Lithospheric modification during crustal extension in the Main

Ethiopian Rift

Tyrone Rooney,1,2 Tanya Furman,1 Ian Bastow,3,4 Dereje Ayalew,5,6 and Gezahegn Yirgu5

Received 22 December 2006; revised 11 May 2007; accepted 27 June 2007; published 5 October 2007.

[1] Quaternary lavas erupted in zones of tectonomagmatic extension within the MainEthiopian Rift (MER) preserve details of lithospheric structure in the East African RiftSystem. Despite observed source heterogeneity, basalts, trachybasalts, and basaltictrachyandesites erupted in the Wonjii Fault Belt (WFB) and the Silti-Debre Zeyit FaultZone (SDFZ) form coherent fractionation paths dominated by variable removal ofobserved phenocryst phases. Crustal assimilation is not widespread, though it is observedat the southern end of the WFB where both fault belts merge; farther north, assimilation ofcumulate phases related to fractional crystallization of previous magmas is identified.Shallow fractionation conditions (!1 kbar) within the WFB do not change from north tosouth. In contrast, lavas erupted within the contemporaneous SDFZ fractionate at variouscrustal depths. These results indicate a better developed magmatic system beneath theWFB where magmas rose quickly before undergoing more significant fractionation at nearsurface levels and a less developed system beneath the SDFZ. The distribution ofmagmatism and extant geophysical data indicate thinned crust and a single rift-centeredzone of magmatic activity northeast of 8!300N, consistent with a transitional lithospherebetween continental and oceanic settings. Southwest of 8!300N, thicker crust and rift-marginal axes of extension suggest lithosphere with continental affinities. The WFB ispropagating southward in response to extension within the Red Sea Rift; the northwardpropagating SDFZ is related to rifting within the East African Rift System. This regionrecords the unification of two rift systems, requiring care in interpreting the MER assimply transitional between continental and oceanic environments.

Citation: Rooney, T., T. Furman, I. Bastow, D. Ayalew, and G. Yirgu (2007), Lithospheric modification during crustal extension inthe Main Ethiopian Rift, J. Geophys. Res., 112, B10201, doi:10.1029/2006JB004916.

1. Introduction

[2] Lithospheric modification during continental rifting isan axiomatic consequence of plate tectonic processes. Theprocesses whereby the continental crust is modified bymagmatism during the progressive evolution from conti-nental rifting to seafloor spreading, are however poorlyconstrained. These ambiguities generate substantial uncer-tainties in detailing thermal structure, rift-related volcanichazards, and hydrothermal resources. Rift generated litho-spheric heterogeneity also frustrates efforts to construct

coherent geodynamic and geophysical models of continen-tal rifting in zones of active tectonism, and to interpret areasof ancient rifting along passive margins. The East AfricanRift System (EARS) stretches for over 3000 km from theRed Sea and Gulf of Aden southward to Mozambique andhas been recognized as a major extensional feature for wellover 100 years [Gregory, 1896]. It is the classic example ofcontinental rifting, generated by subsidence of faultbounded basins coupled with the uplift of rift flanks[Ebinger et al., 1989]. It comprises two branches flankingthe Tanzania craton in central and eastern Africa, and asingle arm that traverses the Ethiopian plateau and craton inthe north (Main Ethiopian Rift). Ongoing research in theMain Ethiopian Rift (MER) has increasingly pointedtoward magmatism as a primary mechanism for extensionand associated crustal modification processes [Keranen etal., 2004; Kendall et al., 2005, 2006; Rooney et al., 2005].This interpretation is consistent with the transition fromfault-dominated rift morphology observed in continentalrifts toward magma-dominated mid-ocean ridge spreadingcenters.[3] We present new geochemical data from several key

locations along the Main Ethiopian Rift between 8! and10!N (Figure 1) in order to assess the interaction of magmas

JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 112, B10201, doi:10.1029/2006JB004916, 2007ClickHere

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1Department of Geosciences, Pennsylvania State University, UniversityPark, Pennsylvania, USA.

2Now at Department of Geological Sciences, Michigan State University,East Lansing, Michigan, USA.

3Department of Geological Sciences, University of South Carolina,Columbia, South Carolina, USA.

4Now at Department of Earth Sciences, University of Bristol, Bristol,UK.

5Department of Earth Sciences, Addis Ababa University, Addis Ababa,Ethiopia.

6Now at Dessie/Kombolcha University, Dessie, Ethiopia.

Copyright 2007 by the American Geophysical Union.0148-0227/07/2006JB004916$09.00

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Figure 1

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with the continental crust in this zone of active extension,and to examine how magmatic processes modify crustalstructure. This approach focuses on geochemical indicatorsof assimilation and fractional crystallization during theresidence of magma within the crust. We integrate studiesof more primitive rocks [Rooney et al., 2005; Furman et al.,2006] from the same eruptive locations and time periods tochart the evolution of magmas at crustal levels. Specifi-cally, we focus on evolved (<6% MgO) lavas eruptedwithin the Wonjii Fault Belt (WFB), considered one of theprimary extensional axes in the MER [Ebinger and Casey,2001; Casey et al., 2006]. Our sampling approach allowsfor an evaluation of magmatic processes throughout thecentral MER, expanding on previous studies of singlevolcanic centers [e.g., Peccerillo et al., 2003]. Our focuson suites of contemporaneous and, often, comagmaticlavas reduces the effects of source heterogeneity that mayhave clouded previous studies [e.g., Trua et al., 1999;Ayalew, 2000].[4] This study of Quaternary eruptives probes ongoing

crustal modification processes. Consequently, our geochem-ical results can be interpreted within the context of substan-tial geophysical investigations of current lithosphericstructure within the MER [Keranen et al., 2004; Bastowet al., 2005; Cornwell et al., 2006; Furman et al., 2006;Maguire et al., 2006; Whaler and Hautot, 2006]. Thealteration of continental crust through assimilation and/orthe presence of fractionating melt in the form of pervasivedikes or large magma chambers will have a significantimpact on crustal structure by replacing existing strata withmagmatic products. We assess small-scale heterogeneities ingeochemical signatures that point to the size and distribu-tion of fractionating magma bodies within the MER. Theseresults will aid in the interpretation of existing geophysicaldata sets that have indicated the presence of melt withinthe crust [e.g., Dugda et al., 2005; Stuart et al., 2006;Whaler and Hautot, 2006] but have been unable todetermine if this melt occupies large shallow magmachambers or more complex diked zones of melt intrusion.Such distinctions are an important indicator of the relativeimportance of magmatism- and fault-based extension incontinental rifting.[5] Our investigation is suggestive of a dual axis of

extension requiring revision of any single-axis model ofEthiopian Rift magmatism. Further, we reveal heterogeneityin the crustal structure, and by inference the active riftingprocesses, beneath the WFB and Silti-Debre Zeyit Fault

Zone (SDFZ) (Figure 1). We attribute these variations to thecoeval propagation of extension into the MER from both thesouth (Kenya) and the north (Red Sea).

2. Geologic and Geodynamic Background2.1. Tectonic Setting

[6] The central MER marks the transition between riftingof thick continental crust in the southern and central EARSand incipient seafloor spreading in northern Afar. Recentvolcanic and tectonic activity in the central MER is gener-ally confined to distinct faults belts [Mohr, 1962, 1967;WoldeGabriel et al., 1990], postulated to be magmaticsegments and precursors to mid-ocean ridge spreadingcenters [Ebinger and Casey, 2001]. These belts of tectono-magmatic activity and associated lithospheric modificationare the WFB [Mohr, 1967] and the SDFZ [WoldeGabriel etal., 1990] (Figure 1). These belts are coincident with crustalextension within the rift, occupying grabens on the easternand western margins, broadly paralleling the rift alignment[WoldeGabriel et al., 1990]. The WFB and SDFZ aredominated by large (up to !90 km3) felsic volcanoes(e.g., Ziqualla, Fantale, Dofan, Bosetti), some with well-developed calderas (Kone, !5 km diameter; Gedemsa,!8 km diameter), spaced 20–45 km apart along the faultbelts. Basaltic activity is restricted to minor flows emanatingfrom cinder cones along faults and fractures, and rare fissureeruptions. The evolution from more widespread magmatismduring the early syn-rift stage toward more restricted axialactivity was accompanied by a shift in locus of strain fromthe rift border faults toward the WFB [Ebinger and Casey,2001] and SDFZ. The current concentration of crustalextension within the SDFZ and WFB makes them excellentlocales for the study of rift-related lithospheric modificationprocesses.

2.2. Geologic History

[7] The volcanics of the SDFZ and WFB erupted througha substantially modified continental crust that has experi-enced considerable volcanism over the past 30 Ma. Exten-sive volcanism in north central Ethiopia, linked to the Afarplume [Pik et al., 1999; Kieffer et al., 2004], is restricted toepisodic eruptions between 30 Ma and 11 Ma (pre-riftstage). Rejuvenated magmatism from !11 Ma to presentis coincident with the onset and progression of continentalrifting in north central Ethiopia [Hart et al., 1989;WoldeGabriel et al., 1990; Chernet and Hart, 1999]. Pre-rift

Figure 1. Shaded relief digital elevation model of the central MER, constructed using NASA-JPL Shuttle RadarTopography Mission (SRTM) data with a 90 m resolution. The individual 1 by 1 degree tiles were corrected for holes andcombined to produce a larger picture of the MER. The projection is based on the WGS-84 datum. (a) Cinder conefrequency in the Wonjii Fault Belt and Silti-Debre Zeyit Fault Zone in the north central MER. Cinder cones are shown assmall dots [Berhe and Wondm-Agennehu, 1978; Benvenuti et al., 2002; Abebe et al., 2005]. The Boru-Toru Structural High[Bonini et al., 2005] is labeled BTSH. Northward toward Dofan, our field investigations indicate that cinder cone activity ismuch reduced compared with the more southerly section of the study area, but no geologic mapping exists for this region.(b) Sample locations (dots), roads (solid lines), towns and cities (starred), and lakes (grey) overlain. The large volcanicfeatures of Dofan, Fantale, Kone, Bosetti, Gedemsa, Tullu Moye, and Ziqualla are labeled and represented by grey circles.The Wonjii Fault Belt (WFB) and Silti-Debre Zeyit Fault Zone (SDFZ) extend beyond the edges of this diagram but areroughly shaded in white.

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volcanic activity is dominated by a brief (!1 my [Hofmannet al., 1997]) episode of continental flood basalt volcanismthat extruded up to a 2 km thick succession over an area of600,000 km2 [Mohr, 1983; Mohr and Zanettin, 1988]. Theflood basalt series are dominated by transitional to tholeiiticmafic lavas [Pik et al., 1998] with interlayered felsic vol-canics in the upper levels [Ayalew et al., 1999]. Isolatedbimodal shield volcanoes, composed of mafic and lesserfelsic eruptives, rise up to 1200 m above the flood basaltplateau [Kieffer et al., 2004]. Shield volcanoes have eruptedepisodically with activity recorded at 30 Ma (Simien), 23 Ma(Choke and Guguftu) and 10 Ma (Guna) [Kieffer et al.,2004]. The onset of rifting within the central Ethiopian Rift at!11 Ma [Wolfenden et al., 2004] was coincident with a shiftin magmatism toward felsic and subordinate basaltic erup-tives located in discrete bimodal volcanic centers within therift [Hart et al., 1989; Chernet and Hart, 1999], which alsogenerated the widespread ignimbrites that currently blanketthe rift floor. Subsequent to !3 Ma magmatic activity wasfurther focused within the rift, now centering on WFB andSDFZ.[8] Magmatism and lithospheric modification have been

coupled during the 30 Ma history of the Ethiopian volcanicprovince. The widespread magmatic event at 30 Ma, asso-ciated with the Afar plume, is linked to infiltration of meltand the thermomechanical erosion of the lower lithosphererecorded in metasomatized mantle xenoliths [Bedini et al.,1997]. Magmatism and lithospheric modification were onceagain coupled as the onset of rifting at 11 Ma was accompa-nied by a shift in regional magmatism, which became focusedon the newly formed rift. More recently, pervasive magmaticintrusions and diking have further modified the lithospherebeneath the WFB and SDFZ [e.g., Rooney et al., 2005;Kendall et al., 2006]. In this study we will concentrate onthe effects of modern magmatism as it modifies the crustalportion of the lithosphere within the WFB and SDFZ. Byfocusing onmore evolved volcanic products we will examinemagmas that have spent time differentiating within the crust,recording lithospheric processes.

