred Hz to of order 0.001 Hz, but in somecases extending to longer periods. Above ~1 Hz,the magnetic field source is sferics associatedwith thunderstorms; at longer periods, solaractivity. The magnetic field is trapped in theionospheric wave guide encircling the Earthand, sufficiently far from the original source(e.g. not adjacent to a thunderstorm), can beapproximated as a harmonic, plane wave atinfinity. Variations in the induced fields thatarise from the different lithologies and struc-tures in the subsurface can be mapped from themeasured MT time series.

The challenge in using MT in geophysics is todeduce useful information about Earth structurefrom these time series. The electric and magneticsignals are related by Z, an impedence tensorknown as the Earth response (also responsefunction or transfer function), which dependson variations in the resistivity of the area sur-veyed, reflecting, in turn, the subsurface struc-ture. The MT method involves modelling thedata to get the best fit for resistivity in one, twoand three dimensions (see “The magnetotelluricmethod” p2.30). Complications arise becausethe source field does not take the simple formassumed and lacks energy in key sections of thefrequency range, and from the effects of noise.

Data acquisitionEven when the source strength is good, the sig-nals measured in MT are extremely small; thusthe equipment must be very sensitive and thesignals are amplified considerably, making themvulnerable to noise. The most significant noisesources are energy at mains frequency and its

harmonics, cultural noise from the movement ofvehicles, people and animals, and wind. There-fore equipment must be deployed in quiet areas– both electromagnetically, so keeping wellaway from power lines and buildings with elec-trical power (be it generator or mains supplied),and vibrationally, avoiding trees or dense vege-tation, and roads. The magnetic sensors are usu-ally buried to reduce wind noise, but this doesnot reduce wind-caused tree-root vibration.Away from the main population centres inAfrica, it is usually easy to avoid electro-magnetic and vehicle noise, but vegetation covercan cause problems.

The electric field, E, is deduced from the volt-age between pairs of electrodes; separations of30–100 m are typical. At higher frequencies innon-arid conditions, metal electrodes are ade-quate, but non-polarizing porous pot type elec-trodes are used for longer periods and where drysoil conditions would make it difficult to makegood electrical contact with the ground. Lead –lead chloride or silver – silver chloride electrodesare stable and maintenance-free over long peri-ods. At high frequencies, the magnetic field, H,is monitored by induction coils consisting ofmany turns of fine wire on a metal core. Atlonger periods, fluxgate magnetometers aremore suitable, although broadband coils arenow available for periods of up to several thou-sand seconds.

I have used equipment loaned from the Nat-ural Environment Research Council’s Geophys-ical Equipment Facility (formerly GeophysicalEquipment Pool): SPAM (Short Period Auto-matic Magnetotelluric) Mk II or III for shorterperiod data acquisition (Ritter et al. 1998),incorporating sampling over a number of fre-

quency bands at different rates, real-time dataprocessing and quality control (though post-processing was also undertaken), and long-period systems based on EDA fluxgatemagnetometers, sampling at 20 s. SPAM IIIincorporates multiple band simultaneousrecording using parallel computing methodsand, latterly, GPS-based timing to an accuracyof ±2 µs, allowing for remote reference process-ing (described in “The magnetotelluricmethod”, page 2.30). Typically, SPAM isdeployed at sites for a day, and long-period sys-tems for about a month.

ModellingA variety of modelling algorithms, both forwardand inverse, have been developed to treat theproblem of inferring resistivity structure fromMT data. The inverse problem is highly non-unique and nonlinear, which causes difficultiesof stability and computational effort. The non-uniqueness means models are usually subject tosensitivity analysis, whereby the robustness oftheir features is tested by determining whetherthey are required by the data. Parker (1980)developed an important existence algorithm forthe 1-D case. He showed that the best-fittingmodel is a series of Dirac delta functions in con-ductivity, separated by insulating layers, referredto as a D+ model. The number, conductance anddepth of the delta functions are determined bythe data. The existence or otherwise of a 1-Dmodel is then determined by the misfit to the D+

model. D+ was developed for modelling thecomplex admittance (E/iωH, where ω is fre-quency); subsequently, Parker and Booker(1996) developed the equivalent algorithm,known as ρ+, for apparent resistivity and phase

2.28 A&G • April 2006 • Vol. 47

WHALER: PRESIDENTIAL ADDRESS

The magnetotelluric (MT) method is a way of probing the electrical resistivity (or its inverse,electrical conductivity) distribution of the subsurface. It is based on the principle ofelectromagnetic induction, in which natural magnetic field sources external to the Earthinduce eddy currents within it whose geometry and decay depend on the subsurfaceresistivity structure. The method is therefore entirely passive, and it can be employed inenvironmentally sensitive areas. The resistivity of the subsurface depends on itstemperature, mineralogy and fluid content, both interstitial and partial melt, and is sensitiveto even small amounts of melt and residual fluids and its connectivity. Thus it can be a gooddiscriminant between different rock types, and complements information available fromother geophysical methods. Here, I outline the basics of the MT method, indicate how dataacquisition, processing and modelling are undertaken, and illustrate the uses of MT toinvestigate tectonic and structural problems through three case studies from Africa.

ABSTRACT

Magnetotelluric fieldwo

data (see “The magnetotelluric method”, page2.30). At the opposite extreme from theextremely “spiky”, non-physical D+ and ρ+

models, most interpretable models are producedby seeking the spatially smoothest, or regular-ized, model fitting the data, a technique firstintroduced in the 1-D case by Constable et al.(1987) and known as Occam modelling. Thusthe objective function minimized consists of aterm measuring the goodness-of-fit and a termmeasuring model smoothness, with a freeparameter determining the relative importanceassigned to the two. The method is iterative, sorequires a starting model; typically, this will be auniform half-space of specified resistivity. Themodels are over-parameterized, allowing a goodfit to the data, and the smoothness constraintminimizes the amount of structure in the model.We then know that the real Earth has at least asmuch structure as the Occam model. “Smooth-ness” is usually characterized by a two-normmeasure of the gradient or second (spatial)derivative of the model (approximated by firstor second differences).

