6.1. Upper Mantle Shear Wave Velocity of 4.0–4.2 km/s
 The primary potential causes of low upper mantle shear wave velocity are composition, increased temperature, the presence of fluids (such as partial melt), grain size reduction [Faul and Jackson, 2005], or some combination of these factors. These factors can be analyzed individually to assess the possible contribution of each to the low observed upper mantle shear wave velocity.
 Composition and grain size can be eliminated as predominant effects because of the large magnitude (5–10%) of the observed decrease. The composition of the mantle is less variable than the composition of the crust, and at most compositional variations cause an ∼1% change in shear wave velocity [Cammarano et al., 2003]. The effect of changing the grain size from 1 cm to 1 mm, the range commonly estimated for the upper mantle [Hirth and Kohlstedt, 2003, and references therein], results in an ∼100 m/s (∼2%) change in velocity at 1200°C (Figure 13), again not enough to account for the 300–500 m/s decrease observed in Ethiopia.
Figure 13. Estimated shear wave velocity versus temperature in the upper mantle from Faul and Jackson  and Priestley and McKenzie . The two curves from the Faul and Jackson  model are for upper mantle grain sizes of 1 and 10 mm. The gray box represents the range of shear wave velocity values observed beneath the MER and the adjacent plateaus. The horizontal dashed black lines show the average upper mantle temperature for the different regions discussed in the text. For comparison, the upper mantle temperature at KIBE is ∼4.5 km/s, which corresponds to a temperature much less than 1000°C for the Priestley and McKenzie  model, which may be the more accurate of the two curves in this temperature range given that the Faul and Jackson  model is only calibrated for temperatures between 1000°C and 1300°C. The two models show different behavior, but for a velocity of 4.1 km/s both predict an upper mantle temperature of ∼1300°C.
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 We next examine the remaining factors: increased temperature and partial melt. These two factors are likely interdependent and combine to reduce the shear wave velocity. With the two effects superimposed, it is difficult to constrain the exact contribution from each. We instead seek to develop reasonable maxima and minima for each factor. Two recent publications [Faul and Jackson, 2005; Priestley and McKenzie, 2006] discuss the relationship between upper mantle temperature and shear wave velocity. The behavior of these two models differs (Figure 13), however both predict that an upper mantle temperature over 1300°C is required to produce a shear wave velocity of 4.1 km/s or lower, and that above 1300°C dVs/dT increases rapidly. On the basis of Faul and Jackson  and Priestley and McKenzie , if there is no contribution from partial melt, the estimated temperature of the upper mantle for seismic velocity ranging from 4.0 to 4.1 km/s (i.e., beneath the MER and NW plateau) is predicted to be 1300–1350°C. To consider the effect of only partial melt with no increase in temperature, we use the estimate of Faul et al.  of a 3.3% reduction in seismic velocity per percent melt. In this melt-only case, 3% melt is required in the upper mantle to reduce the seismic velocity by the observed 10 percent.
 Both high temperature and partial melt are capable of reducing the seismic velocity of the upper mantle. It is implausible however to assume that one of these effects would be present without the other. This is evident in the experimental results of Hirschmann et al. . The authors measured percent melt produced as a function of increasing temperature for upper mantle peridotites, including a peridotite composition similar to that present in Ethiopian xenoliths [Conticelli et al., 1999] (Figure 14). As Figure 14 demonstrates, as upper mantle temperature increases above ∼1270°C there is a rapid increase in the percent partial melt present from near zero to ∼5%. If a temperature of the uppermost mantle of 1300°C is required to reduce shear wave velocity to 4.1 km/s by a purely thermal effect as discussed above, we see that at this temperature ∼5% partial melt would be present (further reducing seismic velocity). The combined effect of high temperature and partial melt would then lower the shear wave velocity below that observed in Ethiopia.
Figure 14. Experimental results for percent melt versus temperature for depleted peridotite, modified from Hirschmann et al.  by permission of Oxford University Press. At 1300°C (our upper bound on temperature if the low upper mantle shear wave velocity below Ethiopia is caused only by high temperature), melt percent would be over 5% on the basis of these experimental results. Black circles are reported measurements, and the solid black line is the predicted curve using the MELTS algorithm.