2.3. Geophysical/Geodynamic Background

[9] Most passive margins world-wide are considered tobe magmatic, on the basis of the observation that they arecharacterized by thick sequences of extruded and intrudedigneous rocks that were emplaced prior to or in conjunctionwith the onset of rifting [e.g., Menzies et al., 2002].However, the breakup history of most magmatic passivemargins remains enigmatic as the ocean-continent boundaryis concealed by thick seaward dipping reflectors [e.g.,Holbrook and Kelemen, 1993] and rendered obscurethrough erosion and thermal subsidence. The northern EastAfrican Rift System provides the perfect setting in which tostudy the early stages of continental breakup. Within theregion magmatic and tectonic features are well-exposed,and in some cases are currently active (e.g., Dabbahu Rift inAfar [Wright et al., 2006]).[10] The processes of lithospheric extension and rupture

have been a source of much debate that centered, in largepart, on active versus passive rifting models [Sengor andBurke, 1978; Turcotte and Emerman, 1983]. Passive riftingmodels, where extension is generated by far-field platestresses, include models of pure shear [McKenzie, 1978]

and simple shear [Wernicke, 1985]. Active rifting modelsexplain extension as the result of asthenospheric upwelling[Morgan, 1971; Burke and Dewey, 1973]. Shortcomings ofthese end-member models prompted recent work that com-bines processes associated with active and passive riftinginto hybrid models [Zeyen et al., 1997; Courtillot et al.,1999; Huismans et al., 2001; Ziegler and Cloetingh, 2004].Strictly kinematic models of continental breakup [e.g.,Wernicke, 1985; Lister et al., 1986] often ignore the effectsof melt in the rifting process, and in doing so they ignore theobservation that the average tectonic force needed to initiateand maintain breakup may be an order of magnitude greaterthan that which is commonly available [Kusznir and Park,1987; Hopper and Buck, 1993; Zeyen et al., 1997; Buck,2004]. In Ethiopia, magmatic intrusion facilitating litho-spheric rupture and extension was first outlined by Mohr[1987], who suggested asthenospheric material is diapira-cally emplaced into the lithosphere. Ebinger and Casey[2001] presented a geodynamic model for the MER thatsuggested the initial stage of lithospheric extension isaccommodated by slip along rift border faults. Over time,extensional stresses become focused in the central portionof the rift valley through dike intrusion and associatedfaulting as the rift border faults become inactive. Thesegeodynamic models are excellent tools in the study ofcontinental rifting and new data collected during the recentEthiopia Afar Geoscientific Lithospheric Experiment(EAGLE) [Maguire et al., 2003] project have brought intosharp focus, the role of magmatism and faulting duringrifting and the transition toward seafloor spreading. Inparticular, these new data shed light on how magmatisminteracts with and modifies the continental crust.[11] EAGLE controlled-source seismological data indi-

cate limited lower-crustal thinning northward toward Afarwith almost complete replacement of the lower crust byintrusive mafic volcanism [MacKenzie et al., 2005;Maguireet al., 2006]. Such magmatic replacement of the crust hasalso been suggested farther south on the basis of teleseismicreceiver function analyses [Dugda et al., 2005]. Magneto-telluric investigations reveal shallow (<1 km) and mid-crustal zones of high conductivity, interpreted as partiallymolten material located beneath the surface manifestation ofQuaternary volcanism within the MER. The upper crustwithin these zones is imaged as seismically fast [Keranen etal., 2004;MacKenzie et al., 2005] and dense [Mahatsente etal., 1999; Cornwell et al., 2006], suggestive of cooled maficintrusions. Partial melt is also thought to reside withinvertically oriented dikes in the lithosphere on the basis ofseismic anisotropy studies across the MER [Kendall et al.,2005]. Low velocity P and S wave anomalies dominate theEthiopian upper mantle structure [Benoit et al., 2003;Bastow et al., 2005], and indicate the presence of partialmelt in the upper 300 km of the mantle [Bastow et al.,2005]. We therefore have the opportunity to couple geo-chemical and geophysical results to produce a synthesismodel of incipient continental breakup.

3. Methods: Sampling and Sample Preparation

[12] Fresh samples of basalt, trachybasalt and basaltictrachyandesite (Figure 2) were collected in 2002 and 2003from cinder cones and associated flows from 41 localities

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along the WFB (Figure 1). The samples were cut into slabs,trimmed to remove surface alteration and polished toremove saw marks. Slabs were then crushed using aporcelain jaw-crusher and powdered for <1 minute with atungsten-carbide disc mill.[13] Twenty-one samples were dissolved and analyzed at

Duke University using a VG PlasmaQuad-3 ICP-MS fortrace elements (including REE) and an ARL-Fisons Spec-traspan 7 DCP for major elements and selected traceelements. All trace element data were obtained by ICP-MS analysis, excluding Sr and Ba which were undertakenby DCP. These data are presented in Table 1; analytical dataon the remaining 20 primitive samples (>6% MgO) arepresented by Furman et al. [2006].[14] Elemental analysis of minerals was undertaken at

The Pennsylvania State University using a Cameca SX-50 electron microprobe equipped with 4 wavelength spec-trometers and one energy dispersive spectrometer. Theanalyses used a 12 nA sample current and a 15 kVaccelerating voltage. The width of the electron beam was20 microns for analysis of all minerals excluding spinelswhere a beam width of !1–3 microns was used. Countingtimes were 20 seconds per element peak and 10 secondsbackground for major elements while minor elements (0.1–1%) had a counting time of 60 seconds per peak and30 seconds background. The data are presented in auxiliarymaterial1 Tables S1–S4 and include unpublished micro-probe analyses from both the primitive samples describedby Furman et al. [2006] and the evolved samples which arethe focus of this study.

4. Results4.1. Petrography

[15] All samples are hypocrystalline and many are vesic-ular. The fabric of most samples is inequigranular butobserved textures include porphyritic, glomeroporphoritic,seriate and rarely poikilitic. Plagioclase crystals in someflows exhibit trachytic textures, but in the majority ofsamples the crystals are unaligned. Intergrowths of all threedominant phenocryst phases (olivine, plagioclase feldsparand subordinate clinopyroxene) are observed. These phe-nocrysts are set in a fine-grained matrix of plagioclasefeldspar, opaque oxides and to a lesser extent clinopyroxeneand olivine. Examples of olivine, clinopyroxene and pla-gioclase feldspar crystals in disequilibrium (distinguishedby resorption textures) are common.[16] Olivine core compositions range from Fo52 in N-03

to Fo77 in N-24 and contain NiO and Cr2O3 below theinstrument detection limit (0.10 wt.%). Clinopyroxenecompositions straddle the diopside-augite boundary andare enriched in Al2O3 ("7.25%) and TiO2 ("3.5%). Pla-gioclase feldspars range primarily from An50–85; notableexceptions are two crystals of anorthosite in N-3. Opaqueoxides in the matrix are rich in both FeO (65–70 wt.%) andTiO2 (19–27 wt.%).

4.2. Major Elements

[17] Quaternary volcanic flows and scoria associated withcinder cones along the WFB are predominantly alkaline to

transitional basalts with subordinate trachybasalts and ba-saltic trachyandesites (Figure 2). A distinct Daly Gap isevident, with intermediate material (54–62 wt.% SiO2) rare.We focus here on products with 46–51 wt.% SiO2.[18] Major element patterns are broadly consistent with

fractionation of the observed mineral phases. Among maficlavas, major element trends appear to reflect removal ofolivine, clinopyroxene and plagioclase feldspar. Samplesstudied here exhibit positive trends of Na2O, K2O, MnOand P2O5 with decreasing MgO, while CaO and TiO2

exhibit negative trends among samples with <5 wt.%MgO. Examination of Figure 3 reveals patterns consistentwith dominant crystal fractionation such as the rise in TiO2

and Al2O3 (not shown) concentrations followed by a fall inboth. The decreasing CaO/Al2O3 values indicate clinopyr-oxene is an important fractionating phase among sampleswith >6% MgO, whereafter plagioclase becomes increas-ingly important (Figure 4). The final fractionating phasesare Fe-Ti oxides, indicated by a drop in TiO2 at about5% MgO.[19] Two fractionation trends are distinguished by CaO/

Al2O3 values: Trend A, which is spatially restricted tosamples surrounding Lake Besaka, and Trend B, whichcomprises the majority of samples in the region. Such trendsare the result of variable clinopyroxene and plagioclasefractionation, addressed in later sections. A third field,defined by low MgO and high CaO/Al2O3 encompassessamples from the fissure eruption at Fantale and other flowsbetween Fantale and Dofan (N-24 from south of TulluMoye also plots in this field).

4.3. Trace Elements

[20] Abundances of compatible trace elements Sc, Cr andNi decrease with decreasing MgO, consistent with olivineand clinopyroxene fractionation (Figure 5). Vanadium abun-

Figure 2. Total alkali-silica diagram [Le Bas et al., 1986]showing all Quaternary MER samples including the moreprimitive WFB (Wonjii Fault Belt) basalts presented byFurman et al. [2006] and SDFZ [Rooney et al., 2005].

1Auxiliary materials are available at ftp://ftp.agu.org/apend/jb/2006jb004916.

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Table 1. (continued)

Sample La Ce Pb Pr Nd Sm Zr Hf Eu Gd Tb Dy Y Ho Er Yb Lu Co Cr Ni V Sc

1026 18.9 44.3 1.48 6.08 26.8 6.41 159 3.75 2.32 6.87 1.06 6.17 31.3 1.20 3.20 2.75 0.41 41.6 58.5 34.0 305 29.71028 35.3 84.4 2.21 11.34 49.1 11.26 275 6.34 4.40 11.90 1.84 10.32 39.6 2.04 5.41 4.75 0.72 14.5 8.1 11.0 36 23.41029 27.9 68.5 1.48 9.92 46.0 10.60 142 3.56 4.01 11.30 1.65 8.79 40.3 1.68 4.17 3.33 0.49 34.1 10.4 19.1 302 35.01024 25.1 57.7 1.63 7.81 34.3 7.88 169 4.10 2.85 8.29 1.26 6.96 33.0 1.36 3.59 3.08 0.46 46.9 20.6 25.2 300 33.21024A 24.8 56.4 2.18 7.72 33.6 7.67 169 4.09 2.74 8.10 1.22 6.73 32.1 1.34 3.45 3.01 0.46 42.2 20.2 24.3 336 32.51025 19.7 44.7 2.19 5.93 25.3 5.80 157 3.64 2.03 6.09 0.93 5.16 28.3 1.00 2.65 2.29 0.35 39.2 65.6 34.0 340 29.01034 29.3 69.7 2.53 9.60 43.2 9.90 170 3.90 3.83 10.38 1.46 8.09 36.0 1.54 3.96 3.29 0.49 36.1 9.7 14.4 300 33.51017 30.2 70.0 1.89 9.86 43.7 9.89 191 4.59 3.77 10.52 1.51 8.29 35.1 1.56 3.95 3.39 0.48 38.4 9.5 16.1 285 33.61021 29.2 68.4 1.89 9.60 42.7 9.57 166 3.99 3.72 10.20 1.46 8.12 34.7 1.51 3.87 3.34 0.47 35.5 10.4 15.7 284 32.6N-19 28.4 66.7 1.9 8.8 41.4 8.84 162 4.0 3.47 9.80 1.51 7.45 37.3 1.42 3.59 2.99 0.43 31.9 11 20 278 28.81040 22.6 57.7 3.30 7.04 30.0 6.51 199 4.68 2.17 6.92 1.02 5.49 31.8 1.07 2.80 2.42 0.37 31.6 7.2 13.2 237 20.21039 15.3 33.4 1.53 4.45 18.8 4.40 111 2.82 1.56 4.27 0.70 3.92 20.6 0.77 1.98 1.71 0.27 62.0 110 55.9 292 32.2N-13 40.5 90.7 2.3 11.5 50.1 10.37 275 6.1 3.40 11.29 1.77 9.07 50.7 1.72 4.75 3.89 0.61 19.4 8 16 143 21.1N-07 32.0 70.9 3.3 9.7 39.6 8.30 230 5.3 2.84 8.87 1.42 7.17 39.6 1.41 3.59 3.15 0.46 26.6 11 20 186 22.7N-06 32.5 72.5 3.3 9.4 40.3 8.53 234 5.4 2.93 9.19 1.45 7.23 40.4 1.44 3.65 3.22 0.47 27.9 10 20 203 23.9N-05 32.1 70.8 3.2 8.9 39.4 8.25 232 5.2 2.84 8.96 1.42 7.07 39.9 1.39 3.52 3.12 0.47 33.1 8 28 182 22.3N-04 28.1 61.5 2.0 8.0 34.3 7.18 192 4.5 2.63 7.64 1.20 6.02 33.4 1.13 3.00 2.60 0.40 34.1 11 22 278 27.2N-03 28.8 63.8 2.4 8.2 35.9 7.50 192 4.4 2.67 8.02 1.26 6.27 34.2 1.23 3.09 2.66 0.41 37.4 10 24 255 26.8N-25 32.2 62.2 4.2 8.4 34.9 7.21 220 5.0 2.27 7.98 1.23 6.14 34.8 1.20 3.21 2.95 0.43 45.8 29 37 368 31.4N-24 17.8 39.2 1.7 5.1 22.2 4.81 139 3.2 1.70 5.22 0.85 4.38 23.3 0.81 2.18 1.80 0.27 48.4 42 45 389 35.7N-23 25.7 56.0 3.9 7.1 30.2 6.26 209 4.8 1.99 6.47 1.05 5.35 29.2 1.01 2.76 2.41 0.37 41.4 43 43 336 28.8

aUSGS standards (W-2, Diabase; DTS-1, Dunite; DNC-1, Diabase; AGV-1, Andesite; G-2, Granite; SDC-1, Mica Schist; BIR-1, Basalt), NBS/NISTstandard (688 Basalt) and two standards developed at Lamont Doherty Earth Observatory (K1919a, Basalt; AII-92, Basalt) were used. The precision basedon duplicate analysis is typically better than 1% for SiO2, MgO, Fe2O3, and Al2O3; better than 2% for CaO, MnO2, TiO2, Na2O, Sc, Rb, Sr, Y, Zr, Nb, Ba,La, Ce, Pr, Nd, Sm, Gd, Ho, and Er; better than 3% for V, Tb, Dy, Yb, Lu, Hf, Ta, Th, and Eu; and better than 5% for K2O, U, Pb, Cs, Ni, and Cr.