There are various implementations of regular-ized modelling in two and three dimensions.Two-dimensional Occam (de Groot-Hedlin andConstable 1990) and rapid relaxation inversion,or RRI (Smith and Booker 1991) are the mostwidely used 2-D algorithms. Over the years,guidelines have been developed for efficientcomputational implementation, including opti-mal gridding of the subsurface, and how toincorporate the data. For instance, joint inver-sion of equally weighted TE and TM mode data(defined in “The magnetotelluric method” page2.30) from a uniform half-space model tends togive a model dominated by structure derivedfrom the TM mode response (e.g. Livelybrookset al. 1993). These iterative methods require pre-dictions of the data by the current model as partof the computational process, for which efficientalgorithms have been developed using bothfinite difference (e.g. Swift 1967, Brewitt-Taylorand Weaver 1976) and finite element (e.g. Wan-namaker et al. 1987) techniques.

In the 3-D case, Maxwell’s equations do notseparate into independent modes, and all fourelements of the impedance tensor must beincluded in modelling. The computational effortis therefore much greater, and only recently has3-D inversion become routinely practical. Hau-tot et al. (2000) developed a regularized 3-Dinversion algorithm, based on the forward mod-

A&G • April 2006 • Vol. 47 2.29

WHALER: PRESIDENTIAL ADDRESS

rk adventures in Africa

28 29 30 31

–17

–16

ZIMBABWE

ZAMBIA

MOZAMBIQUE

LakeKariba

Lake CaboraBassa

N

Mana PoolsMT profile

River Zambezi

River Sanyati

River Sengwa

Rive

r Man

yam

e

River Kafue

internationalboundaries

Mana Poolsseismic line

BGRMTprofile

Lower Zambeziseismic line

MzarabaniFault

1: Karoo sedimentary basins within the mobile belts of southern Africa. Basins 1–4 lie within theDamara orogenic belt, which extends west of the map area into Namibia, and 3 and 4 are the basinsstudied here. The contours indicate approximate basement. Basins are: 1 Passarge, 2 Mid-Zambezi,3 Mana Pools, 4 Lower Zambezi, 5 Western Zambian, 6 Kafue, 7 Launo, 8 Luangwa. The Sijarira horst(S.H.) separates sub-basins within the Mid-Zambezi basin, and the Chewore Inlier (C.I.) separatesthe Mana Pools and Lower Zambezi basins. After Orpen et al. (1989).

2: Relative locations of the Mana Pools and Lower Zambezi MT and reflection seismic profiles. Dot-dash lines mark the approximate basin boundaries. Dashed lines mark international borders.

elling difference equations method of Mackie etal. (1993). Further complications can occur ifconductivity is anisotropic, and distinguishingbetween 3-D and anisotropic subsurface con-ductivity distributions is particularly difficult.Most treatments have been concerned with atmost 2-D anisotropic media (e.g. Pek andVerner 1997).

The electromagnetic fields satisfy a diffusionequation, rather than the wave equation appro-priate for seismology. This means the boundaryconditions for 2- or 3-D modelling have to beapplied carefully, and that structures some dis-tance from a site influence the MT data mea-sured. Thus the area or volume over which a 2-or 3-D model is parameterized must be signifi-cantly larger than that encompassed by the sites.This adds further to the computational burden.

Case studiesHere, I present examples from three (previouslypublished) MT case studies from Africa. The

first is from the Zambezi Valley in northernZimbabwe, where the aim was to infer sedimen-tary thickness in the Karoo basins, and establishwhether an extremely good subsurface conduc-tor running through much of the Damara oro-genic belt in southern Africa was continuous.The second is a 3-D investigation of the sedi-mentary history of the Baringo-Bogoria regionof the Kenyan rift. Finally, the power of a multi-disciplinary study is demonstrated in the inter-pretation of a profile across the northern mainEthiopian rift.

Zambezi Valley, northern ZimbabweA simplistic description of the African continentdivides it into stable cratonic blocks and mobilebelts. Rifting between the late Permian and earlyTertiary led to the formation of several Karoobasins within the mobile belt terranes, whosetectonic and structural development is describedby Orpen et al. (1989). A number of previousstudies (e.g. Haak and Hutton 1986 and refer-

ences therein, Losecke et al. 1988) had estab-lished that there is an extremely good electricalconductor beneath the Damara orogenic belt inNamibia, Botswana, Zambia and parts of Zim-babwe (figure 1). The Mana Pools portion of themobile belt in the Zambezi Valley of northernZimbabwe was a “missing link”, since its elec-trical structure had not been probed before1987, when our study began. In addition, aero-magnetic and gravity data had been used toinvestigate the thickness of sediments in both theMana Pools and Lower Zambezi (sometimesreferred to as Cabora Bassa) basins in the Zam-bezi valley, but the interpretation of the aero-magnetic data was ambiguous (Bosum andGeipel 1988). Although the original explorationaim of investigating the basins was uraniumprospecting, there was also interest in theirhydrocarbon potential, and a reflection seismicsurvey was shot (Hiller and Buttkus 1996).Unfortunately, both basins straddle inter-national boundaries (with Zambia and Mozam-

2.30 A&G • April 2006 • Vol. 47

WHALER: PRESIDENTIAL ADDRESS

The theory behind electromagnetic (EM)induction methods, including MT, is based on(non-relativistic) Maxwell’s equations,assuming that no dielectric or magnetic materialis present, and that permittivity andpermeability take their free-space values(independent of frequency). Assuming harmonictime dependence for both the magnetic field Hand electric field E, and transforming into thefrequency domain, both H and E satisfyHelmholtz equations ∇2H – γ2H = 0 and ∇2E – γ2E = 0 where γ2 = iωσµ, with ω angularfrequency, σ electrical conductivity and µpermeability (neglecting displacement currents,which are generally much smaller thanconduction currents at the frequencies employedin MT [Telford et al. 1990]). These are diffusion(rather than wave) equations, which indicate oneof the limiting factors of the MT method – itsinherent lack of resolution. Fields are attenuatedduring propagation and the field strength is largeenough to measure only up to some penetrationdepth; the diffusive nature of the EM signalssmoothes out Earth structure (in both thevertical and lateral directions).

The 1-D solution to these equations can beshown to have an electric field varying asEx = E0eiωte–γz with x in a horizontal directionand z positive downwards. E0 is the incidentelectric field at the Earth’s surface. This inducesan orthogonal horizontal magnetic fieldHy = H0ei(ωt–φ)e–γz. Hy lags behind Ex with a phasedifference φ between the two fields.

