Mantle transition zone structure and upper mantle S velocity variations beneath Ethiopia: Evidence...

Mantle transition zone structure and upper mantle S velocityvariations beneath Ethiopia: Evidence for a broad,deep-seated thermal anomaly

Margaret H. BenoitDepartment of Geosciences, Penn State University, University Park, Pennsylvania 16801, USA

Now at Earth Resources Laboratory, Massachusetts Institute of Technology, E34-556, Cambridge, Massachusetts 02142,USA ([email protected])

Andrew A. NybladeDepartment of Geosciences, Penn State University, University Park, Pennsylvania 16801, USA

Thomas J. OwensDepartment of Geological Sciences, University of South Carolina, Columbia, South Carolina 29201, USA

Graham StuartSchool of Earth and Environment, University of Leeds, Leeds LS2 9JT, UK

[1] Ethiopia has been subjected to widespread Cenozoic volcanism, rifting, and uplift associated with theAfar hot spot. The hot spot tectonism has been attributed to one or more thermal upwellings in the mantle,for example, starting thermal plumes and superplumes. We investigate the origin of the hot spot by imagingthe S wave velocity structure of the upper mantle beneath Ethiopia using travel time tomography and byexamining relief on transition zone discontinuities using receiver function stacks. The tomographic imagesreveal an elongated low-velocity region that is wide (>500 km) and extends deep into the upper mantle(>400 km). The anomaly is aligned with the Afar Depression and Main Ethiopian Rift in the uppermostmantle, but its center shifts westward with depth. The 410 km discontinuity is not well imaged, but the660 km discontinuity is shallower than normal by �20–30 km beneath most of Ethiopia, but it is ata normal depth beneath Djibouti and the northwestern edge of the Ethiopian Plateau. The tomographicresults combined with a shallow 660 km discontinuity indicate that upper mantle temperatures are elevatedby �300 K and that the thermal anomaly is broad (>500 km wide) and extends to depths �660 km. Thedimensions of the thermal anomaly are not consistent with a starting thermal plume but are consistent witha flux of excess heat coming from the lower mantle. Such a broad thermal upwelling could be part of theAfrican Superplume found in the lower mantle beneath southern Africa.

Components: 8272 words, 11 figures.

Keywords: Ethiopia; Afar; mantle plume; superplume.

Index Terms: 7208 Seismology: Mantle (1212, 1213, 8124); 8120 Tectonophysics: Dynamics of lithosphere and mantle:

general (1213); 8121 Tectonophysics: Dynamics: convection currents, and mantle plumes.

G3G3GeochemistryGeophysics

Geosystems

Published by AGU and the Geochemical Society

AN ELECTRONIC JOURNAL OF THE EARTH SCIENCES

GeochemistryGeophysics

Geosystems

Article

Volume 7, Number 11

16 November 2006

Q11013, doi:10.1029/2006GC001398

ISSN: 1525-2027

Copyright 2006 by the American Geophysical Union 1 of 16

Received 26 June 2006; Accepted 18 July 2006; Published 16 November 2006.

Benoit, M. H., A. A. Nyblade, T. J. Owens, and G. Stuart (2006), Mantle transition zone structure and upper mantle S velocity

variations beneath Ethiopia: Evidence for a broad, deep-seated thermal anomaly, Geochem. Geophys. Geosyst., 7, Q11013,

doi:10.1029/2006GC001398.

1. Introduction

[2] For much of the Cenozoic, Ethiopia has beensubjected to hot spot tectonism (magmatism, rifting,and plateau uplift). Basaltic volcanism began inEthiopia �45–40 Ma [Davidson and Rex, 1980;Zanettin et al., 1980; WoldeGabriel et al., 1990;Kieffer et al., 2004; Ebinger et al., 1993; Georgeet al., 1998], and volcanic activity culminated in theAfar and central portion of the Ethiopia Plateauaround 30 Ma, when up to 2000 m of flood basaltswere erupted within 1–2 Myr [Mohr, 1983; Kiefferet al., 2004; Baker et al., 1996; Ayalew et al., 2002;Coulie et al., 2003; Wolfenden et al., 2004]. Addi-tional shield volcanism occurred across Ethiopia be-tween 30–10 Ma that added 1000–2000 m of localrelief [Kieffer et al., 2004; Coulie et al., 2003], andmagmatic processes in the Afar Depression andMainEthiopian Rift (MER) continue to the present.