Table 1. Major, Minor, and Trace Element Analysis of the MER Basaltsa

Sample Volcano Lat Long SiO2 Al2O3 Fe2O3 MgO CaO Na2O K2O TiO2 MnO2 P2O5 SUM Sr Ba* Cs Rb Th U Nb Ta

1026 Dofan 9.478 40.198 47.16 15.07 14.28 5.42 9.70 3.44 0.65 3.14 0.23 0.62 99.72 395 246 0.15 14.1 1.67 0.47 28.7 1.841028 Dofan 9.308 40.144 52.69 14.35 13.70 2.47 6.16 4.96 1.20 2.16 0.36 0.95 99.00 443 521 0.31 28.7 3.19 0.83 49.9 2.951029 Dofan 9.322 40.176 46.05 13.02 16.37 5.11 9.91 3.68 0.67 3.93 0.31 1.74 100.79 459 377 0.07 11.1 1.65 0.21 30.9 1.971024 Dofan 9.082 39.911 47.04 14.29 16.14 4.32 9.89 3.35 0.70 3.43 0.25 0.80 100.20 464 397 0.08 12.7 1.87 0.39 31.3 1.931024A Dofan 9.082 39.911 47.30 13.99 15.65 4.63 9.36 3.69 0.84 3.35 0.25 0.79 99.84 444 421 0.05 17.8 1.96 0.49 31.2 1.931025 Dofan 9.037 40.059 47.67 16.33 13.21 4.63 10.32 3.37 0.78 2.73 0.19 0.49 99.73 454 376 0.06 15.3 1.61 0.47 25.4 1.541034 Fantale 8.930 39.908 47.02 13.73 15.48 4.68 9.15 3.92 0.90 3.84 0.29 1.13 100.13 495 514 0.16 19.9 2.14 0.57 36.0 2.281017 Fantale 8.905 39.908 47.40 13.72 15.29 4.69 9.12 3.90 0.93 3.82 0.29 1.14 100.29 501 509 0.22 21.3 2.41 0.66 36.2 2.321021 Fantale 8.904 39.888 46.46 13.66 15.28 4.65 9.06 3.90 0.88 3.82 0.29 1.13 99.13 484 507 0.24 21.6 2.21 0.61 35.3 2.26N-19 Fantale 8.903 39.893 46.69 13.73 15.39 4.67 9.03 4.01 1.03 3.86 0.29 1.16 99.86 448 500 0.07 11.6 1.92 0.61 35.4 2.21040 Kone 8.720 39.555 48.61 14.78 13.76 4.34 8.24 4.17 1.31 3.19 0.25 0.78 99.43 390 386 0.14 18.8 2.23 0.59 37.9 2.411039 Kone 8.716 39.553 46.95 18.20 11.07 5.62 11.49 2.85 0.60 1.99 0.16 0.30 99.22 504 191 0.09 10.1 1.64 0.36 19.8 1.28N-13 Kone 8.667 39.483 50.92 14.86 12.17 3.82 7.23 4.58 1.97 2.91 0.25 1.17 99.88 483 504 0.37 35.0 4.34 1.10 47.4 2.9N-07 Gedemsa 8.317 39.108 50.20 15.02 12.94 4.06 8.04 4.32 1.53 3.21 0.25 0.67 100.23 562 359 0.32 24.8 3.19 0.88 42.5 2.6N-06 Gedemsa 8.271 39.146 50.43 15.07 12.94 4.02 7.92 4.38 1.30 3.21 0.25 0.65 100.17 572 356 0.30 23.5 3.23 0.89 43.2 2.6N-05 Gedemsa 8.267 39.142 49.96 14.88 12.83 4.00 7.94 4.24 1.39 3.20 0.25 0.65 99.32 564 351 0.34 24.9 3.11 0.90 43.1 2.5N-04 Gedemsa 8.258 39.175 48.86 14.94 14.64 4.76 8.85 4.03 1.00 3.64 0.23 0.61 101.57 533 324 0.11 21.3 2.61 0.64 37.8 2.3N-03 Gedemsa 8.254 39.188 48.16 14.82 14.28 4.64 8.81 3.89 1.19 3.57 0.24 0.68 100.25 555 325 0.18 19.9 2.73 0.75 38.1 2.4N-25 Gedemsa 8.024 39.036 48.26 14.59 14.57 5.04 8.78 3.37 1.38 3.15 0.21 0.46 99.81 462 424 0.23 26.8 3.18 0.72 35.2 2.1N-24 Gedemsa 8.018 39.044 46.40 14.75 14.77 5.80 10.79 3.15 0.88 2.98 0.20 0.36 100.09 408 224 0.09 11.4 1.86 0.41 25.7 1.5N-23 Gedemsa 8.005 39.091 48.69 15.83 13.20 5.00 9.19 3.40 1.22 2.79 0.19 0.42 99.94 497 379 0.20 21.7 2.96 0.67 32.6 1.9

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dances are nearly constant in samples with >5 wt.% MgOwhereafter they decrease with decreasing MgO (Figure 5),indicting the onset of Fe-Ti oxide fractionation.[21] Large-ion lithophile elements (LILE) exhibit hetero-

geneous behavior: Sr and Ba (not shown) do not formcoherent trends against MgO, whereas Rb exhibits a cleartrend of enrichment throughout the suite, increasing rapidlyat lower MgO contents. LILE ratios such as Rb/Sr and Ba/Rb are also not correlated with MgO content (Figure 6). Theanomalous behavior of the LILE are further highlighted bythe lack of correlation between Sr and Ba versus individualhigh field strength elements (HFSE), in contrast to the clearpositive correlation of Rb with Zr (Figure 6). Interestingly,Sr, Ba and Rb correlate well with La values. Abundances of

the HFSE increase gradually with decreasing MgO amongmafic lavas, but this enrichment becomes more pronouncedin samples with <5% MgO. HFSE abundances are stronglycoupled with one another (e.g., Nb versus Zr) and with therare earth elements (REE; e.g., Nb versus La).[22] REE exhibit similar trends to the HFSE, gradually

increasing in samples with 10–5% MgO; below 5% MgO,REE concentrations rise rapidly. Chrondrite normalizedREE patterns [see Furman et al., 2006, Figure 3] show aminor HREE depletion (Tbn/Ybn = 1.68–2.23) similar tothat observed in previous studies of Ethiopian rift basalts(1.37–2.34 [Trua et al., 1999]) but not of the samemagnitude as the 30 Ma HT2 plume-derived basalts(2.45–2.80 [Pik et al., 1999]). Minor positive Eu anomaliesare observed in some samples from the Fantale-Dofan area(Figure 7; 1028, 1029, 1021, N-19, 1017).[23] Indicators of crustal contamination (e.g., La/Nb, Ce/

Pb; Figure 8) are generally within the range of uncontam-inated mantle-derived lavas; a few anomalously low (N-21,N-23, N-25) and high (N-13, N-19, 1017, 1021, 1024, 1028,1029) Ce/Pb values require some modification. Significant-ly, samples with high Ce/Pb values also have anomalouslyhigh P2O5 and low Hf/Sm, suggesting a role for apatite(Figure 9); a subset of these same samples also has positiveEu anomalies (Figure 7). A group of samples collectedwithin and south of Kone Caldera (N-22, N-20, 1036)exhibit anomalously high concentrations of the most in-compatible trace elements (e.g., Cs, Rb, Nb) while lessincompatible elements show no such heterogeneity (e.g.,Yb, Zr; Figure 5). These same samples also have highnormative nepheline contents (4.9–6.5%) suggesting theywere generated by a smaller degree partial melt and/or adeeper source with respect to other Kone and WFB samples[Furman et al., 2006], similar to samples from the SDFZ.[24] The heterogeneity observed in incompatible trace

element variations, despite an apparent lack of widespreadFigure 3. Major element X-MgO diagrams. The moreprimitive data plotted here (greater than 6% MgO) arepresented by Furman et al. [2006]. SDFZ data are presentedby Rooney et al. [2005]. Data have been normalized on awater-free basis after FeO:Fe2O3was calculated on the basis ofa 85:15 ratio. Clear trends exist above and below !6% MgO.Greater than 6% MgO, olivine is the primary fractionatingphase, highlighted by the relatively constant values of otherelements. These trends are covered more fully in the text. Allvalues are presented in wt.%.

Figure 4. Variation of CaO/Al2O3 with MgO. Symbols arethe same as in Figure 3. The common fractionation trendfrom 10 to 7 wt.% MgO bifurcates into a steeper trend ‘‘A’’and a more shallow trend ‘‘B’’ which dominates thefractionation path. Trend ‘‘C’’ appears to be generated bythe addition of a component with high CaO/Al2O3 andlower MgO. Silti-Debre Zeyit Fault Zone lavas are shownfor comparison [Rooney et al., 2005].

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crustal assimilation, coupled with the compositional varia-tion documented at Kone argues for some degree of sourceheterogeneity in the WFB lavas. However, the consistenttrends in major and most trace elements are broadlyindicative of fractional crystallization of observed phases.

5. Discussion5.1. Crustal Assimilation

[25] In the study of mafic rocks erupted in continentalsettings we acknowledge the perennial problem of detectingand assessing the impact of crustal assimilation. The broadrange of lithologies for potential crustal assimilants hasprompted the application of multiple petrographic (e.g.,embayed feldspars, disaggregated crustal xenoliths) andgeochemical (e.g., Ce/Pb, Ti/Yb, K/P, La/Nb, Sr and Pbisotopes) indicators of crustal contamination. Crustal assim-ilation has been documented to varying degrees throughout

the region from Afar in the north [Deniel et al., 1994] toTurkana in the south [Furman et al., 2004] but is generallyminor. Hart et al. [1989] used K/P, Ti/Yb, and Sr-Ndisotopic enrichment patterns to indicate significant crustalassimilation of basalts erupted from 11-6 Ma on the Ethi-opian Plateau, also noting however, that mixing betweensilicic and basaltic melts can have a significant impact oncomposition elsewhere in the rift (e.g., west central Afar).Other authors have also indicated a role for mixing ofevolved magma or solidified silicic material with fraction-ating basalt to produce the range of observed mafic Qua-ternary lavas in the Ethiopian Rift [Gasparon et al., 1993;Peccerillo et al., 2003]. However, the effect of crustalcontamination on MER lavas, in particular the question ofwhether the potential assimilant was of upper or lowercrustal lithology, remains unresolved [e.g., Boccaletti etal., 1995; Chernet and Hart, 1999; Pik et al., 1999; Trua

Figure 5. Incompatible and compatible trace element X-MgO variation diagrams for Wonjii Fault Beltand Silti-Debre Zeyit lavas. Primitive samples plotted here are presented by Furman et al. [2006]. Valuesfor Debre Zeyit and Butajira basalts are presented by Rooney et al. [2005]. All values are ppm unlessotherwise specified.

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et al., 1999; Kieffer et al., 2004]. Upper crustal contamina-tion has been suggested to produce low Ti/Yb (>3,500 [Hartet al., 1989]) and high K/P (<12 [Hart et al., 1989]),coupled with elevated 87Sr/86Sr (0.706–0.765 [Hart et al.,1989; Pik et al., 1999; Trua et al., 1999]) and 206Pb/204Pb(!18.6 [Trua et al., 1999]). The upper crust, represented byEthiopian meta-volcanics [Pik et al., 1999] or by Sudanesegranites [Trua et al., 1999], is interpreted as playing only aminor role in Quaternary MER lava petrogenesis [Boccalettiet al., 1995; Chernet and Hart, 1999; Trua et al., 1999].Lower crustal assimilation has been suggested to play amore substantial role in the contamination of mafic Ethio-pian lavas [Boccaletti et al., 1995; Chernet and Hart, 1999;Pik et al., 1999; Trua et al., 1999; Kieffer et al., 2004]; thisassimilant is typically identified by high Ti/Yb (in excess of2000 [Hart et al., 1989]) and La/Nb (greater than 2 [Pik etal., 1999]), and low K/P (!3–12 [Hart et al., 1989]), Ce/Pb(>1 [Pik et al., 1999]) and 206Pb/204Pb (16.93–17.63 [Truaet al., 1999]). The lower crust, represented by Sudanesegranulite [Hart et al., 1989; Boccaletti et al., 1999] orSudanese granite [Trua et al., 1999] is modeled to con-taminate Quaternary basalts in the MER by !10–13%[Boccaletti et al., 1995; Trua et al., 1999]. However, thesestudies did not account for heterogeneity in parentalmagmas [Rooney et al., 2005; Furman et al., 2006] anddid not use Ethiopian crustal compositions, rendering suchnumerical estimates unreliable.[26] Previous studies in the southern WFB recorded

basaltic enclaves within rhyolitic eruptives and felsic xen-oliths within basaltic lavas and scoria [Peccerillo et al.,2003]. Xenoliths were not observed in our study, but there

is petrographic evidence for magmatic interaction. FeO-richolivines (Fo52–57) in sample N-7 have anomalous Fe-Mgexchange coefficients (KD = FeOcrystal #MgOhost/FeOhost #MgOcrystal) of 0.55–0.67 and are apparently xenocrysts.Samples N-7 and N-17 contain xenocrysts of more evolvedminerals that are similar to the phenocryst assemblage ofN-19, suggesting mixing between primitive and evolved

Figure 6. Key trace element ratio and correlation plots for Wonjii Fault Belt and Silti-Debre Zeyit lavas.Primitive samples plotted here are presented by Furman et al. [2006]. Values for Debre Zeyit and Butajirabasalts are presented by Rooney et al. [2005]. All values are ppm unless otherwise specified.