An EM field will be reduced in amplitude by afactor 1⁄e at a distance δ within the Earth,known as the skin depth, given by

δ == 503where f is the frequency in Hz (e.g. Vozoff1972). Thus the MT penetration depth dependson both the period of the signal being measuredand the conductivity of the medium. The longerthe period measured or the lower theconductivity of the medium, the deeper the EMfields penetrate. The MT impedance Z isdefined as the ratio

Z = HEx

y = eiϕ (1)

Hence

Z = 1

2+ i (2)

Thus in a homogeneous Earth the compleximpedance Z is independent of frequency; itsmagnitude determines the conductivity, and thephase difference φ is 45°.

In a 2-D Earth, Maxwell’s equations decoupleinto two sets of equations. One set, known asthe E-polarization or transverse electric (TE)mode, relates the field components Ex, Hy andHz observed when Ex is parallel to the structureand in the direction of constant conductivity,called the geoelectrical strike direction. Theother set, the H-polarization or transversemagnetic (TM) mode, relates Hx, Ey and Ez

when the currents are perpendicular togeoelectrical strike. The impedances for the twomodes are different because conductivity is nolonger homogeneous.

In general, the horizontal electric andmagnetic field ratios define an impedancetensor, Z, satisfying

E = ZH (3)or

Ex= Zxx ZxyHxEy Zyx Zzz Hy

Z is also known as the Earth response, responsefunction, or transfer function. It is far fromstraightforward to obtain good estimates of theimpedance tensor elements from measured timeseries of the electric and magnetic fields, forseveral reasons. The source field is not theassumed harmonic plane wave at infinity, andsuffers from lack of energy, particularly in theso-called “dead band” at around 1–10 s period.Even with careful data acquisition, the data arecontaminated (biased) by non-zero mean, non-Gaussian noise. Thus robust methods oftransfer function estimation are employed (e.g.Egbert and Booker 1986, Chave and Thomson1989). In addition, if two or more sites recordsynchronously, their transfer functions can beestimated by a cross-power spectral methodknown as remote referencing (Gamble et al.1979), by assuming the signal is correlatedbetween the two sites but the noise is not.Remote reference estimation can be used toobtain usable transfer functions even when thedata suffer severe noise contamination in somecircumstances.

Comparing (1) and (3) it can be seen that, in a1-D Earth, Z takes the form

Z1D = 0 Z–Z 0

where Z is given by (2) (e.g. Zhang et al. 1987).The sign change indicates that the phase is in thethird quadrant for the equation relating Ey toHx. As already noted, the equations relating theelectric and magnetic fields decouple for a 2-D

ωµσ

E0H0

1fσ

2ωµσ

The magnetotelluric method

bique), and the geophysical information is con-fined to Zimbabwe.

We collected audio-frequency MT data at 11sites along a profile through the Zimbabwe por-tion of the basin (which comprises most of it)and onto the Zimbabwe craton to the south; atfive sites, longer period data were also collected(figure 2). Robust response function estimationproduced acceptable values at most sites,although the electric and magnetic field timeseries were insufficiently coherent at longeraudio-frequency MT periods and the shortestperiods obtainable from the long-period datafor usable data to be obtained. This left a gapbetween periods of about 10 and 100 s, so the ρ+

modelling technique was used to check that thedata were mutually consistent. Both Groom-Bailey and Bahr distortion analyses were under-taken, which indicated that the data werereasonably consistent with a 2-D structureapproximation (Bailey 1998), with a geoelectri-cal strike direction of approximately 80° appro-

priate for most periods at most sites. Two-dimensional RRI of the data was under-

taken (Bailey 1998, Bailey et al. 2000a). Theoptimum model has a good conductor to depthsof between 5 and 10 km beneath the valley,thinning and becoming less conductive to thesouth, a resistor at the surface beneath the Zim-babwe craton, and a relatively uniform base-ment resistivity beneath both the valley andcraton. The shape of the good conductorbeneath the valley corresponds reasonably wellwith the basin structure inferred from the seis-mic reflection survey shot slightly to the west(Hiller and Buttkus 1996, see figure 2), whichalso indicated that the southern boundary faultof the basin is steeply dipping to the north andreaches a depth of at least 7 km. We expected asimilar resistivity contrast at the northernescarpment in Zambia about 10 km north ofour most northerly site. With the boundingfaults included as vertical discontinuities(across which the smoothness constraint was

not applied), the fit to the data at stations closeto the basin edges was significantly improved,both numerically and in terms of matching theshape and trend of the data curves. The resis-tivity of the material to the north of thenortherly bounding fault was set to 2000 Ωm(higher than that inferred for the Zimbabwecraton to the south); increasing it improved thefit to the TE mode apparent resistivity, but thefit to the phase deteriorated, and vice versawhen the resistivity was decreased. Furtherimprovement was obtained when only thephase data from the most southerly site on thecraton were included in the inversion; there wasa suggestion that the apparent resistivity datawere static shifted. The structure of the finalgeoelectric section (figure 3) is in good agree-ment with the seismic reflection model (Hillerand Buttkus 1996) and supports Orpen et al.’s(1989) interpretation of the basin as a half-graben with an asymmetric cross-section and adepocentre offset to the northwest. The mater-

A&G • April 2006 • Vol. 47 2.31

WHALER: PRESIDENTIAL ADDRESS

Earth, with the electric field component Ex onlyinducing an orthogonal magnetic field Hy, andsimilarly Ey inducing the orthogonal Hx. Thusthe 2-D impedance tensor also has only non-zero off-diagonal elements, but now theirmagnitudes will differ:

Z2D = 0 Zxy (4)Zyx 0

(e.g. Zhang et al. 1987). In the general, 3-Dcase, all elements of the impedance tensor arenon-zero. This is also true for a 2-D Earth if thegeoelectrical strike direction does not coincidewith one of the horizontal co-ordinate axes but,in this case, the impedance tensor can be rotatedinto horizontal co-ordinates parallel andperpendicular to strike: Z′(θ) = R(θ)ZRT(θ)where R is the rotation matrix

R = cosθ sinθ–sinθ cosθ

and θ is the anticlockwise angle of rotation (e.g.Swift 1967). In this new reference frame, thediagonal elements of Z′ vanish. Thecomponents of the impedance tensor can beexpressed in terms of a quantity with units ofresistivity, based on the amplitude, and phase:

ρaij(ω) =µωZ

1

ij(ω)2 (5)

ϕij(ω) = arctanRIm

eZZ

i

i

j

j

((ωω

))

(6)

ρa is known as the apparent resistivity, and is theactual resistivity of a uniform Earth (as can bededuced from equation [2]). MT data are oftenpresented as apparent resistivity and phase as afunction of period, as plots from individual

stations, or, for data along (or projected onto) aprofile, as data pseudo-sections.