[3] Rifting in Ethiopia is associated with the for-mation of the triple junction in Afar, where the RedSea rift, Gulf of Aden rift and the MER join.Crustal extension in the present-day location ofthe Red Sea and Gulf of Aden initiated during theOligocene when Africa and Arabia began to sep-arate. The development of the MER began later atapproximately 18 Ma in southwestern Ethiopia and11 Ma in the central and northern regions ofEthiopia [Wolfenden et al., 2004; Chernet andHart, 1999; Ebinger et al., 2000; WoldeGabrielet al., 1997]. Plateau uplift across Ethiopia beganbetween 30 and 20 Ma [Pik et al., 2003], withelevations over a large portion of Ethiopia todayexceeding 1500 m (Figure 1).

[4] The origin of the hot spot tectonism in Ethiopiaremains enigmatic. Many authors have attributedthe tectonism to one or more mantle plumes [e.g.,Schilling et al., 1992; Ebinger and Hayward, 1996;Burke, 1996; Marty et al., 1996; Ebinger andSleep, 1998; George et al., 1998; Courtillot etal., 1999; Debayle et al., 2001; Montelli et al.,2004; Pik et al., 2006], often invoking the startingthermal plume of Griffiths and Campbell [1990],which has a large head followed by a narrow tail. A

number of these studies suggest the presence of aplume head under Ethiopia today [e.g., Schilling etal., 1992; Burke, 1996; Debayle et al., 2001;Montelli et al., 2004], while others [e.g., Ebingerand Sleep, 1998] suggest that the plume headmaterial may no longer exist in any great amountwithin the upper mantle under Ethiopia.

[5] It has also been proposed that the hot spottectonism in Ethiopia could be the surface manifes-tation of the so-called African Superplume [Ritsemaet al., 1999; Benoit et al., 2006]. The AfricanSuperplume is a low-velocity anomaly that extendsupward from the core-mantle boundary beneathsouthern Africa in most global tomography models[e.g., Masters et al., 1996; Su et al., 1994; Li andRomanowicz, 1996], and in some models the anom-aly can be traced into the upper mantle beneath east-ern Africa [e.g., Ritsema et al., 1999; Grand et al.,1997; Grand, 2002; Trampert et al., 2004]. Inthese models, the superplume anomaly has a width>500 km beneath Ethiopia within the transition zone,consistent with the regional tomographic P wavemodel of Benoit et al. [2006].

[6] In contrast to the regional [Benoit et al., 2006]and global [Ritsema et al., 1999; Grand et al.,1997; Grand, 2002; Trampert et al., 2004] tomo-graphic models suggesting a broad, deep seatedmantle thermal anomaly beneath Ethiopia, resultsfrom receiver function imaging of the 410 and660 km discontinuities show little evidence for thin-ning of the transition zone beneath the Afar or centralEthiopia [Nyblade et al., 2000; Chevrot et al., 1999],indicating that a deep seated thermal anomaly, aswould be expected for a superplume, might notexist. These discontinuities are generally consid-ered to be mineral phase transformations of olivine(i.e., a-spinel to b-spinel at 410 km, g-spinel toperovskite and magenesiowustite at 660 km). TheClapeyron slopes of the equilibrium phase bound-aries indicate that in regions where temperaturesare warmer than normal, the 410 km discontinuitywill be depressed (i.e., occur at a depth deeper than410 km), and the 660 km discontinuity will beelevated [Bina and Helffrich, 1994].

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[7] To evaluate starting plume and superplumemodels for the hot spot tectonism in Ethiopia, inthis paper we examine further the nature of the 410and 660 km discontinuities beneath Ethiopia andmodel the S wave velocity structure of the uppermantle using broadband data from the EthiopiaBroadband Seismic Experiment [Nyblade andLangston, 2002] and the Ethiopia Afar Geoscien-tific Lithosphere Experiment (EAGLE) [Maguireet al., 2003; Bastow et al., 2005]. By imaging theupper mantle S wave velocity structure and reliefon the 410 and 660 km discontinuities, we placenew constraints on the width and depth extent ofthe upper mantle thermal anomaly beneath Ethio-

pia inferred from regional and global tomographicimages, and we in turn use the new constraints toaddress models for the origin of the hot spottectonism.

2. S Travel Time Tomography

[8] The data used for the S wave travel timetomography came from the Ethiopia BroadbandSeismic Experiment (Figure 1) that operated be-tween 2000 and 2002. The experiment consisted of22 stations located on the western and eastern sideof the Ethiopian Plateau, and 5 stations located inthe Main Ethiopian rift and the Afar Depression.