Figure 7. Chondrite normalized REE pattern for WonjiiFault Belt lavas. The shaded field represents all samplesfrom the Wonjii Fault Belt and includes primitive samplespresented by Furman et al. [2006]. Note mild positive Euanomaly for sample 1028. Normalizing factors are fromBoynton [1984].

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basaltic magmas. Sample N-3 exhibits substantially hetero-geneous populations of clinopyroxene and feldspars that areclearly xenocrystic. These compositions are observed infelsic eruptives investigated nearby [e.g., Trua et al., 1999;Peccerillo et al., 2003]. Interestingly, these samples (N-3,N-7 and N-17) do not exhibit a marked decrease in Ce/Pbthat is characteristic of nearby felsic lavas (Figure 8),suggesting the volume of material mixed into these lavas

is likely to be small. Notably, no pan-African crustalxenoliths have ever been reported in studies of MER lavas.[27] Ce/Pb and La/Nb values (Figure 8) can be used to

evaluate potential contamination by both upper crustal(basement rocks from Ethiopia [Kebede et al., 1999; Sifetaet al., 2005; Tadesse and Allen, 2005; Yihunie et al., 2006])and lower crustal lithologies (LT basalts [Pik et al., 1999;Kieffer et al., 2004]). The majority of the WFB and SDFZ

Figure 8. Crustal contamination indicators, Ce/Pb and La/Nb, are plotted for Ethiopian samples.(a) Crustal rocks from Ethiopia [Kebede et al., 1999; Sifeta et al., 2005; Tadesse and Allen, 2005; Yihunieet al., 2006] and 30 Ma flood basalts that exhibit a lower crustal contamination [Pik et al., 1999; Kieffer etal., 2004]. The WFB and SDFZ lavas lie predominantly within the mantle field but extend to lowervalues of Ce/Pb, which may indicate crustal contamination. Importantly, a subset of samples plots at highCe/Pb and cannot represent contamination with crustal materials. (b) The WFB and SDFZ lavas areplotted with samples from Tullu Moye [Trua et al., 1999] and Gedemsa [Peccerillo et al., 2003]. Samplesfrom Tullu Moye that lie within the mantle array (25 ± 5 [Hofmann et al., 1986]) plot at higher La/Nb,consistent with some lower crustal contamination of these lavas noted by Trua et al. [1999]. Gedemsasamples form an array with from low Ce/Pb to moderate Ce/Pb at relatively unenriched La/Nb (excepting3 enclaves analyzed). Wonjii Basalts from Gedemsa and evolved rocks (>65% SiO2) from the Tulu Moyeregion are marked by filled symbols.

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samples have La/Nb and Ce/Pb values that plot within therange observed in mantle-derived basalts [Hofmann et al.,1986]. Samples with high Ce/Pb cannot be the result ofEthiopian crustal assimilation (Figure 8). Evolved productsof the southern WFB have low Ce/Pb and moderate to lowLa/Nb values [Trua et al., 1999; Peccerillo et al., 2003],suggesting they are an appropriate assimilant for our sam-ples with anomalously low Ce/Pb. Basalts erupted withinthe Gedemsa caldera which show petrographic evidence ofinteraction with intrusive sialic rocks form an array ofdecreasing Ce/Pb at relatively constant La/Nb (Figure 8);samples N-25, N-23 and N-21 plot within this array,suggesting contamination with evolved magmas that areco-genetic or with the Ethiopian crust.[28] Samples with elevated Ce/Pb values (1028, 1029,

1024, N-19, 1017, 1021) are spatially restricted to theregion between Fantale and Dofan. These samples arecharacterized by unusually high CaO/Al2O3 and lowMgO, and exhibit small positive Eu anomalies consistentwith feldspar assimilation. The unidentified feldspar-bearingassimilant also contains apatite, suggested by anomalouslyhigh P2O5 and Ce/Pb, low Hf/Sm (Figure 9). Apatite isfound in silicic rocks in the region and the fractionation ofa plagioclase-rich assemblage with accessory apatite hasbeen invoked to produce silicic rocks from mafic magmasin the MER [Trua et al., 1999]. These data suggest that theunidentified assimilant is a cumulate related to the frac-tionation of more primitive lavas. This interpretation sup-ports a model whereby at least some of the silicic productsof the MER are derived through fractional crystallizationof primitive magma [e.g., Ayalew et al., 1999] consistentwith ponding and fractionation of plume-derived melts atthe base of the crust as predicted from coupled petrologic-numerical modeling [Farnetani et al., 1996], and geophys-ical evidence of high velocity lower crust [MacKenzie etal., 2005].

[29] We observe no correlation between Ce/Pb and206Pb/204Pb, weak correlation between Ce/Pb and 87Sr/86Sr,and no correlation between 207Pb/206Pb and 87Sr/86Sr(Figure 10), suggesting any potential contaminant hadisotopic ratios very similar to those of the primary basalts.Unfortunately, the lack of isotopic data for the Ethiopianpan-African crust makes it difficult to rule out this compo-nent as a potential contaminant for Quaternary lavas.

Figure 9. Variation of P2O5 with Hf/Sm ratio. Smallquantities of apatite will significantly lower the Hf/Sm ratiowhile elevating P2O5. Samples identified with high Ce/Pbvalues show deflection toward apatite assimilation. Theapatite is an accessory mineral in the fractionatingassemblage required to generate evolved MER lavas fromprimitive basalts [Trua et al., 1999].

Figure 10. Trace element and isotopic ratios sensitive tocrustal contamination. A weak correlation is observedbetween Ce/Pb and Sr isotopes, but no such correlation isobserved with Pb isotopes. Trace element data for the SDFZare from Rooney et al. [2005]. Isotopic data for WFB arefrom Furman et al. [2006] and Rooney [2006].

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[30] An important conclusion from these data is that themajority of WFB lavas do not show clear evidence ofinteraction with the continental crust. This interpretationcontrasts with observations on older lava series (low-Ti30 Ma flood basalts [Pik et al., 1999]; 11-6 Ma pre-rift lavas[Chernet and Hart, 1999]). This comparison suggests eitherthat (1) crustal modification by basaltic intrusives within theWFB has replaced the pre-rift continental crust with morerecent rift-associated magmas and/or (2) crustal residencetime is substantially less for Quaternary lavas than for theolder flood and fissure basalts [e.g., Chernet and Hart,1999]. Geophysical evidence supports the hypothesis ofsubstantial intrusions and crustal modification within theMER [e.g., Keranen et al., 2004; Dugda et al., 2005].[31] To document the impact of magmatism on litho-

spheric structure, we first determine the conditions offractionation. The maximum possible storage and fraction-ation depth for the SDFZ and WFB lavas is limited by thedepth of melt generation. Previous studies indicated thatthese Quaternary magmas are generated at !15–25 kbar(!50–90 km) [Rooney et al., 2005; Furman et al., 2006].This melting zone overlies low velocity anomalies inter-preted as partial melt in the depth range 75–250 km[Bastow et al., 2005]. The presence of a high velocitylower crustal unit interpreted as dense mafic cumulates[MacKenzie et al., 2005] raises the possibility that magmasmay fractionate at the base of the crust (!30–35 km).Clinopyroxene barometry [Nimis and Ulmer, 1998] indi-cates that neither the WFB nor the SDFZ show appreciablealong-axis variations in the depth of fractionation. How-ever, WFB lavas fractionate at depths of less than 5 kmand at modest temperatures (1050–1150!C; Table 2) while

the SDFZ lavas undergo fractionation at higher temperatures(1126–1362!C) and at multiple levels within the crust[Rooney et al., 2005]. These data indicate that at least somefractionation occurs at deeper levels beneath the SDFZ incomparison to the WFB, consistent with the increased clino-pyroxene fractionation recorded in mass balance models.[32] In order to further refine our working model of low-

pressure fractionation, we employed a coupled linear leastsquares mass balance (Table 3) and thermodynamic modeling(Table 4) approach using MELTS [Ghiorso and Sack, 1995].We chose our most primitive sample (N-21; 10.5% MgO) asa parental composition. Although this sample has low Ce/Pb(11), other incompatible trace element ratios indicate nocrustal contamination (e.g., La/Nb = 0.74) and compatibletrace element abundances suggesting only limited removal ofolivine from a primary mantle melt (165 ppm Ni; 556 ppmCr; 33 ppm Sc). Sensitivity tests indicate that to initiateplagioclase growth (as required from the observed phases andmass balance results) requires a water content for N-21 of!0.2 wt.% and pressure of !1 kbar. At lower MgO contents(<7% MgO), two distinct fractionation trends are evident:Trend A, which is restricted to samples surrounding LakeBesaka, and trend B which defines the majority of sampleswithin the WFB. Variable proportions of fractionating phasesgenerate the alternative evolution pathways for the WFBlavas: Trend A is dominated by clinopyroxene fractionationand trend ‘‘B’’ reflects increased plagioclase and eventuallyFe-Ti oxide growth. To reduce the plagioclase:clinopyroxeneratio and thus replicate Trend A lavas at 1 kbar, a highermagmatic water content (!0.5%) is required. The inferredheterogeneous magmatic water contents likely reflect varia-tions in the parental magmas. A significant result of thismodeling is the requirement for low pressure (!1 kbar) tosatisfy the relative proportions and total mass of fractionatedphases predicted using the mass balance method.[33] Mass balance models of fractionation for lavas from

the SDFZ (Figure 1) highlight an even greater role forclinopyroxene fractionation relative to that inferred for anyof the Trend A or B WFB lavas (Table 2). Plagioclase isabsent from the modeled fractionating assemblage in DebreZeyit but is present in Butajira. The increased role ofclinopyroxene fractionation in these basalts is consistent withSc abundances, which form much tighter arrays with de-creasing MgO contents in comparison to the WFB (Figure 5).The substantial clinopyroxene fractionation may indicatedeeper fractionation conditions and/or increased water con-tent in SDFZ lavas compared to the WFB lavas.[34] Previous thermodynamic studies of Quaternary mag-

matic systems within the MER have attributed the evolutionamong basalts to primarily shallow fractional crystallization[Trua et al., 1999; Peccerillo et al., 2003]. Trua et al.[1999], using a more primitive parental sample than thisstudy (12% MgO), modeled limited fractionation initially at5 kbar followed by extensive crystallization at 2 kbar (42%;clinopyroxene, olivine, plagioclase and titanomagnetite). Itis reasonable that fractionation of quite primitive lavas(>10 wt.% MgO) may occur at levels deeper than thosemodeled in our data set. Peccerillo et al. [2003] extendfractionation to lower MgO contents (0.27%) by !40%fractionation of olivine, clinopyroxene, plagioclase andtitanomagnetite. The dominance of olivine and plagioclasein the fractionating assemblage is consistent with the low

Table 2. Xenolith and Host Lava Components With Temperatureand Pressure Estimatesa

Sample Wo En Fs AcP,

kbarDepth,km Fo Fa Temp

N-20 (C)b 46.45 41.46 10.68 1.41 1.4 5.0N-20 45.57 41.99 11.03 1.41 0.3 1.1N-22 47.85 39.24 11.49 1.41 1.1 4.0N-22 46.64 42.76 9.46 1.14 0.2 0.6N-22 46.30 40.57 11.65 1.48 0.9 3.1N-22 48.34 40.48 9.91 1.26 2.0 6.9N-22 46.82 36.26 13.77 3.15 1.0 3.7N-21 43.87 42.74 11.93 1.46 0.4 1.5N-14 44.64 41.66 12.07 1.63 1.5 5.2N-1 46.35 42.07 10.41 1.18 1.4 5.0N-24 (R)b 45.62 36.89 15.94 1.55 0.1 0.4N-7 52.1 46.6 977N-17 47.4 51.6 1034N-17 64.8 34.7 1171N-16 73.9 25.7 1201N-22 71.0 28.4 970N-22 81.2 18.5 1064N-22 73.9 25.6 1111N-1 77.5 22.1 1080N-1 81.8 17.9 1156N-1 80.5 19.2 1089N-1 78.4 21.2 1100

aPressure calculated using Nimis and Ulmer [1998] with a standard errorof 1.7 kbar. Temperatures calculated using geothermometer (standard erroris ±6!C; analytical uncertainty is generally less than ±35!) of Loucks[1996].

bValues are not systematically distributed (e.g., core and rim) exceptwhere marked core (‘‘C’’) and rim (‘‘R’’).