In general, there is no rotation angle θ forwhich the diagonal elements of the impedancetensor vanish, even if the data have been collectedover an approximately 2-D region. Mostalgorithms seek the rotation angle for which theyare a minimum. The extent to which anyminimum is well-defined indicates how good the2-D approximation is. For a 1-D Earth, thediagonal elements vanish for all rotation angles,and for a 3-D Earth, for no rotation angle.However, this angle can be a function of period,indicating a change in strike direction with depth,and can change from station to station. For the2-D approximation to be realistic, there shouldbe a well-defined rotation angle for which thediagonal elements are small that is approximatelyconstant for the survey and does not varysignificantly with period. In such circumstances,the data can be rotated into the new co-ordinatesystem, and the diagonal impedance tensorelements ignored. It is also necessary todetermine which of the two off-diagonalimpedance tensor components corresponds tothe TE mode and which to the TM mode (there isa 90º ambiguity in θ). There are severalrotational invariants that can be calculated fromimpedance tensor components (see e.g. Simpsonand Bahr 2005 for details). Quantities such as thedeterminant or Berdichevsky average are usefulfor 1-D modelling, since they are thought to beless affected by noise and distortion. Others areuseful indicators of the (minimum) dimension ofEarth structure.

The 2-D approximation most often fails due tolocal 3-D distortion, frequently galvanic, by

small-scale, near-surface inhomogeneities. Thiscan be tested for and the distortion removed (asfar as possible) to leave the signal resulting fromthe dominant 2-D structure. Several methodshave been proposed to undertake this tensordecomposition and assessment. A commonlyapplied technique is that of Groom and Bailey(1989), who resolved the distortion of an under-lying 2-D structure into parts describing twist,shear and anisotropy which have a geophysicalinterpretation. Bahr (1991) showed how the datacan be divided up into classes defined by thedegree of distortion, which indicate the type ofdecomposition (based on an earlier parameter-ization of the distortion tensor [Bahr 1988]) andsubsequent modelling that is appropriate. Bahr(1991), Smith (1995) and Balasis et al. (1997)showed that Groom and Bailey’s (1989)decomposition of the distortion tensor isequivalent to Bahr’s (1988) parameterization.

In the field, galvanic distortion may be revealedby static shift, a DC offset between the apparentresistivity curves at the shortest periods in thetwo orthogonal directions. Correcting for staticshift, i.e. moving one or both of the curves up ordown such that they coincide at short periods,can be achieved using independent data thatprovide a model of the shallow resistivitystructure. Most commonly, this information isobtained from DC resistivity sounding or thetransient electromagnetic method (TEM) (e.g.Meju 1996). Alternatively, it is possible to solvefor a static shift term in conjunction withinversions for subsurface structure. Anotherpossibility is to neglect the apparent resistivitydata and model only the phase information,which is unaffected by static shift.

2.32 A&G • April 2006 • Vol. 47

WHALER: PRESIDENTIAL ADDRESS

ial beneath has a resistivity of about 20 Ωm,which is low for basement material, but nolocalized good conductor similar to that seenelsewhere beneath the Damara orogenic belt isobserved. However, this may be because of thedata gap at periods between 10 and 100 s,which corresponds to depths around the top ofthe basement.

A much smaller portion of the adjacent LowerZambezi basin lies within Zimbabwe, althoughit is of larger extent. Losecke et al. (1988) pro-duced 2-D models of profiles across it from theirMT survey; the example in figure 4 is from the

profile shown on figure 2. They used a combi-nation of 1-D inversion and 2-D forward mod-elling, noting the results of the earlieraeromagnetic survey. Their model is divided intothree layers of broadly similar resistivity, whosedepths and thicknesses vary beneath each site;the model therefore has some lateral smooth-ness, but this was not a modelling requirement.In this case, there is a very large differencebetween the apparent resistivities of the twomodes within the Zambezi Valley, with the TMmode apparent resistivity curves having a steepnegative slope at periods beyond around 2 s and

dropping to extremely small values, and the TEmode apparent resistivity levelling out atbetween 10 and 100 Ωm (see Bailey et al.2000b). The penetration depths of the twomodes are thus very different. Losecke et al.’s(1988) model is dominated by the TM mode,such that an extremely good conductor (resis-tivities in the range 0.2–1.6 Ωm) begins atdepths of between 5 and 10 km in the valley andcontinues to the maximum depth extent of themodel. The subsequent seismic survey of Hillerand Buttkus (1996) indicates that the lower ofthe two magnetic horizons inferred by Bosumand Geipel (1988) (figure 4) corresponds to thebase of the upper Karoo. Thus the good con-ductor straddles the seismic basement, indicat-ing that, unlike the situation in the Mana Poolsbasin, it is a basement feature, as elsewherealong the Damara orogenic belt. Above thegood conductor is a resistive (1000–3000 Ωm)wedge, thinning from south to north, with amore conductive surface layer. However,Losecke et al.’s (1988) model is a poor fit to theTE mode data, so we investigated alternativemodels (Bailey 1998, Bailey et al. 2000b).Unfortunately, we only had access to the off-diagonal tensor component elements in the geo-electrical strike co-ordinate system determinedby Losecke et al. (1988), so were unable toundertake tensor decomposition and dimen-sionality analysis, or estimate distortion effects.There were indications of static shift in the data(such that only phase data were included fromone site during inversion). Inverting the twomodes separately with RRI gave two very dif-ferent models (figure 5), with that from the TMmode data having a good conductor similar toLosecke et al.’s (1988), and that from the TEmode data reproducing its resistive wedge.However, there are regions of the subsurfacethat are conductive in the TM mode model andresistive in the TE mode model, indicating thedifficulty in fitting the two modes simultane-ously. Both models had a much higher resistivityfor the Zimbabwe craton than Losecke et al.’s(1988); the enormous contrast with the veryconductive region in the valley presents numeri-cal difficulties in modelling (the iterative processdoes not fully converge). For the joint inversion,we included a vertical escarpment fault acrosswhich the smoothness constraint was not imple-mented; the final model (figure 5) fits the TEmode data better overall and is closer in appear-ance to the model from those data alone. This isquite unusual – the TM mode structure oftendominates in joint inversions (Livelybrooks etal. 1993). A comparison of the data predictionsindicates that the fits of ours and Losecke et al.’s(1988) models are similar. Both models have sig-nificant discrepancies from the seismic model,and attempts to characterize seismic layers withlayers of constant resistivity were unsuccessful(Bailey 1998, Bailey et al. 2000b). The lower of

borderMozambique

0

5

10

15

20

5.6

level 2

2

1

41.3458.43 8.4

N S

zone 1

zone 2

zone 3

dept

h (k

m)