Figure 1. Map of Ethiopia showing surface elevation and the outline of the Main Ethiopian rift (MER) and the Afardepression (bold dashed line). Ethiopia Broadband Seismic Experiment station locations are denoted by black squares(stations operating between 2000 and 2002) and black triangles (stations operating between 2001 and 2002). Thelocations of permanent broadband stations ATD and FURI are shown as white squares.


benoit et al.: ethiopian mantle transition zone structure 10.1029/2006GC001398benoit et al.: ethiopian mantle transition zone structure 10.1029/2006GC001398

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Further details about the Ethiopia Broadband Seis-mic Experiment are given by Nyblade and Langston[2002]. Additionally, we used data from ATD andFURI, permanent broadband stations located inDjibouti and Ethiopia, respectively.

[9] For the travel time tomography, we analyzed Swaves from earthquakes at distances of 30� to 84�and SKS waves for events between 87� and 140�,yielding a total of 915 S rays and 1012 SKS rays(Figure 2). The S wave data were band-passedfiltered between 0.04–0.1 Hz. Most of the eventsanalyzed had magnitudes larger than 5.5 and mainlyoccurred to the east in the Indian and westernPacific subduction zones. To expand the azimuthaldistribution, we selected several well-recordedevents to the west, north, and south of the Ethiopianetwork with magnitudes as low as 5.0 with clear Sarrivals.

[10] Relative S wave travel time delays were cal-culated with respect to the IASP91 global referencemodel [Kennett and Engdahl, 1991] using themultichannel cross correlation method (MCCC)of VanDecar and Crosson [1990]. The MCCCmethod is composed of three steps. First, a peakor trough of a given phase that is coherent on all ofthe waveforms is selected. Next, a cross correlationwindow around the phase (15 s for S and SKS) isselected, and the maximum of the waveform crosscorrelation function for all possible pairs of stationsi and j (tijmax) is computed. Combining the lagtime from the cross correlation with arrival timeestimates from step 1 (ti and tj) results in the crosscorrelation-derived relative delay time betweenstations i and j, Dtij = ti � tj � tijmax.

[11] Because of noise, relative delay times are notnecessarily consistent (i.e., Dt13 6¼ Dt12 + Dt23),and to account for this, the cross correlationderived delay times are used to generate a systemof linear equations solved by a least squaresoptimization scheme that adjusts the delay timesto minimize the inconsistencies so that Sti = 0.

[12] The average variation in the delay times com-puted from the MCCC was about 5 s for S and SKSwaves, however, the delay times varied by as muchas 8 seconds. Generally, the largest variation indelay times is found east-west across the network.Variations in the S wave relative travel times arecomparable to those found in a similar study fromEthiopia [Bastow et al., 2005] and for Tanzania(i.e., 7–8 s) where Cenozoic rifting is also occur-ring [Ritsema et al., 1998]. The uncertainties in thedelay times are 0.1–0.3 seconds for both S and

SKS. These estimates are determined using thestandard error analysis in the MCCC and byobserving the variations in the delay times due tochanges in the MCCC parameters.

2.1. Travel Time Inversion

[13] After the relative delay times were computedwe inverted them for a 3-D velocity model usingthe method of VanDecar [1991]. We used anintermediate spline tension of 10 to maintain asmooth model that still characterized local struc-ture. The model parameterization consisted of agrid of 34 knots in depth, 51 knots in latitude and48 knots in longitude for a total of 83,232 knots. Inthe upper 100 km of the model the knots werespaced 25 km apart in depth, between 100–700 kmdepth the knots were spaced 50 km apart, and from700 to 1000 km depth, the knots were 100 kmapart. The horizontal knot spacing was 1/3� in theinterior of the model and increased to 1.5� in theoutermost portion of the model. Station terms wereincluded in the inversion to absorb near stationvelocity variations, and event locations to absorberrors from event mislocations and anomalousstructure outside of the model domain [Bostockand VanDecar, 1995]. The starting model for theinversion was the IASP91 reference earth model[Kennett and Engdahl, 1991].

[14] To regularize the slowness model, the codeuses constraint equations that minimize spatialgradients (the first spatial derivative) and modelroughness (the second spatial derivative) [Bostockand VanDecar, 1995]. The final models werechosen on the basis of a trade-off between modelsmoothness and the fit of the data, resulting in an Smodel that accounts for 87% of the RMS misfitreduction. About 1% of the RMS misfit reductionwas absorbed by the station terms, and 2% of theresidual reduction was absorbed by the eventrelocation terms in the inversion. Station termsare shown in Figure 3b.

2.2. S Velocity Models

[15] Shown in Figures 3 and 4 are the S wavevelocity model obtained from the travel time in-version, along with the P wave models from Benoitet al. [2006] for comparison. Results are shownfrom 150 km to 400 km, because the model lackssufficient crossing ray coverage to clearly resolvestructure shallower than 150 km and deeper than400 km.