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Table

3.MassBalance

Fractional

CrystallizationModelingoftheWonjiiFaultBeltBasaltsa

WonjiiBelt

Debre

Zeyit

Butajira

From

N-21

toN-18

Difference

b

From

N-18

to1030

Difference

b

From

1030

toN-17

Difference

b

From

N-18

to1026

Difference

b

From

1026

toN-6

Difference

b

From

DZ-1009

toDZ-1007

Difference

b

From

DZ-1007

toDZ-1004

Difference

b

From

DZ-1004

toDZ-1006

Difference

b

From

BJ-1045

toBJ-1048

Difference

b

From

BJ-1048

toBJ-1047

Difference

b

From

BJ-1047

toBJ-1042

Difference

b

SiO

20.13

$0.12

0.01

$0.22

00.25

$0.52

$0.35

$0.03

$0.12

$0.24

TiO

2$0.07

0.06

$0.05

$0.15

$0.28

$0.01

0.1

0.28

0.31

$0.18

0.06

Al 2O3

$0.11

0.01

$0.01

0.09

$0.03

$0.38

0.18

0.17

0.01

$0.08

0.3

FeO

$0.02

$0.06

0.02

0.07

0.11

$0.11

0.55

0.21

$0.09

0.38

0.13

MnO

00

0$0.01

$0.01

00.02

00

0.01

0MgO

$0.04

0.07

$0.02

0.05

$0.12

$0.06

$0.01

0.05

0.04

$0.1

0.03

CaO

$0.07

0.07

00.14

0.07

$0.11

0.34

0.2

00.1

0.12

Na 2O

0.05

0.36

$0.02

0.27

$0.06

0.29

0.07

0.21

0.07

$0.13

$0.2

K2O

$0.26

0.43

$0.04

0.64

$0.08

0.23

$0.3

0.12

0.18

$0.39

$0.01

P2O5

$0.04

0.17

$0.03

0.06

0.29

$0.12

0.12

0.07

$0.06

0.01

0.09

Fractionated

phases

PL

An73:3.50%

An71:2.70%

An65:8.90%

An71:14.10%

An54:23.40%

An71:5.4%

CPX

En43:8%

En40:6.70%

En38:6.80%

En40:7.70%

En37:14%

En41:7%

En45:3.90%

En 4

5:8.9%

En 4

4:5.6%

En45:8.1%

En34:8.7%

OL

Fo 8

5:7.10%

Fo 8

1:0.80%

Fo 7

4:2.20%

Fo 8

1:5.70%

Fo61:5.30%

Fo85:2%

Fo83:1.50%

Fo 8

2:2.2%

Fo 8

6:2.4%

Fo 8

2:2.6%

Fo80:2.7%

MT

Fe-Ti:5.20%

Residual

liquid

81.80%

88.90%

82.40%

71.70%

52.20%

91.10%

94.60%

88.30%

86%

90%

84.80%

Sr2

0.112

0.375

0.06

0.594

0.2

0.385

0.843

0.38

0.145

0.384

0.237

a Theprimitivesamplesusedhere(N

-21;N-18;1030)

arepresentedbyFurm

anet

al.[2006].Themodel

was

usedto

simulate

discretesteps(e.g.,N-21to

N-18,N-18to

1026)andnotacontinuousfractionation

trend.Thefractionated

phases

wereclinopyroxene(CPX),olivine(O

L),feldspar

(PL),andtitanomagnetite(M

T).

bObserved

minus

calculated,times

theweightfactor.Alloxides

are1except

SiO

2(0.4)andAl 2O3(0.5).

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Table

4.Thermodynam

icModelingofWonjiiFaultBeltBasaltsUsingMELTSa

Trend

‘‘A’’

Trend‘‘B’’

N-21

N-18

1030

N-17

N-18

1026

N-6

Parent

Calc

Measured

Calc

Measured

Calc

MeasuredParent

Calc

Measured

Calc

Measured

Liquidus,!C

1256

1172

1177

1163

1166

1138

1149

1186

1138

1160

1106

1118

Pressure,kbar

11

11

11

11

11

11

Log (10)fO

2$6.7

$7.63

$7.57

$7.73

$7.69

$8.03

-$8.46

$9.03

$8.77

$9.44

-Buffer

QFM+1

QFM+1

QFM+1

QFM+1

QFM+1

QFM+1

QFM+1

QFM

QFM

QFM

QFM

QFM

SiO

247.16

47.35

46.9

46.77

47.27

47.23

46.73

46.95

46.71

47.71

51.29

50.65

TiO

21.91

2.31

2.27

2.48

2.34

2.75

2.72

2.27

2.92

3.18

2.65

3.22

Al 2O3

14.19

15.46

15.85

15.91

16.43

15.28

16.40

15.87

15.44

15.25

13.53

15.10

Fe 2O3

2.54

2.84

2.74

2.9

2.78

3.14

3.02

1.89

2.26

2.13

2.21

2.05

FeO

8.33

9.02

8.54

9.14

9.27

10.46

10.16

9.3

11.05

11.08

10.66

9.82

MnO

0.18

0.23

0.19

0.22

0.2

0.24

0.21

0.19

0.27

0.23

0.39

0.25

MgO

10.51

7.35

7.45

6.83

6.87

6.24

6.28

7.44

5.44

5.48

4.05

4.03

CaO

11.09

10.46

10.93

10.1

10.06

9.29

9.06

10.95

9.32

9.81

8.08

7.94

Na 2O

2.84

3.43

3.23

3.51

3.12

3.38

3.35

3.24

3.96

3.48

4.41

4.39

K2O

0.67

0.85

1.14

1.28

0.78

0.92

0.98

1.14

1.56

0.66

1.06

1.30

P2O5

0.37

0.47

0.51

0.58

0.38

0.46

0.50

0.51

0.71

0.63

1.06

0.65

H2O

0.2*

0.25

0.25*

0.28

0.5*

0.6

0.6*

0.25

0.35

0.35

0.6

0.6*

Fractionated

phases

OL

Fo 8

6:6.53%

Fo83:0.86%

Fo82:2.62%

Fo80:4.55%

Fo 7

4:2.21%

CPX

Di 52,Ce 1

5,He 1

3:9.51%

Di 50,Ce 1

4,He 1

3:5.61%

Di 44,Ce 1

6,He 1

6:2.70%

Di 49,Ce 1

3,He 1

6:10.42%

Di 36,Ce 1

9,He 2

3:11.12%

PL

An75:3.64%

An 7

5:3.43%

An75:10.25%

An73:12.18%

An63:19.60%

MT

Mt 32,Sp42,Usp

58:5.69%

a GhiorsoandSack

[1995].Theprimitivesamplesusedhere(N

-21;N-18;1030)

arepresented

byFurm

anetal.[2006].FeO

:Fe 2O3forthesamplesiscalculatedonthebasisoftherelevantoxygen

buffer

andliquidus

conditions.Nomeasuredwater

contentsareavailable,andallwater

valuesaremodeled.Measuredvalues

forsamplesarenorm

alized

to100%

given

themodeled

water

contentsandFeO

:Fe 2O3.Themodelwas

usedto

simulatediscretesteps(e.g.,N-21to

N-18,N-18to

1026)

andnotacontinuousfractionationtrend.Thestartingmaterialforeach

step

isthemeasuredvalueforthatsampleandnottheendpointoftheprevious

model

step.AQFM

+1oxygen

fugacitydroppingto

QFM

was

used,similar

tothatusedbyTruaetal.[1999].Initialwater

contentsandpressure

weredetermined

bymodelsensitivityanalysis.The

modelfollow

stwobroad

paths

correspondingto

trends‘‘A’’and‘‘B’’in

Figure4.Anasterisk

(*)indicates

that

thesewater

contentsarenotmeasuredandaretaken

from

modeled

values.Fo,fosterite;

Di,diopside;

Ce,

clinoenstatite;

He,

hedenbergite;

An,anorthite;

Mt,magnetite;

Sp,spinel;Usp,ulvospinel.

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modeled pressure of !0.5 kbar for samples from Gedemsa,slightly shallower than modeled in this study as expected forlavas erupted within an existing caldera system.

5.2. Spatial Variations in Crustal Structure

[35] The data presented here allow for comparison offractionation depths across the rift (east-west, between theWFB and SDFZ) and along the length of the rift (northeast-southwest, along the WFB). This three-dimensional frame-work of fractionation depths within the rift has directrelevance to existing geophysical work that has pointed tothe presence of melt within the crust. Analysis of tele-seismic receiver functions provides information on bulkcrustal properties: thickness and Vp/Vs ratio. Stuart et al.[2006] and Dugda et al. [2005] indicate that measured Vp/Vs ratios within the rift are high (%2 in some cases) andpoint to melt within the crust, but this method is unable tolocate this melt precisely. Kendall et al. [2005, 2006] alsosuggest the presence of melt within the crust to explainobservations of seismic anisotropy beneath the MER. Anal-yses of magnetotelluric data have suggested melt within thecrust beneath the SDFZ and WFB at both shallow (<5 km)and mid-crustal depths [Whaler and Hautot, 2006]. Thegeochemical constraints presented here provide direct evi-dence that melt resides in the crust, strengthening existinggeophysical arguments. Moreover, we have shown that thedepth of melt fractionation does not vary along the length ofthe MER, but does differ across it. The consistently shallowfractionation conditions of Quaternary products within theWFB are in contrast to the basalts erupted in the SDFZ thatexhibit fractionation throughout the crust (Figure 11). Thisvariation between the two tectonomagmatic zones mayreflect a more mature magmatic plumbing system withinthe WFB, where magma quickly rises through existingconduits to fractionate at shallow depths. Conversely, inthe SDFZ these conduits are less well defined and melt maytake multiple routes to the surface, on occasion stagnatingas indicated by Al-augite xenoliths and the wide range infractionation depths observed in the SDFZ (Figure 12)[Rooney et al., 2005].[36] Further details of crustal structure in the MER may

be gained by probing the method by which Quaternarybasalts traversed the crust. Within the WFB, spatiallyrelated primitive basalts exhibit substantially heterogeneoustrace element and isotopic compositions [Furman et al.,2006]. These diverse geochemical signatures point to aheterogeneous source region and to a complex magmaticplumbing system, which allows the transport of magmathrough the crust without the homogenization typicallyassociated with residence in larger magma-chambers.Smaller conduits are also consistent with the low volumeof basaltic material erupted in the MER. Where multipledikes/conduits have been ‘‘captured’’ by an existing silicicplumbing system, larger volumes of extrusives may bepossible, e.g., Kone caldera and the fissure eruption atFantale. Keranen et al. [2004] suggested extension in thecentral MER is accommodated by magmatic intrusion intoboth a ductile middle-lower crust and upper brittle crust andcontrolled source reflection/refraction studies [MacKenzie etal., 2005; Maguire et al., 2006] indicate extensive replace-ment of the MER lower crust by high velocity maficmaterial. Mafic intrusions are also consistent with observed

gravity data [Mahatsente et al., 1999; Cornwell et al.,2006]. At deeper levels, Kendall et al. [2005] predict aregion of magma injection in the mantle lithosphere beneatha relatively unstretched, but heavily intruded crust. Similarmethods have also constrained the presence of melt atcrustal levels. Dike emplacement into the crust is consistentwith dike-derived xenoliths found in lavas from the SDFZ[Rooney et al., 2005]. These data indicate that the modifi-cation of the crust beneath these active tectonomagmaticbelts is primarily through dike intrusion associated with themagmatic plumbing system of the rift.[37] We may gain insight into the magmatic plumbing

system from surface observations of eruptive activity andthe relationship of these eruptives to tectonic features. Thebasaltic cinder cones and flows within the WFB and SDFZare associated with the tail cracks of faults/fractures[Chorowicz et al., 1994; Korme et al., 1997, 2004],suggesting a direct link between tectonics and magmatism.Basaltic cinder cones observed in the western rift alsoshare this tectonomagmatic relationship [e.g., Ebinger,1989a, 1989b]. The numerous fractures and faults exposedin the WFB are not deep-seated features and propagatefrom the surface to a maximum estimated depth of only!2 km [Acocella et al., 2003a, 2003b], though seismicityindicates some deeper fault activity (14 km [Keir et al.,2006]). It is therefore plausible that these fractures inter-sect shallow dikes or magma storage chambers, promptingeruption of the variably fractionated contents.