8

8000

80

30

3000

80

escarpment

ZIMBABWE

LZ1 LZ2 LZ3 LZ4 LZ5 LZ6 LZ7 LZ8

2000 30001500 1000 2000 1300 1000

10magnetic level 1

4.6

0.2

1.1

1.31.60.60.6

0 6.1 12.7 18.5 23.4 27.5 33.2 38.9 46.3 49.2–4.1

Zambezi Valley

distance from site LZ1 in km

craton

3: The resistivity structure of the Mana Pools basin, with vertical breaks in structure allowed acrossthe locations of the northern and southern escarpment faults. Triangles and circles represent theTM and TE polarization data points plotted at one skin depth respectively, indicating the penetrationdepth of the data. Red points were down-weighted in the inversion because of poor coherency.

4: The resistivity structure of the Lower Zambezi basin after Losecke et al. (1988). Figures indicateresistivity in Ωm. Magnetic levels 1 and 2 indicate the positions of the two horizons deduced from anaeromagnetic survey of the area by Bosum and Geipel (1988).

A&G • April 2006 • Vol. 47 2.33

WHALER: PRESIDENTIAL ADDRESS

the two magnetic horizons interpreted byBosum and Geipel (1988) corresponds to thebase of the upper Karoo in the seismic model(figures 3 and 4). Losecke et al. (1988) inter-preted their resistive wedge as a layer of sedi-ments and intercalated sills within the basin. Ifthe intercalations are thin north–south orienteddykes instead, they would be invisible seismi-cally, and give rise to macroscopic resistivityanisotropy, explaining the separation betweenthe two MT modes, which occurs in the phase aswell as apparent resistivity (Kellett et al. 1992).Alternatively, an oblique contact between theresistive craton and conductive valley preservestheir separation over a longer lateral distancethan the vertical fault introduced into our RRImodel (figure 5) (which is easier to implementnumerically) without introducing anisotropy(see figure 17 of Bailey et al. 2000b).

More insight into the origin of the conductiv-ity anomalies beneath the Damara orogenic belthas come from a recent high-resolution MT sur-vey in Namibia (Ritter et al. 2003). This surveysuggests that the conductor there consists oftwo narrow, subvertical upper to mid-crustallevel zones that correlate with the major tec-tonic boundaries. Ritter et al. (2003) interpretthe high conductivity in terms of graphite (orother mineralization) enrichment along theshear zones.

Baringo-Bogoria basin, KenyaThe Baringo-Bogoria basin (BBB) lies within theeastern (Kenya) branch of the East African rift.The lakes to the north (Baringo) and south(Bogoria) are currently fresh water and alkaline,respectively, and both have hydrothermal fea-tures such as hot springs. In the past, the lakeshave been connected, and they have switchedbetween being fresh water and alkaline. Sedi-ment deposition has been controlled by two dif-ferent structural trends: N–S border faults andNW–SE transverse zones. Alluvial fans havedeveloped adjacent to border faults, and otherdeposition has been by river channels and inlakes. Several episodes of volcanism haveoccurred, most extensively in the early/mid-Miocene. The resulting subsurface resistivitypattern was expected to be 3-D, so a 3-D MTsurvey was carried out, with a station spacing oftypically 2 km over the central part of an area ofapproximately 15 × 25 km. MT data were col-lected at 24 sites; at these and a further sevensites, DC resistivity soundings were also carriedout to determine near-surface structure. The tec-tonic setting and station distribution are shownin figure 6. This and subsequent figures in thissection are reproduced from Hautot et al.(2000), which gives more details of all aspectsof the project.

The impedance tensor components werederived using the robust remote reference methodof Chave and Thomson (1989). Although we

5: The resistivity structure of the Lower Zambezi basin from RRI. (a) From TE polarization data only;(b) from TM polarization data only; (c) from both polarizations simultaneously, with a southernescarpment fault added. Triangles and circles indicate the TM and TE polarization data point skindepths respectively.

(a)

(b)

(c)

undertook 3-D modelling, tensor decompositionwas carried out using the method of Counil et al.(1986), which revealed two dominant trends:N120°–140° at short periods, and N180°–200°at longer periods, corresponding to the transferzone and main rift axis trends, respectively (fig-ure 6). This also enabled us to derive (using RRI)a representative 2-D model from MT sites alongan east–west profile which was used as a startingmodel for the deeper structure in 3-D inversion,overlying a 3-D resistivity distribution derivedfrom the DC resistivity soundings for the shallowstructure. Three-dimensional modelling requirescareful parameterization of the region studied,and consideration of the effects of the topogra-phy of the rift flanks. The horizontal and verticalgrid layout (with resistivity of each block theparameter to be determined) is shown in figure7. A central 3-D portion of the model volumecovers the east–west profile of the 2-D RRImodel (16.5km wide) and extends 25.5km in thenorth–south direction; the depth extent was ini-tially 8km, but increased to 17km during inver-sion. The size of the blocks in the horizontaldirections is determined by the station distribu-tion; in the vertical direction, they increase withdepth. Each block is subdivided into a finer mesh(see figure 7) to improve the accuracy of the for-ward modelling calculation. The 3-D portion ofthe model is embedded in a 2-D region to thenorth and south, in which the model is assumedto have a north–south strike, and a 1-D region tothe east and west and below the whole modelarea. The inversion was carried out in severalsteps, but allowing all parameters, includingthose of the 1-D portions of the model space, tovary, until a minimum misfit was obtained, using

2.34 A&G • April 2006 • Vol. 47

WHALER: PRESIDENTIAL ADDRESS

1

23

4 5

67

8

910

11

12

14

15

16

17

18

19

20

21

22

23

24N

Y

1-D 1-D

X 2-D

2-D

0 5000 10000 15000 m

12 11 10 9 17 20 22 19

0 5000 10000 15000 distance (m)

0.1 0.2 0.4 0.8 1.6 3 6 12 25 50 100 resistivity (Ωm)

dept

h (m

)

0

–2000

–4000

–6000

X

Z

7: Vertical (a) and horizontal (b) grid layout for 3-D inversion of the BBB data set. The opencircles are the MT sites. The finer grid (dashed lines) is used for more accurate forwardmodelling calculation. The 2-D RRI model is superimposed on the vertical grid. From Hautotet al. (2000).