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[16] The first order features in the S and P wavemodels are similar. Both models show strongvelocity variations between 150 and 400 km depth.At 150 km depth, the S and P models show a low-velocity region beneath the Afar Depression, MainEthiopian rift, and slightly west of the rift under thewestern portion of the Ethiopian Plateau. Deeper inthe models (200–400 km depth), the center of thelow-velocity structure shifts westward across theWestern Ethiopian Plateau, offset from the strike ofthe Main Ethiopian rift. Additionally, the S and Pmodels both reveal faster than average velocitiesbeneath the Eastern Ethiopian Plateau and thenorthwest section of Western Ethiopian Plateau.

[17] While the first order features of the S and Pmodels are similar, there are some second orderdifferences. The S wave model has a maximumvelocity variation of ±4% compared to the ±2.5%variation in the P wave model. This velocityvariation primarily reflects differences of delaytimes for the S and P waves. S wave delay times,

on average, varied by �4 s, while the P wave delaytimes varied by �1.5 s.

[18] The second difference between the S and Pmodels is that the S wave model shows lower wavespeeds under the Afar depression in the 150–200 kmdepth interval than the P wave model. We attributethis difference to greater vertical (upward) smear-ing in the S wave model compared to the P wavemodel. The smearing in the S wave model is due,in large part, to the inclusion of SKS waves in ouranalysis, because SKS waves have very steeplydipping travel paths that smear velocities in a nearvertical direction more than travel paths for direct Swaves.

[19] The third difference between the models is thatthe S wave model has a more ‘‘patchy’’ nature thanthe P wave model. This is also likely from theinclusion of SKS waves in the S model, and thefact that there were fewer S waves than P wavesused in the travel time inversion.

Figure 2. Event locations for S wave travel times. Black dots represent event locations for S waves, and red dotsshow the event locations for SKS waves. The star denotes the location of the Ethiopia Broadband SeismicExperiment. Concentric circles denote great circle angular distances in degrees from the center of the EthiopiaBroadband Seismic Experiment network.



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Figure

3.

Horizontalcross

sectionssliced

through(a–d)Sand(e–h)Pwavevelocity

modelsfrom

travel

time

tomography.

Horizontalcross

sectionsthroughthemodel

areshownfordepthsof150,200,300,and400km.The

locationsoftheverticalcross

sectionsin

Figure

4areshownin

Figure

3a.Politicalboundariesareshownin

white,and

approxim

aterift

boundariesareoutlined

ingray.

The2,000m

elevationcontourlineis

shownin

blue.

Station

locationsaredepictedas

smallwhitesquares,andregionsofthemodel

withlow

raycoveragevolumeareshaded

inblack

accordingto

hitcount,withahitcountof10havinglightshadingandahitcountof0havingcompleteshading.

ThePwavevelocity

model

isafterBenoitet

al.[2006].



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[20] The shallow portion of our model (<200 kmdepth) is similar to the results of Bastow et al.[2005] using data from the EAGLE experiment.Both models show low velocities with similaramplitudes beneath the MER, although our modelalso shows low velocities west of the rift. Similarly,the deeper portion of our model (>200 km depth)shows that the low-velocity zone extends furtherwestward than the low-velocity zone in the studyby Bastow et al. [2005]. This westward extensionof the low-velocity zone is likely due to the factthat the Ethiopia Broadband Seismic Experimenthad greater station coverage west of the rift thanthe EAGLE experiment.

2.3. Model Resolution

[21] Model resolution is influenced by severalfactors, including limited ray coverage and thesteep incidence angles of teleseismic S and SKSwaves. In addition, the inversion technique of

VanDecar [1991] inherently penalizes steep spatialgradients, thereby making sharp velocity transi-tions difficult to resolve. To evaluate model reso-lution, we performed a number of tests. Synthetictravel time data were generated for many inputmodels using the same ray paths of our actualmodel and then inverted using the same modelparameterization and regularization parameters thatwe used to invert our data. Results for an inputcheckerboard model of alternating spherical Gauss-ian pulses are shown in Figure 5. This figureillustrates that the horizontal resolution in ourmodel is at least 100 km (equal to 3 nodes in ourmodel parameterization), and that lateral resolutionappears to be best between 200 and 400 km depth.

[22] The vertical resolution in the model is morelimited than the horizontal resolution. Figure 6shows the vertical cross sections through the inputmodel where the spheres are centered at 200 km indepth. Figures 6b and 6c show vertical cross

Figure 4. Vertical cross sections sliced through (a–c) S and (d–f) P wave models. Cross section locations areshown in Figure 3a. Surface features are shown above each cross section, and the shading used is the same as inFigure 3.



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sections of the recovered models, illustrating thatabout 100 km of vertical smearing occurs bothdownward and upward.