5.3. Synthesis Model of Crustal Structure andImplications for Rifting

[38] The geochemical and geophysical data gathered inthe north central MER support a model of focused mag-matic and tectonic activity during the Quaternary, centeredon the WFB and SDFZ. The spatial distribution of thesefault belts, which are the surface expression of riftingrelated crustal modification, have significant utility indetermining how extension evolves in a rift environment.Figure 1a shows the location of Quaternary cinder conesand associated lava flows within the MER, reflecting thedistribution of Quaternary magmatic and tectonic activity.Shallow seismicity within the rift [Keir et al., 2006] oftencoincides with Quaternary lava flows; epicenters cluster onthe edges of flows and are less common directly beneath theflows (Figure 13), suggesting linkage between the rheolog-ical characteristics of the upper crust and magmatism withinthe rift. There is an abrupt discontinuity in the alignment ofcinder cones around 8!300N, coincident with a change inoverall cinder cone alignment from !N24!E in the south to!N41!E farther north. There is also a change in strike of theMER border faults at 8!300N from N32!E ± 9! in thesouthwest to N47!E ± 12! in the northeast [Acocella andKorme, 2002]. Segmentation of seismic anomalies at 75 kmin the upper mantle presented by Bastow et al. [2005]mirrors this change in cinder cone alignment.[39] Existing lithospheric structures exert control on ex-

tension throughout the EARS [Ebinger et al., 1997; Petitand Ebinger, 2000; Nyblade and Brazier, 2002]; in partic-ular, existing basement structures in the MER have beenreactivated during the Cenozoic [Korme et al., 2004]. East-west trending structures of the Yerer-Tullu Wellel volcano-tectonic lineament (YTVL) are observed on the rift flanks

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[Abebe et al., 1998] and may exert influence in the rift itself[Mazzarini et al., 1999; Bonini et al., 2005]. Bastow et al.[2005] noted a seismic low velocity zone in the uppermostupper mantle beneath the YTVL, suggesting it is a featurethat exists to asthenospheric depths. Significantly, thenorthern limit of the YTVL corresponds to the shift inalignment of the rift border faults (discussed above) and tothe Boru-Toru structural high, interpreted as a transfer zonebetween MER rift segments [Bonini et al., 2005]. The Boru-Toru Structural High thus separates the northern EARS intotwo zones: one to the northeast that is part of the Red SeaRift system and one to the southwest, dominated by theMER (Figure 14). Regional geodynamic models supportthis division of the northern EARS, indicating that rifting inAfar commenced along the Red Sea at !26–29 Ma andpropagated southward only as far as !10!N at that time[Wolfenden et al., 2005]. To the south, rifting of the MERthat began at !25 Ma in Turkana [Hendrie et al., 1994;Morley, 1994] may have propagated northward forming riftborder faults in the northern MER at !11 Ma [Wolfenden etal., 2004]. Interestingly, two low velocity seismic anomaliesare imaged in the upper mantle beneath this region [Bastowet al., 2005], which are distinct at 75 km depth and locatedon either side of the structural high, but merging withincreasing depth. Within the upper crust, the structural highalso marks the site of a change in orientation of high

velocity anomalies, interpreted as cooled mafic intrusionsthat underlie the WFB and SDFZ [Keranen et al., 2004].[40] Figure 14a outlines a synthesis model for the north

central MER at !10 Ma; the northern EARS is divided intotwo zones associated with the northward propagating MERand the southward propagating Red Sea Rift. While somemodels suggest the influence of the Gulf of Aden Riftsystem [Chernet et al., 1998], recent work has highlightedthe dominance of the Red Sea Rift [Wolfenden et al., 2004,2005]. Although we do not completely rule out the possi-bility that Gulf of Aden structures extend into this region,geophysical evidence does not support such a model.Regional SKS shear wave splitting measurements have fastpolarization directions that are oriented NNE-SSW, almostperpendicular to the inferred ESE-WSW Gulf of Aden trendat 9!N [Gashawbeza et al., 2004; Kendall et al., 2005,2006]. Furthermore, seismicity in the region shows adominance of Red Sea trends with little activity on theeastern rift margin north of 9!N [Keir et al., 2006]. A modelof current crustal structure is presented in Figure 14b; thetwo zones apparent at !10 Ma are now connected acrossthe Boru-Toru Structural High. We acknowledge the crustalvariations along the length of MER in both models andgroup our observations into two regions, separated by atransfer zone at 8!300N (Figures 14a and 14b).[41] Northeast of the transfer zone the distinct rift valley

graben morphology that is evident in the southwest givesway to the less well-defined Afar depression and a singlerift-centered zone of magmatism dominates (Figures 14aand 14b). Geophysical surveys confirm a south-to-norththinning of the crust from 40 km to 26 km, coincident with

Figure 11. Plot of latitude versus calculated clinopyroxenephenocryst equilibration depths based on Nimis and Ulmer[1998]. Pressure is converted to depth assuming 1 kbar =3.5 km. Source data are from Table 2 and Rooney et al.[2005]. Phenocryst equilibration depths in the Silti-DebreZeyit Fault Zone extend throughout the crust, while thosefrom the Wonjii Fault Belt are restricted to the upper crust.

Figure 12. A cartoon of crustal structure beneath the Silti-Debre Zeyit Fault Zone (SDFZ) and Wonjii Fault Belt(WFB). This representation suggests that the magmaticplumbing system beneath the WFB is more developed thanthat beneath the SDFZ. This results in magmas within theWFB rising toward the surface more rapidly, fractionatingclose to the surface. Within the SDFZ the magmas mayfractionate at various crustal depths prior to eruption.

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the transfer zone [Maguire et al., 2006], consistent with theinfluence of seafloor spreading propagating southward fromAfar. Crustal modification within this region is dominatedby basaltic and silicic activity within the WFB (the SDFZdoes not extend north of the transfer zone). The fraction-ation data for the WFB presented here indicate a well-developed magmatic system that allows magma to rise toshallow levels before fractionating in predominantly smallmagma bodies (dikes) and some larger magma chambers(e.g., nested calderas). Within this crustally thinned northern

region, the WFB may be a precursor to a seafloor spreadingcenter [e.g., Ebinger and Casey, 2001; Keranen et al., 2004;Casey et al., 2006].[42] Southwest of the transfer zone, the distribution of

recent cinder cones defines two parallel zones of magma-tism corresponding to the SDFZ in the west and the WFB inthe east, neither occupying the central portion of the RiftValley (Figure 14b). Magmatism in the SDFZ extends fromthe embayment of the rift near the Boru-Toru StructuralHigh in the north toward Butajira in the south, eventually

Figure 13. Distribution of seismicity overlain on Quaternary basaltic magmatism in the MER. TheWonjii Fault Belt (WFB) is shaded in blue, while the Silti-Debre Zeyit Fault Zone (SDFZ) is shaded ingreen [Berhe and Wondm-Agennehu, 1978; Benvenuti et al., 2002; Abebe et al., 2005]. Seismicity is fromKeir et al. [2006]. The area north of 9!N has not been mapped, but fieldwork there during 2003 indicatedlittle basaltic magmatism between Fantale and Dofan (see Figure 1). The diagram illustrates a strikinganticorrelation between Quaternary basaltic magmatism and modern seismicity within the rift. Within theWFB, seismicity is concentrated between Fantale and Dofan (1), at the Ankober Border Fault (2), and atthe very southern end of the fault belt (3). Less intense seismicity is also distributed throughout the WFBbut is concentrated at the edge of basaltic flows (white arrows). At the northern end of the SDFZ,seismicity is also observed (dashed oval). The seismicity to the west of the SDFZ at 9!N may be relatedto the YTVL.

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occupying a central zone of magmatism near Lake Shala,merging with the WFB (Figure 13). Magmatism in the WFBdecreases substantially south of the transfer zone ceasingnear Lake Ziway as it merges with the SDFZ. This southerntermination of the WFB before it merges with the SDFZ isalso the site of active extension; a fissuring event recentlyoccurred near Lake Ziway, generating subsidence andrupturing the main Addis Ababa to Awassa road (G. Yirgu,personal communication, 2006]. The crustal thickness(!40 km [Maguire et al., 2006]) in the southwest is greaterthan in the northeast (26 km). Evidence of melt andsolidified intrusions in the crust to depths of !30 km areobserved beneath both belts [e.g., Keranen et al., 2004;Whaler and Hautot, 2006], indicating that magmatic mod-ification of the crustal structure is concentrated on the WFBand SDFZ. These two belts are not equivalent. Fractionationwithin the SDFZ occurs at various depths within the crustand is followed by rapid eruption of the overall lessfractionated products [Rooney et al., 2005]. Within theWFB, fractionation is predominantly a shallow processand generates a greater range in erupted compositions.The differences between these two zones of focused Qua-

ternary tectonomagmatic activity point to heterogeneitywithin the rift-based magmatic plumbing system. Geophys-ical evidence also points to variation between the WFB andSDFZ, specifically magnetotelluric results indicate a deeper(25–30 km) and more pronounced zone of melt beneath theSDFZ in comparison to shallower melt beneath the WFB[Whaler and Hautot, 2006]. Despite suggested models thatthe majority of extension is accommodated within the WFB[Ebinger and Casey, 2001; Casey et al., 2006], GPSmeasurements within or across the rift have been unableto locate precisely where extension is accommodated[Bilham et al., 1999; Pan et al., 2002; Bendick et al.,2006; Pizzi et al., 2006]. Interestingly, no significantvariation in fractionation depth is observed along thelength of the WFB from north to south regardless of theobserved variation in crustal thickness.[43] It is clear that distinct differences in crustal structure

exist between the two regions described above. The dual-focus magmatism near the rift border faults and the minimalcrustal thinning to the southwest passes northeastward into aregion with much thinner crust and a single centralized beltof magmatism away from the rift border faults. The com-

Figure 14. Cartoon of the north central Ethiopian Rift centered on the Boru-Toru Structural High(BTSH). The modern crustal thickness and velocity anomalies are based on geophysical data discussed inthe text. YTVL is the Yerer-Tullu Wellel volcanotectonic lineament [Abebe et al., 1998]. Large volcanoesare silicic centers (e.g., Fantale, Ziqualla); smaller volcanoes are basaltic cinder cones. (a) The !10 Maconfiguration of the Ethiopian Rift. This configuration represents the rift prior to the breach of the BTSHby the Wonjii Fault Belt. The propagation direction of the Main Ethiopian Rift (MER) and Red Sea Riftare based on Wolfenden et al. [2004, 2005] and Bonini et al. [2005]. (b) Present-day structure of theMER. The Boru-Toru Structural High (BTSH) is breached by the Wonjii Fault Belt. To the south of theBTSH the Silti-Debre Zeyit Fault Zone and Wonjii Fault Belt dominate Quaternary magmatism, but tothe north, only the Wonjii Fault Belt is present. This model suggests crustal modification along thelengths of the Wonjii Fault Belt and Silti-Debre Zeyit Fault Zone.

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bined geochemical, geophysical and geodynamic data set allpoint to a model whereby the crustal structure to thenortheast of the transfer zone is dominated by incipientseafloor spreading processes while to the southwest, conti-nental rifting still controls crustal modification processes.This division into two zones separated by a transfer zone iscomplicated by the WFB, which passes through the transferzone. We suggest that the WFB represents the southwardpropagation of the Red Sea Rift system, while the SDFZrepresents the northward propagation of the MER. Thisinterpretation defines the Quaternary zones of tectonomag-matic activity as overlapping spreading centers, generatedby the merger of the Red Sea Rift and MER. This hypoth-esis is supported by seismic data [Keir et al., 2006] thatshow ‘‘earthquake swarms’’ along the Ankober BorderFault (linkage between the Red Sea Rift and MER) andwithin the WFB. In particular, at the southern limit ofmagmatism in the WFB a distinct ‘‘earthquake swarm’’has been noted (Figure 13), coincident with the recentfissuring event at Lake Ziway. This seismically activesouthern tip of the WFB also displays evidence of crustalcontamination in WFB basalts [Trua et al., 1999]. Theincreased seismic activity, in tandem with the poorly devel-oped magmatic plumbing system and evidence of crustalassimilation is indicative of the southward propagating tipof the WFB. Similarly, the enhanced seismic activity andpoorly developed young explosive magmatism at the north-ern limit of the SDFZ at Chefe Donsa, suggests thenorthward propagation of the SDFZ. These two zones oftectonic and magmatic activity thus record more thansimply the transition between rifting and seafloor spreading;they are the result of the recent linkage between the twodifferent structural and tectonic domains defined by thesouthward propagating Red Sea Rift and northward propa-gating MER/EARS.

6. Summary and Conclusions

[44] The combined geochemical and geophysical resultsof the EAGLE project point to two significant advances inour understanding of continental extension and rifting inEast Africa: Extension is accommodated largely throughdike intrusion in the lithosphere and this extension isfocused within regions of Quaternary magmatic activity inthe Main Ethiopian Rift (Wonjii Fault Belt and Silti-DebreZeyit Fault Zone).[45] The geochemical results of the EAGLE project

significantly enhance existing geophysical images of mag-matic intrusion. Coupled mass balance and thermodynamicmodeling indicate that WFB lavas require shallow fraction-ation conditions, indicative of a well-developed magmaticplumbing system. Conversely, within the SDFZ, melt ispresent throughout the crust and magmatic composition isdominated by fractionation at deeper levels. The preserva-tion of source-derived geochemical heterogeneities [Rooneyet al., 2005; Furman et al., 2006; this work] and xenolithevidence [Rooney et al., 2005], points to a complex mag-matic plumbing system dominated by pervasive dikingbeneath the WFB and SDFZ. There is no clear evidenceof pan-African crustal assimilation in these lavas (exceptinglavas at the southern termination of the WFB), howeveranomalous major and trace element ratios indicate the

occasional assimilation of magmatic fractionation products.These results are consistent with abundant geophysical datasuggesting the widespread replacement of the continentalcrust by magmatic intrusion.[46] The geochemical results of Quaternary magmas,

when interpreted within the existing geophysical frame-work, preclude a single axis of extension model for theMER. The data indicate differing magmatic histories andorigins for the WFB and SDFZ lavas that are best explainedby a model where two rifts join: the Southern Red Sea Rift,in which incipient seafloor spreading may be underway haspropagated south and joined with the northward propagat-ing continental East African Rift System in the MainEthiopian Rift.