6: Tectonic setting and data distribution of the BBB area of the Kenyan rift. Solid circles are siteswhere both MT and DC resistivity soundings were undertaken; crosses are sites with just DC resistivitysoundings. The box encloses the area of the 3-D model of figure 8. The thick line is the profile alongwhich 2-D RRI gave the starting model for 3-D inversion. WMTZ, Wasages-Marmanet TransverseZone; KF, Karau Fault; IBF, Ilosowuani-Bechot Fault; KAF, Kapthurin Fault. From Hautot et al. (2000).

(a) (b)

a regularization term to control the resistivitycontrast between blocks.

The final model, presented in depth slices infigure 8, which is a good fit to the data (see fig-ure 10 of Hautot et al. 2000), was subjected to aseries of sensitivity tests to establish that the keystructures were constrained by the data. Thisallowed us to propose a new model of rift devel-opment in the region. The resistive layer 10 offigure 8, between 1.4 and 4.9 km, was initiallythought to be basement, but the conductorbeneath (layer 11) suggests it is a thick layer ofvolcanic rocks underlain by sediments. Thuslayer 11 could represent the initial synrift sedi-mentary infill. Hautot et al. (2000) suggest thatthe volcanic succession is most probably lowerto middle Miocene in age, and hence that theproto-Baringo basin would be Palaeogene tolower Miocene, i.e. contemporary with the deepKerio basin to the west. The structure of layers 8and 9 is different from that above and below.Here the internal basin structure is dominatedby two narrow, north–south oriented conduc-tive structures, labelled Y and Z, and the N140°trending Wasages–Marmanet transverse zone(WMTZ) is only weakly indicated comparedwith its expression in the layers above. How-ever, layer 7 has a similar structure to the nar-row conductor U at the southernmost limit ofthe model. In layers 6 and 7, the resistivity dis-tribution is dominated by a conductive north–south axis, with resistive features linked to vol-canic formations affected by the Kapthurin fault(KAF), Ilosowuani-Bechot fault (IBF) andWMTZ. The north-eastward migration of theWMTZ from its surface position (marked onlayer 5) is consistent with a flexural structure.The resistive features surrounding the conduc-tive area labelled X can be correlated with vol-canic rocks belonging to the Hanningtonformation, associated with the fluvio-lacustrineKapthurin formation (Kfm), and forming theIBF structure. The resistivity contrast along themarked N140° trend following the surface posi-tion of the WMTZ may represent the contactbetween the recent sediments and the Hanning-ton trachyphonolites (HT). Feature X correlateswith the present sedimentary basin, and twomain axes of deposition can be defined: anorth–south one, several kilometres wide in thenorth and thinning southwards, marking thechannels of the main rivers flowing within therift floor and the Lake Bogoria sedimentaryinfill, bordered by alluvial fans, and aneast–west one, corresponding to lateral sedi-mentary fluxes from the western and easternbasin borders. The resistivity distribution in lay-ers 1–4 can be related to the Recent sedimentarycover, comprising silts of ~20 m thickness andalluvial fans which have developed on bothfaulted margins.

The interpretation of this resistivity modelsuggests a marked modification to previous

models of rift propagation (e.g. Mugisha et al.1997), summarized in figure 13 of Hautot et al.(2000). Rifting was believed to have nucleatedon the western side of the rift as early as Palaeo-gene time, with the only expression in the BBBarea being moderate flexure in the late Miocene.During Pliocene to Recent time, rift deformationwas assumed to have shifted eastwards, focus-ing along the Baringo-Bogoria zone that belongsto the active axial trough of the Kenyan rift. Theexistence of a Palaeogene deeply buried basinbeneath the Baringo-Bogoria area (layer 11) sug-gests that the initial rift system was almost twiceas wide as previously thought, consisting of twowesterly tilted half grabens. The resistivity dis-tribution suggests the initial geometry of the

BBB was controlled by both north–south (KAF)and transverse (WMTZ) faults on the successivedepocentres. The initial length of the BBB wasmuch greater, with its current rhomb-shapedstructure accounted for by subsequent flexureand fault segmentation.

Northern main Ethiopian riftIn 2002/3, a number of geophysical techniqueswere used in the Ethiopia Afar GeoscientificLithospheric Experiment (EAGLE) to investi-gate the transition from continental rifting tosea-floor spreading that is taking place in thenorthern main Ethiopian rift. These included thedeployment of broadband and short-period seis-mic receivers, the latter along profiles across and

A&G • April 2006 • Vol. 47 2.35

WHALER: PRESIDENTIAL ADDRESS

HT

(a)

0°30′N

0°25′N

0°20′N

0°30′N

0°25′N

0°20′N

(b)

0°30′N

0°25′N

0°20′N

layer 3 (z = 3–8 m) layer 4 (z = 8–18 m) layer 5 (z = 18–70 m)

layer 6 (z = 70–170 m) layer 7 (z = 170–370 m) layer 8 (z = 370–720 m)

36°00′E 36°05′E 36°00′E 36°05′E 36°00′E 36°05′Elayer 9 (z = 720–1420 m) layer 10 (z = 1420–4920 m) layer 11 (z = 4920–16920 m)

0.1 0.5 1 2 4 8 16 31 62 125 250 500 1000 resistivity (Ωm)

km0 5

KFmX

WMTZ

HT

IBF

KAFY Z

Y Z

U

8: Depth slices through the final 3-D model of the BBB, beginning at 3m below the surface. Opencircles indicate the surface positions of the MT sites. The south-west part of the model has beenmasked because there are no MT sites in the area. See caption to figure 6 for key to most tectonicunits. KFm, Kapthurin formation; HT, Hannington Trachyphonolites. Features U, X, Y and Z aredescribed in the text.

along the rift, and in a 2-D array around theirpoint of intersection, and gravity and MT sur-veys along the cross-rift profile. The seismicdata were used to study seismicity, produce 2-Dand tomographic images, investigate crustalthickness and Poisson’s ratio from receiver func-tion analysis, and analyse seismic anisotropyfrom shear wave splitting. The MT sites werelocated in the same locations as broadband seis-mic receivers where feasible, to enable a com-parison of results. Eighteen MT sites wereoccupied, at an average spacing of 20 km out-side the rift, decreasing to 5 km within it, wherehigher subsurface conductivities restricting pen-etration were expected. The station distribution

is shown in figure 9. Robust response processing(using the methods of Chave and Thomson1989 and Ritter et al. 2003) showed that onewas badly contaminated by cultural noise andwas not considered further, but the others allprovided usable transfer functions, althoughsome only at shorter periods.