[23] Model resolution can also be influenced byanisotropy. SKS wave splitting measurements fromthe Ethiopia Broadband Seismic Experiment[Gashawbeza et al., 2004] and the EAGLE net-work [Kendall et al., 2005] show splitting timesbetween 0.5–1.7 s with most of the fast polariza-tion directions slightly oblique to the rift directionand almost parallel to the strike of the SW-NE trend-ing velocity anomaly shown in Figures 3 and 4.Gashawbeza et al. [2004] and Kendall et al. [2005]conclude that the splitting is most likely due toalignment of melt in the rifted lithosphere and theorientation of structural trends in the Precambrianorogenic basement. The average travel time variationfor S waves in our data set is �4s, and a splittingtime of �1 s could account for roughly 25%–30%of this travel time variation. Therefore the S velocity

anomaly amplitude could be somewhat altered byanisotropy. However, all the shear wave splittingfast directions are roughly parallel to one another;thus we do not expect anisotropy to change theoverall shape of the velocity anomaly. To examinethe effect of including SKS waves in our S model,we inverted the S wave measurements alone with-out the SKS measurements and recovered a veloc-ity model with a similar velocity pattern to themodelwhere both sets of measurements (S and SKS) wereused.

3. Receiver Function Analysis

[24] The low-velocity anomaly imaged in both theP and S body wave models suggests that a thermalanomaly may be present throughout the uppermantle beneath much of Ethiopia. The verticalsmearing present in the tomographic models, how-ever, makes it difficult to assess whether or not thedepth extent of the anomaly is restricted to the

Figure 5. Examples of horizontal resolution using a checkerboard model. (a) Input models consisted of 9 sphericallyshaped Gaussian pulses with a peak velocity of ±10% centered at 150, 200, 300, and 400 km. Cross sections A and Bare shown in Figure 6. (b–e) Output from the resolution test.



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region above the mantle transition zone. To helpconstrain the depth extent of the thermal anomaly,as stated in the introduction, we next apply ageographic stacking method to receiver functionsto investigate relief on the 410 and 660 kmdiscontinuities (hereafter referred to as the 410and 660, respectively).

3.1. Data and Stacking Method

[25] For this part of the study, we combined datafrom earthquakes in the 30 to 90 degree distancerange recorded by the Ethiopia Broadband Seismic

Experiment and the EAGLE network [Maguire etal., 2003; Bastow et al., 2005]. Figure 7 shows amap of stations and pierce points at 410 and 660 kmdepths from the 1275 station-event pairs used. Weapplied a Gaussian filter with a width of 1.0 tofilter the data before deconvolution, and thencomputed receiver functions using a frequencydomain deconvolution method with water levelstabilization [Langston, 1979]. The iterative decon-volution method of Ligorria and Ammon [1999]was also tested but did not yield receiver functionswith better signal-to-noise ratio than the water leveldeconvolution.

[26] We used the procedure from Owens et al.[2000] to stack the receiver functions. For eachreceiver function a travel time, pierce point, andamplitude were calculated at 10 km intervals from40 to 800 km depth for a specified mantle velocitymodel. By doing this, a specific depth in thesubsurface can be correlated with an amplitudeon a receiver function. The amplitudes were binnedas a function of pierce points, using a variable-sizecircular bin at each depth interval, and summed toproduce a receiver function stack.

[27] To determine the optimal bin size for stacking,user-defined stacking criteria must be met or elsethe bin is incremented in size until the criteria aremet. We found that including traces from at least4 stations together with a minimum of 30 receiverfunctions per bin produced stacks with discernablePs arrivals.

[28] Figure 8 shows the region over which thereceiver functions were stacked. Bin sizes were0.5�–0.75� in the center portion of the study areaand increased to 1.5� at the edges of the area. Westacked the receiver functions using both a 1-Dmantle model (i.e., IASP91 [Kennett and Engdahl,1991]) and a 3-D mantle model from the S and Pteleseismic tomography, combined with the crustalmodel for Ethiopia from Dugda et al. [2005].

3.2. Results

[29] The results of the receiver function stackingindicate that the transition zone discontinuity struc-ture beneath Ethiopia is complicated. P to S con-versions that could be interpreted as the 410 arevisible only on a limited number of stacks, mainlyunder the southern part of the Afar and eastern sideof the Ethiopian Plateau. This result is illustrated inFigure 9, where the P to S conversion can be seenas a weak arrival on the cross section at 41�longitude at �430 km depth but is not clear on

Figure 6. Examples of vertical resolution. (a) Inputmodel containing spherical Gaussian pulses with peakvelocities of ±10% at 200 km depth. (b and c) Output ofresolution tests. The locations of the cross sections areshown in Figure 5.