[47] Acknowledgments. This research was supported by NationalScience Foundation grant EAR 0207764 (T. Furman) as part of the cross-disciplinary Ethiopia Afar Geoscientific Lithosphere Experiment (EAGLE).A George H. Deike Jr. grant to T. Furman provided additional support foranalytical work. T.R. thanks Kassahun Ejeta, guide and driver, and RoelandDoust, field assistant during fieldwork in Ethiopia. Margaret Nitz and LeighPatterson were helpful in sample preparation. We thank Julie Bryce forcontributing to an earlier version of the manuscript and Clifton Rooney forhelp in drafting a figure. We are grateful to Mark Angelone for help withthe Electron Microprobe at Penn State and Emily Klein, Gary Dwyer, andMark Rudnicki of Duke University for performing DCP and ICP-MSanalysis. Thanks also to Andrew Nyblade, Richard Parizek, DerekElsworth, David Eggler, Barry Hanan, Peter LaFemina, and Wendy Nelsonfor thoughtful comments on earlier manuscript versions. We thank GidayWoldeGabriel and Roger Buck for careful peer reviews of the manuscriptand Peter Cawood for editorial handling. Finally, we would like to thank theEAGLE group and in particular Cindy Ebinger for the many stimulatingdiscussions that improved this manuscript.

ReferencesAbebe, T., F. Mazzarini, F. Innocenti, and P. Manetti (1998), The Yerer-Tullu Wellel volcanotectonic lineament: A transtensional structure incentral Ethiopia and the associated magmatic activity, J. Afr. Earth Sci.,26(1), 135–150.

Abebe, T., P. Manetti, M. Bonini, G. Corti, F. Innocenti, F. Mazzarini, andZ. Pecksay (2005), Geological map (scale 1:200,000) of the NorthernMain Ethiopian Rift and its implications for the volcano-tectonic evolu-tion of the rift, Geol. Soc. of Am., Boulder, Colo.

Acocella, V., and T. Korme (2002), Holocene extension direction along theMain Ethiopian Rift, East Africa, Terra Nova, 14(3), 191–197.

Acocella, V., T. Korme, and F. Salvini (2003a), Formation of normal faultsalong the axial zone of the Ethiopian Rift, J. Struct. Geol., 25(4), 503–513.

Acocella, V., T. Korme, F. Salvini, and R. Funiciello (2003b), Ellipticcalderas in the Ethiopian Rift: control of pre-existing structures, J. Vol-canol. Geotherm. Res., 119(1–4), 189–203.

Ayalew, D. (2000), Crustal control in the genesis of Plio-Quaternary bimo-dal magmatism of the Main Ethiopian Rift (MER): Geochemical andisotopic (Sr, Nd, Pb) evidences: Discussion [modified], Chem. Geol.,168(1–2), 1–3.

Ayalew, D., G. Yirgu, and R. Pik (1999), Geochemical and isotopic (Sr, Ndand Pb) characteristics of volcanic rocks from southwestern Ethiopia,J. Afr. Earth Sci., 29(2), 381–391.

Bastow, I. D., G. W. Stuart, J. M. Kendall, and C. J. Ebinger (2005), Upper-mantle seismic structure in a region of incipient continental breakup:northern Ethiopian rift, Geophys. J. Int., 162(2), 479–493.

Bedini, R. M., J. L. Bodinier, J. M. Dautria, and L. Morten (1997), Evolu-tion of LILE-enriched small melt fractions in the lithospheric mantle:A case study from the East African Rift, Earth Planet. Sci. Lett.,153(1–2), 67–83.

Bendick, R., S. McClusky, R. Bilham, L. Asfaw, and S. L. Klemperer(2006), Distributed Nubia-Somalia relative motion and dike intrusion inthe Main Ethiopian Rift, Geophys. J. Int., 165, 303–310.

Benoit, M. H., A. A. Nyblade, J. C. VanDecar, and H. Gurrola (2003),Upper mantle P wave velocity structure and transition zone thicknessbeneath the Arabian Shield, Geophys. Res. Lett., 30(10), 1531,doi:10.1029/2002GL016436.

Benvenuti, M., et al. (2002), The Ziway-Shala lake basin (main Ethiopianrift, Ethiopia): a revision of basin evolution with special reference to theLate Quaternary, J. Afr. Earth Sci., 35(2), 247–269.

B10201 ROONEY ET AL.: LITHOSPHERE IN THE ETHIOPIAN RIFT

19 of 21

B10201

Berhe, S. M., and B. Wondm-Agennehu (1978), Nazret, Ethiopian MappingAgency, Addis Ababa, Ethiopia.

Bilham, R., R. Bendick, K. Larson, P. Mohr, J. Braun, S. Tesfaye, andL. Asfaw (1999), Secular and tidal strain across the main Ethiopian rift,Geophys. Res. Lett., 26(18), 2789–2792.

Boccaletti, M., A. Getaneh, R. Mazzuoli, L. Tortorici, and T. Trua (1995),Chemical variations in a bimodal magma system: The Plio-Quaternaryvolcanism in the Dera Nazret area (Main Ethiopian Rift, Ethiopia), Afr.Geosci. Rev., 2(1), 37–60.

Boccaletti, M., R. Mazzuoli, M. Bonini, T. Trua, and B. Abebe (1999), Plio-Quaternary volcanotectonic activity in the northern sector of the MainEthiopian Rift: Relationships with oblique rifting, J. Afr. Earth Sci.,29(4), 679–698.

Bonini, M., G. Corti, F. Innocenti, P. Manetti, F. Mazzarini, T. Abebe, andZ. Pecskay (2005), Evolution of the Main Ethiopian Rift in the frame ofAfar and Kenya rifts propagation, Tectonics, 24, TC1007, doi:10.1029/2004TC001680.

Boynton, W. V. (1984), Cosmochemistry of the rare earth elements: Meteor-ite studies, in Rare Earth Element Geochemistry, edited by P. Henderson,pp. 63–114, Elsevier, New York.

Buck, W. R. (2004), Consequences of asthenospheric variability on con-tinental rifting, in Rheology and Deformation of the Lithosphere atContinental Margins, edited by G. Karner et al., pp. 1–30, ColumbiaUniv. Press, New York.

Burke, K., and J. F. Dewey (1973), Plume-generated triple junctions: Keyindicators in applying plate tectonics to old rocks, J. Geol., 81, 406–433.

Casey, M., C. Ebinger, D. Keir, R. Gloaguen, and F. Mohamed (2006), Strainaccommodation in transitional rifts: Extension by magma intrusion andfaulting in Ethiopian rift magmatic segments, in The Afar Volcanic Pro-vince Within the East African Rift System, edited by G. Yirgu, C. Ebinger,and P. Maguire, Geol. Soc. Spec. Publ., 259, 143–164.

Chernet, T., and W. K. Hart (1999), Petrology and geochemistry of volcan-ism in the northern Main Ethiopian Rift: Southern Afar transition region,Acta Vulcanol., 11(1), 21–41.

Chernet, T., W. K. Hart, J. L. Aronson, and R. C. Walter (1998), New ageconstraints on the timing of volcanism and tectonism in the northernMain Ethiopian Rift-southern Afar transition zone (Ethiopia), J. Volcanol.Geotherm. Res., 80(3–4), 267–280.

Chorowicz, J., B. Collet, F. F. Bonavia, and T. Korme (1994), Northwest tonorth-northwest extension direction in the Ethiopian Rift deduced fromthe orientation of extension structures and fault-slip analysis, Geol. Soc.Am. Bull., 106(12), 1560–1570.

Cornwell, D. G., G. D. Mackenzie, R. W. England, P. K. H. Maguire, L. M.Asfaw, and B. Oluma (2006), Northern Main Ethiopian Rift crustal struc-ture from new high-precision gravity data, in The Afar Volcanic ProvinceWithin the East African Rift System, edited by G. Yirgu, C. Ebinger, andP. Maguire, Geol. Soc. Spec. Publ., 259, 307–322.

Courtillot, V., C. Jaupart, I. Manighetti, P. Tapponnier, and J. Besse (1999),On causal links between flood basalts and continental breakup, EarthPlanet. Sci. Lett., 166(3–4), 177–195.

Deniel, C., P. Vidal, C. Coulon, and P. J. Vellutini (1994), Temporal evolu-tion of mantle sources during continental rifting: The volcanism ofDjibouti (Afar), J. Geophys. Res., 99(B2), 2853–2869.

Dugda, M. T., A. A. Nyblade, J. Julia, C. A. Langston, C. J. Ammon, andS. Simiyu (2005), Crustal structure in Ethiopia and Kenya from receiverfunction analysis: Implications for rift development in eastern Africa,J. Geophys. Res., 110, B01303, doi:10.1029/2004JB003065.

Ebinger, C. J. (1989a), Geometric and kinematic development of borderfaults and accommodation zones, Kivu-Rusizi rift, Africa, Tectonics,8(1), 117–133.

Ebinger, C. J. (1989b), Tectonic development of the western branch of theEast-African rift system, Geol. Soc. Am. Bull., 101(7), 885–903.

Ebinger, C. J., and M. Casey (2001), Continental breakup in magmaticprovinces: An Ethiopian example, Geology, 29(6), 527–530.

Ebinger, C. J., T. D. Bechtel, D. W. Forsyth, and C. O. Bowin (1989),Effective elastic plate thickness beneath the East-African and Afarplateaus and dynamic compensation of the uplifts, J. Geophys. Res.,94(B3), 2883–2901.

Ebinger, C., Y. P. Djomani, E. Mbede, A. Foster, and J. B. Dawson (1997),Rifting Archaean lithosphere: The Eyasi-Manyara-Natron rifts, EastAfrica, J. Geol. Soc., 154, 947–960.

Farnetani, C. G., M. A. Richards, and M. S. Ghiorso (1996), Petrologicalmodels of magma evolution and deep crustal structure beneath hotspotsand flood basalt provinces, Earth Planet. Sci. Lett., 143(1–4), 81–94.

Furman, T., J. G. Bryce, J. Karson, and A. Iotti (2004), East African RiftSystem (EARS) plume structure: Insights from Quaternary mafic lavas ofTurkana, Kenya, J. Petrol., 45(5), 1069–1088.

Furman, T., J. G. Bryce, T. Rooney, B. B. Hanan, G. Yirgu, and D. Ayalew(2006), Heads and tails: 30 million years of the Afar plume, in The Afar

Volcanic Province Within the East African Rift System, edited by G. Yirgu,C. Ebinger, and P. Maguire, Geol. Soc. Spec. Publ., 259, 95–120.

Gashawbeza, E. M., S. L. Klemperer, A. A. Nyblade, K. T. Walker, andK. M. Keranen (2004), Shear-wave splitting in Ethiopia: Precambrianmantle anisotropy locally modified by Neogene rifting, Geophys. Res.Lett., 31, L18602, doi:10.1029/2004GL020471.

Gasparon, M., F. Innocenti, P. Manetti, A. Peccerillo, and A. Tsegaye(1993), Genesis of the Pliocene to Recent bimodal mafic-felsic volcanismin the Debre Zeyt area, central Ethiopia: Volcanological and geochemicalconstraints, J. Afr. Earth Sci., 17(2), 145–165.

Ghiorso, M. S., and R. O. Sack (1995), Chemical mass-transfer in mag-matic processes. 4. A revised and internally consistent thermodynamicmodel for the interpolation and extrapolation of liquid-solid equilibria inmagmatic systems at elevated-temperatures and pressures, Contrib.Mineral. Petrol., 119(2–3), 197–212.

Gregory, J. (1896), The Great Rift Valley, 405 pp., John Murray, London.Hart, W. K., G. Woldegabriel, R. C. Walter, and S. A. Mertzman (1989),Basaltic volcanism in Ethiopia: Constraints on continental rifting andmantle interactions, J. Geophys. Res., 94(B6), 7731–7748.

Hendrie, D. B., N. J. Kusznir, C. K. Morley, and C. J. Ebinger (1994),Cenozoic extension in Northern Kenya—A quantitative model of riftbasin development in the Turkana region, Tectonophysics, 236(1 –4),409–438.

Hofmann, A. W., K. P. Jochum, M. Seufert, and W. M. White (1986), Nband Pb in oceanic basalts — New constraints on mantle evolution, EarthPlanet. Sci. Lett., 79(1–2), 33–45.