Tensor decomposition using the technique ofCounil et al. (1986) showed that the data areconsistent with 2-D interpretation, with a geo-electric strike direction roughly parallel to therift border direction (N30°–40°), though turn-ing slightly northwards at sites within the rift.This rotation matches the orientation of the enechelon magmatic segments within it (see figure

9) and follows the change in orientation of theshear-wave splitting fast direction (e.g. Kendallet al. 2005). Some of the sites suffered from sta-tic shift; at the two most badly affected, thenear-surface resistivity structure was determinedfrom DC resistivity soundings, and the short-period data this structure predicted was used toremove the static shift. The data were modelledby iterative regularized inversion, and the resultis presented in figure 10, taken from Whaler andHautot (2006) where more details of both themodelling and interpretation can be found.

The shallow conductive near-surface structureof figure 10 reflects the near-surface geology andmatches its geometry deduced from surfacemapping (Wolfenden et al. 2004) and imaged bythe cross-rift seismic survey (Mackenzie et al.2005) – Late Miocene-Early Pliocene ign-imbrites and rhyolites, and Quaternary sedi-ments and volcanics within the rift, and a fewkilometres thickness of pre-rift Jurassic sedi-ments and Oligocene flood basalts beneath,which crop out elsewhere along the profile. TheAdama basin appears to have a maximum thick-ness of 3 km in the MT model, slightly less thaninferred from geological mapping and the seis-mic survey. Beneath the Boset volcano in the riftis an extremely conductive (resistivity 0.8 Ωm)shallow lens which could be caused by a shal-low magma zone. There is a good resistivitycontrast between the upper and lower crustbeneath the northern plateau, at depths match-ing those inferred seismically, but materialbeneath the southern plateau is uniformly muchmore resistive. Figure 10 has two conductivebodies at lower crustal depths, one beneath theBoset magmatic segment and the other close tothe rift flanks to the north-west, and slightlydeeper. That beneath the segment coincides witha high seismic velocity (Keranen et al. 2004) anddensity (Cornwell et al. 2006) body, thought tobe caused by 1.8Ma gabbroic mafic intrusion,feeding a series of dykes intruding into the uppercrust. The high conductivity would suggest thatit is at least partially molten, which makes itshigh seismic velocity and density even moreremarkable. Keir et al. (2006) note that seismic-ity within the rift is confined to the magmaticsegment and ceases at a depth of ~10 km. Fromthis they infer that it is caused by dyking into thebrittle upper crust, a mechanism that relies onmelt still being present. There is no suggestionfrom other data of a mafic intrusion associatedwith the second high-conducitvity body to thenorth-west, but the Bishoftu volcanic chain(BVC on figure 10) to the south-west may beresponsible, bearing in mind that the MTmethod is sensitive to structure some distanceaway, particularly to conductive features whichchannel electric current. In fact, there is wide-spread evidence for partial melt beneath thenorth-west plateau, including late teleseismicarrivals from all azimuths (i.e. not just from

2.36 A&G • April 2006 • Vol. 47

WHALER: PRESIDENTIAL ADDRESS

38 39 40 41 7

8

9

10

FURI

AAUS

NAZA

Line 1

Line

2

0 500 1000 1500 2000 2500 3000 3500 4000 4500elevation (m)

9: MT sites (red hexagons) along the cross-rift (line 1) profile of the EAGLE project, superimposed onthe topography, border faults (solid thick black lines), faulted monoclines (dashed thick black lines),Bishoftu Volcanic Chain (linked open circles) and magmatic segments (grey shading) of the northernmain Ethiopian rift. Also shown are the seismic stations, both broadband (black inverted trianglesand white diamonds, EAGLE phase I and II respectively; blue inverted triangles, EAGLE line 1; greysquares, Pennsylvania State University [Dugda et al. 2005]) and short period (black dots, EAGLEalong-rift line and central 2-D array). From Whaler and Hautot (2006).

those arrivals passing through the rift) at thepermanent station FURI (figure 9), shear wavesplitting inferred to arise from oriented meltpockets, and high Poisson’s ratios at a numberof nearby sites, though unfortunately not fromthe sites over the second conductor. The deeperMoho beneath the northern plateau is inferredto be caused by a layer of underplated materialof Oligocene and/or Recent age (Mackenzie etal. 2005). The two conductors lie within thisunderplated layer, and may arise from recentlow-pressure fractionation involving it partiallymelting. Two shallow high-resistivity bodiescoincide with local peaks in the gravity profile,one of which has been interpreted as a furthermafic intrusion (Cornwell et al. 2006). Its topsurface coincides with the top of the resistor, butit has greater depth extent. Possibly the MTmodel only images its solidified top, with thewarmer, deeper part not having a discernibleresistivity contrast with the surroundings.

Joint inversion of the MT and gravity or seis-mic data sets are planned to explore some ofthese possible interpretations of the data.Already, the MT data have provided indepen-dent evidence for extensive partial melt beneathboth the rift and the northern plateau, and a dif-ference in crustal structure between the north-ern and southern plateaux. The rotation ingeoelectric stike direction within the rift sup-ports the magma-assisted rifting hypothesis ofEbinger and Casey (2001) and Buck (2004),whereby the border faults have been abandonedand strain is now taken up within the magmaticsegments in the transition from continental rift-ing to sea-floor spreading.