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much of the cross section at 11.5� latitude. Becausethe 410 is not well imaged, we do not attempt tointerpret it and instead focus on the 660.

[30] The 660 can be imaged beneath much of thestudy region and exhibits considerable relief. Forexample, in Figure 9 the 660 is visible on bothcross sections at depths between 625 and 675 km.The depth of the 660 beneath the network is shownin Figure 10. The 660 is at normal depths of�660–670 km in the northwestern portion of theWestern Plateau and beneath Djibouti. Beneath theEastern Plateau and central and southwestern partsof the Western Plateau the 660 is between 630–640 km depth, and is shallowest beneath the south-ern part of the MER, where it is found between625 and 635 km depth. The 660 cannot be clearlyimaged beneath much of the central and northernparts of the MER.

[31] To further illustrate the regional differences inthe depths to the 660, we constructed large-radiusstacks in different subregions of the study area.Figure 11 shows stacks for bins with a radiusbetween 1�–2�, and the subregions for the stacksare shown in Figure 10. The 660 is visible on all ofthe stacks and varies considerably in depth. In theNW Plateau area (subregion 1), it is at 670 kmdepth, and beneath Djibouti (subregion 2) it is at

Figure 7. Maps of pierce point locations for depths of (a) 410 and (b) 660 km. Black dots represent the piercepoints. Inset maps show station locations: red triangles, station locations for the Ethiopia Broadband SeismicExperiment; blue squares, locations of the EAGLE broadband stations.

Figure 8. Map of bin sizes used for the receiverfunction stacks at 660 km depth. Green area representsregion with bin sizes �1�. Blue area denotes the regionwith bin size of 0.75�, and pink area shows the regionwith bin size of 0.5�. Dashed lines show the outline ofthe MER and Afar Depression.



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660 km depth. The ‘‘normal’’ depth of the 660 insubregion 1 corresponds to a region with fasterthan average velocities in the tomographic models(Figure 3). For all other subregions (3–5), the 660is shallower than normal and appears to be shal-lowest (�625 km depth) beneath the southernportion of the MER (subregion 5).

[32] Uncertainties in the depth estimates of the 660can arise from a number of factors. To assess theuncertainty from possible biases in the data, westacked the two different data sets (Ethiopia Broad-band Seismic Experiment and EAGLE) separately.We found no perceptible differences in structure ofthe 660 for the different data sets. Another sourceof uncertainty could be due to anisotropy in theupper mantle. To investigate this, we stackedtransverse component receiver functions, and thestacks exhibited significantly smaller amplitudeswith few coherent arrivals. Other uncertaintiescould arise from upper mantle velocity variationsnot well resolved in the P and S wave bodywave tomography models, noise in the data,and the depth increment used in the stackingprocedure. Considering these sources, we estimatethe uncertainty in the depth of the 660 shown inFigures 9–11 to be ±10 km.

[33] P to S conversions at a number of other depthscan be seen in the receiver function stacks. The520 km discontinuity is seen in many of the stacks,and appears to vary in depth by ±10 km with long

Figure 9. Receiver function stacks for cross sections at (a) 11.5� latitude, (b) 41� longitude. The blue and reddashed lines show the depths of the 410 and 660 km discontinuities, respectively. The green dashed line representsthe 520 km discontinuity. Locations of cross sections are shown in Figure 10.

Figure 10. Maps showing depth to the 660 kmdiscontinuity determined from stacking receiver func-tions using a 3-D velocity model. The circles representlocations of subregional stacks shown in Figure 11.Dashed lines show the outline of the MER and AfarDepression. The dotted lines show locations of the crosssections shown in Figure 9.



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wavelength (>100 km) relief (Figure 9). The originof the 520 km discontinuity is controversial; it mayresult from an olivine phase transformation from a-spinel to b-spinel [Rigden et al., 1991], or alterna-tively, the bottom of a garnet layer [Gasparik,1993]. A strong positive arrival is also visible ataround 480 km depth in some places, however theorigin of this phase is uncertain (Figure 9). Addi-tionally, we find an arrival at depths between 700and 720 km in several stacks (Figure 9). Previousstudies that have found a discontinuity at this depthhave been near subduction zones and the discon-tinuity has generally been attributed to a phasechange in garnet [Niu and Kawakatsu, 1996;Simmons and Gurrola, 2000]. Examining moveoutcurves of P to S converted phases from the receiverfunctions does not convincingly identify any of thephases at 480, 520, or 710 km depths as multiplesfrom shallower discontinuities.