Hofmann, C., V. Courtillot, G. Feraud, P. Rouchett, G. Yirgu, E. Ketefo,and R. Pik (1997), Timing of the Ethiopian flood basalt event and im-plications for plume birth and global change, Nature, 389(6653), 838–841.

Holbrook, W., and P. Kelemen (1993), Large igneous province on the USAtlantic margin and implications for magmatism during breakup, Nature,364, 433–436.

Hopper, J., and W. Buck (1993), The initiation of rifting at constant tectonicforce: The role of diffusion creep, J. Geophys. Res., 98, 16,213–16,221.

Huismans, R. S., Y. Y. Podladchikov, and S. Cloetingh (2001), Transitionfrom passive to active rifting: Relative importance of asthenosphericdoming and passive extension of the lithosphere, J. Geophys. Res.,106(B6), 11,271–11,291.

Kebede, T., C. Koeberl, and F. Koller (1999), Geology, geochemistry andpetrogenesis of intrusive rocks of the Wallagga area, western Ethiopia,J. Afr. Earth Sci., 29(4), 715–734.

Keir, D., C. J. Ebinger, G. W. Stuart, E. Daly, and A. Ayele (2006), Strainaccommodation by magmatism and faulting as rifting proceeds to break-up: Seismicity of the northern Ethiopian rift, J. Geophys. Res., 111,B05314, doi:10.1029/2005JB003748.

Kendall, J.-M., G. W. Stuart, C. Ebinger, I. D. Bastow, and D. Keir (2005),Magma-assisted rifting in Ethiopia, Nature, 433(7022), 146–148.

Kendall, J.-M., S. Pilidou, D. Keir, I. D. Bastow, G. Stuart, and A. Ayele(2006), Mantle upwellings, melt migration and the rifting of Africa: In-sights from seismic anisotropy, in The Afar Volcanic Province Within theEast African Rift System, edited by G. Yirgu, C. Ebinger, and P. Maguire,Geol. Soc. Spec. Publ., 259, 55–72.

Keranen, K., S. L. Klemper, R. Gloaguen, and E. W. Grp (2004), Three-dimensional seismic imaging of a protoridge axis in the Main Ethiopianrift, Geology, 32(11), 949–952.

Kieffer, B., et al. (2004), Flood and shield basalts from Ethiopia: Magmasfrom the African superswell, J. Petrol., 45(4), 793–834.

Korme, T., J. Chorowicz, B. Collet, and F. F. Bonavia (1997), Volcanicvents rooted on extension fractures and their geodynamic implications inthe Ethiopian Rift, J. Volcanol. Geotherm. Res., 79(3–4), 205–222.

Korme, T., V. Acocella, and B. Abebe (2004), The role of pre-existingstructures in the origin, propagation and architecture of faults in the mainEthiopian rift, Gondwana Res., 7(2), 467–479.

Kusznir, N. and R. Park (1987), The extensional strength of the continentallithosphere: Its dependence on geothermal gradient, and crustal composi-tion and thickness, in Continental Extensional Tectonics, edited by P. L.Hancock, Geol, Soc. Spec. Publ., 28, 35–52.

Le Bas, M. J., R. W. Le Maitre, A. Streckeisen, and B. Zanettin (1986), Achemical classification of volcanic rocks based on total Alkali-Silicadiagram, J. Petrol., 27(3), 745–750.

Lister, G., M. Etheridge, and P. Symonds (1986), Detachment faulting andthe evolution of passive continental margins, Geology, 14, 246–250.

Loucks, R. R. (1996), A precise olivine-augite Mg-Fe-exchange geotherm-ometer, Contrib. Mineral. Petrol., 125(2–3), 140–150.

MacKenzie, G., H. Thybo, and P. Maguire (2005), Crustal velocity structureacross the Main Ethiopian Rift: Results from 2-dimensional wide-angleseismic modelling, Geophys. J. Int., 162, 994–1006.

Maguire, P., et al. (2003), Geophysics project in Ethiopia studies continen-tal breakup, Eos Trans. AGU, 84(35), 342–343.

B10201 ROONEY ET AL.: LITHOSPHERE IN THE ETHIOPIAN RIFT

20 of 21

B10201

Maguire, P., et al. (2006), Crustal structure of the Northern Main EthiopianRift from the EAGLE controlled source survey: A snapshot of incipientlithospheric break-up, in The Afar Volcanic Province Within the EastAfrican Rift System, edited by G. Yirgu, C. Ebinger, and P. Maguire,Geol. Soc. Spec. Publ., 259, 269–292.

Mahatsente, R., G. Jentzsch, and T. Jahr (1999), Crustal structure ofthe Main Ethiopian Rift from gravity data: 3-dimensional modeling,Tectonophysics, 313(4), 363–382.

Mazzarini, F., T. Abebe, F. Innocenti, P. Manetti, and M. T. Pareschi (1999),Geology of the Debre Zeyt area (Ethiopia) (with a geologic map at scale1:100,000), Acta Vulcanol., 11(1), 131–141.

McKenzie, D. P. (1978), Some remarks on the development of sedimentarybasins, Earth Planet. Sci. Lett., 40, 25–32.

Menzies, M. A., S. L. Klemperer, C. J. Ebinger, and J. Baker (2002),Characteristics of volcanic rifted margins, in Volcanic Rifted Margins,edited by M. A. Menzies et al., Spec. Pap. Geol. Soc. Am., 362, 1–14.

Mohr, P. A. (1962), The Ethiopian rift system, Bull. Geophys. Observ., 3(1),33–62.

Mohr, P. A. (1967), Major volcano-tectonic lineament in the Ethiopian riftsystem, Nature, 213(5077), 664–665.

Mohr, P. (1983), Ethiopian flood basalt province, Nature, 303(5918), 577–584.

Mohr, P. (1987), Structural style of continental rifting in Ethiopia: Reversedecollements, Eos Trans. AGU, 68, 721, 729–730.

Mohr, P. and B. Zanettin (1988), The Ethiopian flood basalt province, inContinental Flood Basalts, edited by J. D. MacDougall, pp. 63–110,Kluwer Acad., Dordrecht, Netherlands.

Morgan, J. P. (1971), Convection plumes in the lower mantle, Nature, 230,42–43.

Morley, C. K. (1994), Interaction of deep and shallow processes in theevolution of the Kenya rift, Tectonophysics, 236(1–4), 81–91.

Nimis, P., and P. Ulmer (1998), Clinopyroxene geobarometry of magmaticrocks Part 1: An expanded structural geobarometer for anhydrous andhydrous, basic and ultrabasic systems, Contrib. Mineral. Petrol., 133(1–2), 122–135.

Nyblade, A. A., and R. A. Brazier (2002), Precambrian lithospheric con-trols on the development of the East African rift system, Geology, 30(8),755–758.

Pan, M., L. E. Sjoeberg, L. M. Asfaw, E. Asenjo, A. Alemu, andA. Hunegnaw (2002), An analysis of the Ethiopian Rift Valley GPScampaigns in 1994 and 1999, J. Geodyn., 33, 333–343.

Peccerillo, A., M. R. Barberio, G. Yirgu, D. Ayalew, M. Barbieri, and T. W.Wu (2003), Relationships between mafic and peralkaline silicic magma-tism in continental rift settings: A petrological, geochemical and isotopicstudy of the Gedemsa volcano, central Ethiopian rift, J. Petrol., 44(11),2003–2032.

Petit, C., and C. Ebinger (2000), Flexure and mechanical behavior of cra-tonic lithosphere: Gravity models of the East African and Baikal rifts,J. Geophys. Res., 105(B8), 19,151–19,162.

Pik, R., C. Deniel, C. Coulon, G. Yirgu, C. Hofmann, and D. Ayalew(1998), The northwestern Ethiopian Plateau flood basalts: Classificationand spatial distribution of magma types, J. Volcanol. Geotherm. Res.,81(1–2), 91–111.

Pik, R., C. Deniel, C. Coulon, G. Yirgu, and B. Marty (1999), Isotopic andtrace element signatures of Ethiopian flood basalts: Evidence for plume-lithosphere interactions,Geochim. Cosmochim. Acta, 63(15), 2263–2279.

Pizzi, A., M. Coltorti, B. Abebe, L. Disperati, G. Sacchi, and R. Salvini(2006), The Wonjii Fault Belt (Main Ethiopian Rift): structural and geo-morphological constraints and GPS monitoring, in The Afar VolcanicProvince Within the East African Rift System, edited by G. Yirgu, C. J.Ebinger, and P. K. H. Maguire, Geol. Soc. Spec. Publ., 259, 191–207.

Rooney, T. (2006), Continental rifting in central Ethiopia: Geochemical andisotopic constraints from lavas and xenoliths, Ph.D. dissertation thesis,Penn. State Univ., Univ. Park.

Rooney, T., T. Furman, G. Yirgu, and D. Ayalew (2005), Structure of theEthiopian lithosphere: Xenolith evidence in the Main Ethiopian Rift,Geochim. Cosmochim. Acta, 69(15), 3889–3910.

Sengor, A. M. C., and K. Burke (1978), Relative timing of rifting andvolcanism on Earth and its tectonic implications, Geophys. Res. Lett.,5(6), 419–421.

Sifeta, K., B. P. Roser, and J. I. Kimura (2005), Geochemistry, provenance,and tectonic setting of Neoproterozoic metavolcanic and metasedimen-tary units, Werri area, Northern Ethiopia, J. Afr. Earth Sci., 41(3), 212–234.

Stuart, G., I. D. Bastow, and C. Ebinger (2006), Crustal structure ofthe northern Ethiopian Rift from receiver function studies, in The AfarVolcanic Province Within the East African Rift System, edited by G. Yirgu,C. J. Ebinger, and P. K. H. Maguire, Geol. Soc. Spec. Publ., 259, 253–267.

Tadesse, G., and A. Allen (2005), Geology and geochemistry of the Neo-proterozoic Tuludimtu Ophiolite suite, western Ethiopia, J. Afr. EarthSci., 41(3), 192–211.

Trua, T., C. Deniel, and R. Mazzuoli (1999), Crustal control in the genesisof Plio-Quaternary bimodal magmatism of the Main Ethiopian Rift(MER): Geochemical and isotopic (Sr, Nd, Pb) evidence, Chem. Geol.,155(3–4), 201–231.

Turcotte, D. L., and S. H. Emerman (1983), Mechanisms of active andpassive rifting, Tectonophysics, 94(1–4), 39–50.

Wernicke, B. P. (1985), Uniform-sense normal simple shear of the conti-nental lithosphere, Can. J. Earth Sci., 22(1), 108–125.

Whaler, K. and S. Hautot (2006), Magnetotelluric studies of the northernEthiopian rift, in The Afar Volcanic Province Within the East African RiftSystem, edited by G. Yirgu, C. Ebinger, and P. Maguire, Geol. Soc. Spec.Publ., 259, 293–306.

WoldeGabriel, G., J. L. Aronson, and R. C. Walter (1990), Geology, geo-chronology, and rift basin development in the central sector of the MainEthiopia Rift, Geol. Soc. Am. Bull., 102(4), 439–458.

Wolfenden, E., C. Ebinger, G. Yirgu, A. Deino, and D. Ayalew (2004),Evolution of the northern Main Ethiopian rift: Birth of a triple junction,Earth Planet. Sci. Lett., 224(1–2), 213–228.

Wolfenden, E., C. Ebinger, G. Yirgu, P. R. Renne, and S. P. Kelley (2005),Evolution of a volcanic rifted margin: Southern Red Sea, Ethiopia, Geol.Soc. Am. Bull., 117(7–8), 846–864.

Wright, T. J., C. Ebinger, J. Biggs, A. Ayele, G. Yirgu, D. Keir, and A. Stork(2006), Magma-maintained rift segmentation at continental rupture in the2005 Afar dyking episode, Nature, 442, 291–294.

Yihunie, T., M. Adachi, and K. Yamamoto (2006), Geochemistry of theNeoproterozoic metabasic rocks from the Negele area, southern Ethiopia:Tectonomagmatic implications, J. Afr. Earth Sci., 44(3), 255–269.

Zeyen, H., F. Volker, V. Wehrle, K. Fuchs, S. V. Sobolev, and R. Altherr(1997), Styles of continental rifting: Crust-mantle detachment and mantleplumes, Tectonophysics, 278(1–4), 329–352.

Ziegler, P. A., and S. Cloetingh (2004), Dynamic processes controllingevolution of rifted basins, Earth Sci. Rev., 64(1–2), 1–50.

$$$$$$$$$$$$$$$$$$$$$$$D. Ayalew, Dessie/Kombolcha University, P.O. Box 1145, Dessie,

Ethiopia.I. Bastow, Department of Earth Sciences, University of Bristol, Bristol

BS8 1RJ, UK.T. Furman, Department of Geosciences, Pennsylvania State University,

University Park, PA 16802, USA.T. Rooney, Department of Geological Sciences, Michigan State

University, East Lansing, MI 48824, USA. (rooneyt@msu.edu)G. Yirgu, Department of Earth Sciences, Addis Ababa University, P.O.

Box 1176, Addis Ababa, Ethiopia.

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