K A Whaler, Institute of Earth Science, School ofGeoSciences, University of Edinburgh, UK;kathy.whaler@ed.ac.uk.Acknowledgments. Many people have collaboratedwith me or otherwise contributed to the success of theprojects outlined above, through help with the logis-tics, fieldwork, equipment, data processing, model-ling and interpretation. These include, but are nomeans confined to, Laike Asfaw, Atalay Ayele, DavidBailey, Ian Bastow, Johan de Beer, Colin Brown,Dave Cornwell, Graham Dawes, MohammednurDesissa, Cindy Ebinger, Oswald Gwavava, VolkerHaak, Sophie Hautot, Abiy Hunegnaw, Phil Jones,Derek Keir, Mike Kendall, Graeme Mackenzie, PeterMaguire, Oliver Ritter, Graham Stuart, the late Sam-son Takaedza, Pascal Tarits, Jean-Jacques Tiercelin,Ute Weckmann, Teddy Zengeni and the EAGLEWorking Group. In addition, funding has come fromthe NERC through the Geophysical Equipment Facil-ity, a PhD studentship to David Bailey and severalresearch grants, and an EU (Marie Curie) Fellowshipto Sophie Hautot. The British Council, Elf PetroleumNorge AS, the National Oil Corporation of Kenya,the Ministry of Mines and Energy of Ethiopia, andthe universities of Addis Ababa and Zimbabwe havealso provided funding or assistance in kind.

ReferencesBahr K 1988 J. Geophys. 62 119–127.Bahr K 1991 Phys. Earth Planet. Interiors 66 24–38.Bailey D S 1998 Magnetotelluric studies of the Zambezimobile belt of Northern Zimbabwe, PhD thesis,University of Edinburgh.Bailey D et al. 2000a J. Geophys. Res. 105 11185–11202.Bailey D et al. 2000b Geophys. J. Int. 142 898–914.Balasis G et al. 1997 Geophys. J. Int. 129 472–473.Bosum W and Geipel H 1988 Reinterpretation of anaeromagnetic profile in the Zambezi Valley of Zimbabweon the basis of magnetotelluric and gravimetric datatechnical report (Bundesanstalt FürGeowissenschaften und Rohstoffe, Hanover).Brewitt-Taylor C R and Weaver J T 1976 GJRAS 47375–396.Buck W R 2004 in Karner G D et al. (eds) Rheology anddeformation of the lithosphere at continental margins(Columbia Univ. Press, New York) 1–30.Chave A D and Thompson D J 1989 J. Geophys. Res. 9414 202–14215.Constable et al. 1987 Geophysics 52 289–300.Cornwell D G et al. 2006 in Yirgu G et al. (eds) Structureand Evolution of the East African Rift System in the Afarvolcanic province Geol. Soc. London Spec. Pub. 259309–323.Counil J–L et al. 1986 Ann. Geophys. 4 115–130.deGroot-Hedlin C and Constable S 1990 Geophysics 551613–1624.Dugda M T et al. 2005 J. Geophys. Res. 110doi 10.1029/2004JB003065.Ebinger C and Casey M 2001 Geology 29 527–530.Egbert G D and Booker J R 1986 Geophys. J. R. astr. Soc.87 173–194. Gamble T D et al. 1979 Geophysics 44 53–68.Groom R W and Bailey R C 1989 J. Geophys. Res. 941913–1925.Haak V and Hutton R 1986 in J B Dawson et al. (eds)The Nature of the Lower Continental Crust Geol. Soc.London Spec. Pub. 24 35–49.Hautot S et al. 2000 J. Geophys. Res. 105 23493–23518.Hiller K and Buttkus B 1996 Z. angew. Geol. 42

132–137.Keir D et al. 2006 J. Geophys. Res. in press.Kellett R L et al. 1992 Geophys. J. Int. 111 141–150.Kendall J-M et al. 2005 Nature 433 146–148.Keranen K et al. 2004 Geology 32 949–952.Livelybrooks D et al. 1993 Phys. Earth Planet. Interiors81 67–84.Losecke W et al. 1988 Technical Report 84.2171.1(Bundesanstalt für Geowissenschaften und Rohstoffe,Hanover).Mackenzie G D et al. 2005 Geophys. J. Int. 162994–1006. Meju MA 1996 Geophysics 61 56–65.Mugisha F et al. 1997 Tectonophys. 278 61–81.Orpen J L et al. 1989 J. Afr. Earth Sci. 8 215–229.Parker R L 1980 J. Geophys. Res. 85 4421–4428.Parker R L and Booker J R 1996 Phys. Earth Planet.Interiors 98 269–282. Pek J and Verner T 1997 Geophys. J. Int. 128 505–521.Ritter O et al. 1998. Geophys. J. Int. 132 535–548.Ritter O et al. 2003 Phys. Earth. Planet. Int. 138 71–90.Simpson F and Bahr K 2005 Practical magnetotellurics(Cambridge University Press, Cambridge).Smith J T 1995 Geophys. J. Int. 122 219–226.Smith J T and Booker J R 1991 J. Geophys. Res. 963905–3922.Swift C M 1967 in Vozoff K (ed.) MagnetotelluricMethods Society of Exploration Geophysicists,Geophysics Reprints 5 156–166.Telford W M et al. 1990 Applied Geophysics (CambridgeUniversity Press).Vozoff K 1972 Geophysics 37 98–141.Wannamaker P E et al. 1987 Geophys. J. Roy. Astr. Soc.88 277–296.Whaler K A and Hautot S 2006 in Yirgu G et al. (eds) TheStructure and Evolution of the East African Rift Systemin the Afar Volcanic Province Geol. Soc., London, Spec.Pub. 259 295–307.Wolfenden E et al. 2004 Earth Planet. Sci. Letts. 224213–228.Zhang P et al. 1987 Geophysics 52 267–278.

A&G • April 2006 • Vol. 47 2.37

WHALER: PRESIDENTIAL ADDRESS

–40

–35

–30

–25

–20

–15

–10

–5

0

z (k

m –

scal

e x

2)

–120 –100 –80 –60 –40 –20 0 20 40 60 80 100 120distance (km)

18 19 03 02 05 04 06 07 08 09 101211 13 14 15 16 17

NW SE

Adama basin

Boset magmatic segment

0.2 0.4 0.6 0.8 1.0 1.2 1.4 1.6 1.8 2.0 2.2 2.4 2.6 2.8 >3.0log resistivity (Ω.m)

10: 2-D model of the resistivity structure across the main Ethiopian rift, from an inversion of both theTE and TM mode data on a finite difference grid. Station numbers, the extent of the Boset magmaticsegment and Adama basin, and the projection of the Bishoftu Volcanic Chain (BVC) onto the model,are indicated along the top. The Adama basin terminates to the south-east against the Arboye riftborder fault (Wolfenden et al. 2004). Data from site 12 were not included in the inversion becausethey were too badly contaminated by cultural noise. From Whaler and Hautot (2006).