4. Discussion

[34] In this section we examine the geodynamicimplications of our findings for understanding the

origin of the hot spot tectonism in Ethiopia. Theresults of our S wave tomography and receiverfunction analysis are consistent with the P wavetomographic results from Benoit et al. [2006]. TheS wave tomographic model shows an elongated,deep-seated low-velocity structure that broadensacross the Western Plateau with depth. Resolutiontests indicate that this broad structure is not causedby a narrow feature smearing laterally over dis-tances � 500 km, such as a narrow (�100–200 kmdiameter) plume tail. The vertical resolution in theS model is not sufficient to constrain the depthextent of the low-velocity anomaly, althoughresults from the P wave study from Benoit et al.[2006] suggest that the anomaly extends to at least400 km depth. The P wave model favors a broadthermal upwelling related to the African Super-plume, and the S wave model is consistent withthat interpretation.

[35] The results of our receiver function stackscorroborate the interpretation of the tomographicmodels for a broad thermal upwelling most likelylinked to the African Superplume. The 660 is

Figure 11. Receiver function stacks for the subregions shown with circles in Figure 10. The dashed blue and redlines show depths to the 410 and 660 km discontinuities, respectively. The green dashed line denotes the 520 kmdiscontinuity.



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elevated everywhere in our study area (�630 kmdepth), except under the northwestern EthiopianPlateau and Djibouti, indicating that the thermalanomaly in the upper mantle likely extends as deepas the base of the transition zone over a wideregion. We cannot trace the thermal anomalythrough the transition zone into the lower mantle,but the fact that a thermal anomaly exists at thebase of the transition zone suggests a thermalconnection with lower mantle structure.

[36] It is notable that we are unable to resolve the410. This is consistent with the observations ofNyblade et al. [2000], who were not able to imagea 410 beneath central Ethiopia. Additionally, weare not able to resolve the 660 under much of theMER. Our inability to resolve the 410 and 660 inthese places could be due to many factors. Com-plicated crustal and upper-most mantle structure[Dugda et al., 2005], such as sharp lateral changesin Moho depth across the rift, and large variationsin surface topography [Rondenay et al., 2005],could cause wavefield scattering. The presence ofmelt beneath the rift [Bastow et al., 2005; Kendallet al., 2005] could attenuate the Ps convertedphases, and topography on the discontinuities withamplitudes on the order of 15–25 km (in depth)could cause inherent waveform distortion andfocused and defocused amplitudes on the Ps con-verted phases [Van der Lee et al., 1994].

[37] Our observations of the 660 also differ insome respects to those of Nyblade et al. [2000].While we are unable to image the 660 beneathmost of the MER, Nyblade et al. [2000] were ableto image the 660 using data from station AAE, onthe northwest flank of the MER. The pierce pointsfor the 660 Ps conversions stacked by Nyblade etal. [2000] were scattered over a large area(�500 km) that included regions to the west ofthe rift where we are able to image the 660. It maybe that the receiver function stack from Nyblade etal. [2000] is more representative of the WesternPlateau, where we image the 660, than the MER,where we do not image the 660. Nyblade et al.[2000] were also able to image the 410 beneath thestation ATD in Djibouti and reported that the depthof the 410 was not anomalous. The discrepancy inresults for the Djibouti may reflect the differentdata sets used, as Nyblade et al. included data withpierce points to the east and north of Djibouti thatwere not used in this study.

[38] A shallower-than-usual 660 beneath much ofthe study area indicates that the thermal anomalyimaged at �400 km depth in our P and S tomog-

raphy likely extends across the transition zone. Atemperature increase of �280–400 K would beneeded to deflect the 660 upward by 20 or 30 km(Figures 9–11), assuming a Clapyron slope of�2.1 MPa/K for the phase transformation [Binaand Helffrich,1994]. This temperature increase isconsistent with the temperature perturbation neededto produce a ±4% variation in S wave speeds inour tomographic model. Using the method of Fauland Jackson [2005], our shear wave velocityvariation corresponds to a temperature increase of�283 K, assuming an average mantle grain size of10 mm. This temperature variation will be some-what larger if the full amplitude of the velocityanomaly is not entirely recovered in our S-wavemodel.

[39] Our findings of a wide (�500 km) low-velocity anomaly beneath Ethiopia that extends to�400 km depth and an elevated 660 km disconti-nuity are difficult to explain with a starting plumemodel for the hot spot tectonism. Althoughmany authors have attributed the hot spot tecto-nism in Ethiopia to one or more mantle plumes[e.g., Schilling et al., 1992; Marty et al., 1996;Burke, 1996; Ebinger and Sleep, 1998; Courtillotet al., 1999; Debayle et al., 2001], the horizontalresolution of our tomographic models is sufficientto rule out the most diagnostic feature of a plumemodel, a narrow (�100 km diameter) plume tail, asthe source of the wide (�500 km) thermal anomalybeneath Ethiopia. Additionally, the depth extent ofthe thermal anomaly beneath Ethiopia, constrainedto be at least 660 km deep by the receiver functionstacks, is too deep to be caused by another diag-nostic feature of a plume model, the plume head.

[40] Instead, our findings are consistent with abroad thermal upwelling model for the hot spot,such as the African Superplume. As noted in theintroduction, the African Superplume is a low-velocity feature that extends from the core-mantleboundary beneath southern Africa to the upper man-tle beneath eastern Africa in some global tomo-graphic models, and has a width >500 km beneathEthiopia within the transition zone [Ritsema et al.,1999; Grand et al., 1997; Grand, 2002; Trampertet al., 2004]. Such a broad anomaly would causethe 660 to be regionally elevated and create low ve-locities across a broad area of the upper mantle, asfound in this study.

[41] The global tomographic results of Montelli etal. [2004] suggest a vertical, diapiric shaped mantleupwelling associated with the Afar hot spot. In thismodel, the center of the mantle upwelling is



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located beneath the Afar and extends from�1450 km depth in the lower mantle to uppermantle depths. This diapiric upwelling model is notconsistent with our tomographic and receiver func-tion results because the thermal anomaly indicatedby our models is much broader, extending severalhundred kilometers to the south and west of theedge of the upwelling in the model of Montelli etal. [2004].

[42] Finally, the degree of material transfer fromthe lower to upper mantle beneath Ethiopia is notwell understood or constrained by this study. Ourresults are consistent with an excess flux of heatfrom the lower mantle into the transition zone,causing the 660 to be elevated and P and S wavevelocities to be reduced in the upper mantle be-neath Ethiopia, but our results to do necessarilyimply or require that there be a flux of materialfrom the lower mantle into the upper mantle. In thisregard, our results are consistent with the sugges-tion by Marty et al. [1996] that there may be littletransfer of material across the mantle transitionzone beneath Ethiopia, and that only heat is trans-ferred from the lower mantle to the upper mantle.

[43] From the occurrence of high 3He at many sitesin Ethiopia, Pik et al. [2006] have proposed amultiplume model whereby there existed a largeplume originating in the lower mantle that createdthe flood basalts c. 30 Ma in Ethiopia, but smallerthermal upwellings originating from within theupper mantle that created most of the other volca-nic provinces in eastern Africa. The model pro-posed by Pik et al. [2006] is consistent with ourfindings in that a flux of material from the lowermantle into the upper mantle is permitted but notrequired to explain our seismic images. We onlyneed excess heat coming from the lower mantle toaccount for our seismic images.

5. Summary

[44] A new regional S wave tomography imagereveals a broad low-velocity anomaly beneathEthiopia that shifts to the west with depth andextends to at least 400 km depth. Images of the660 km discontinuity from 3-D stacking of receiverfunctions suggest that the 660 km discontinuity iselevated by �30 km under much of Ethiopia. Tem-perature variations in the mantle, inferred from thetomographic models and the relief on the 660 kmdiscontinuity, are �300–400 K. The 410 km dis-continuity is not well imaged anywhere, and insome places the 600 km discontinuity is also not

well imaged, possibly for several reasons, includ-ing scattering, attenuation and defocusing due tocomplicated crustal and upper mantle structure.The results of the body wave tomography andimages of the 660 km discontinuity from receiverfunction stacks suggest that the origin of the hotspot tectonism in Ethiopia is associated with abroad (�500 km) and deep (>660 km) thermalupwelling in the mantle. The width and depthextent of the thermal anomaly is not consistentwith a starting thermal plume head or tail presentlybeneath Ethiopia, but is consistent with a flux ofexcess heat coming from the lower mantle modi-fying upper mantle structure beneath Ethiopia overa wide region. Such a broad thermal upwellingcould be part of the African Superplume found inthe lower mantle beneath southern Africa, suggest-ing that the origin of the hot spot tectonism inEthiopia may be linked to a superplume rooted inthe lower mantle.

Acknowledgments

[45] We thank Laike-Mariam Asfaw, Atalay Ayele, and the

technical staff of the Geophysical Observatory of Addis Ababa

University for their help with the Ethiopian Broadband Seis-

mic Experiment and the EAGLE team for providing us with

access to their data. We also thank Cindy Ebinger and Ken

Dueker for their constructive reviews. This research has been

funded by the National Science Foundation (grants EAR

993093 and 0003424).

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Mantle transition zone structure and upper mantle S velocity variations beneath Ethiopia: Evidence...

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