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 Thus the high temperature that would be required to account for the observed low seismic velocity cannot be present without a contribution from partial melt. In contrast, in principle the observed seismic velocity reduction could be the effect of transient partial melt without anomalous upper mantle temperature. To constrain the lower bound on the upper mantle temperature, we consider independent evidence regarding the thermal structure. Although we lack heat flow data in our study area, mantle xenoliths found in Oligocene and Quaternary basalts near Lake Tana on the Ethiopian Plateau [Conticelli et al., 1999] from a depth range of 40–50 km can be used as estimates of the temperature of the uppermost mantle at the time of eruption. These studies indicate minimum upper mantle temperatures between 920 and 1045°C for the Oligocene and Quaternary samples. This temperature is well above a steady state continental geotherm [Chapman, 1986]. We conclude that the uppermost mantle temperature is elevated by over several hundred degrees above stable continental regions, and that the observed seismic velocity reduction is the result of both high temperature and partial melt. We conclude that the most likely lower bound for the temperature is ∼1250°C, where melt begins to form [Hirschmann et al., 1998], and the most likely upper bound is ∼1300°C (at which the amount of melt produced would create a velocity effect greater than that permitted by observations). This temperature range is compatible with a melt percent of up to 3%. The cause of the high temperature in the uppermost mantle is interpreted to be the thermal perturbation associated with the Afar plume. The strong evidence from previous geophysical and geochemical studies for a thermal perturbation in the mantle down to depths of 600 km is summarized in section 1. Dugda et al.  model instantaneous thinning of the lithosphere due to plume impingement at the time of the flood basalt volcanism (∼30 Ma), with hot plume material remaining to the present. Their models show that the temperature increase (still present today) resulting from the transfer of heat from the plume material to the lithosphere is enough to account for the observed reduction in shear wave velocity [Dugda et al., 2007].
6.2. Lower Crustal Shear Wave Velocity of 3.7–3.8 km/s
 Unlike the upper mantle, where possible compositional effects on seismic velocity are small relative to the observed velocity reduction [Cammarano et al., 2003], in the crust one would commonly assume that the observed lower crustal velocity decrease (to 3.7–3.8 km/s) implies a felsic lower crust. In the following discussion we instead show that the composition is unlikely to be felsic and that the low velocity in the lower crust is most likely the result of a combination of high temperature and partial melt as in the upper mantle.
 Little information on the temperature of the lower crust is available for Ethiopia. Interpreting temperature from seismic velocity in the lower crust is difficult because the value of dVs/dT at high temperatures (>600°C) for crustal lithologies is not well constrained. For temperatures up to 600°C (still significantly below that plausible for the Ethiopian lower crust based on the previous arguments for a 1250–1300°C upper mantle) Kern et al.  show an average of ∼0.21 m/s/K decrease in shear wave velocity as temperature increases. Using a value of 0.21 m/s/K for dVs/dT requires a temperature increase of 1000–2000°C above that of the lower crust in Kenya (∼1500–2500°C) to create the observed decrease of 0.2–0.4 km/s. That temperature range is clearly unrealistic, likely because the dVs/dT value at lower temperatures is not applicable for higher temperatures. Kampfmann and Berckhemer  show a sudden decrease in the shear wave modulus at ∼800°C for mafic rocks, which would cause dVs/dT to rapidly increase above 800°C. Without better constraint on dVs/dT we are unable to quantitatively convert observed seismic velocity in the lower crust to temperature, but using either the Kern et al.  results or the Kampfmann and Berckhemer  results, the temperature in the lower crust is predicted to be 800°C or greater. A hot lower crust is also consistent with extrapolation of the predicted anomalously high upper mantle temperature.
 An independent constraint on the state of the lower crust comes from electrical conductivity data collected along a cross-rift line (Figure 2) by Whaler and Hautot . These data indicate a strong similarity in crustal structure between the NMER and the NW Ethiopian Plateau; the resistivity structure beneath the NMER and the Ethiopian Plateau is nearly the same (as are the seismic velocity profiles) with lower crustal conductivity of ∼4.10−2 S/m. In contrast the Somali Plateau shows a lower conductivity, below 1.10−2 S/m. The integrated lower crustal conductivity (conductance) beneath the rift and the NW Ethiopian Plateau is at least 400 S (siemens), the average conductance for Phanerozoic lower crust, while the conductance below the Somali Plateau is only ∼100 S [Whaler and Hautot, 2006]. High electrical conductivity in the lower crust implies saline fluids, graphite, or melt. Although we cannot a priori rule out any of these, the high temperatures implied by the previous discussion suggest anhydrous equilibrium conditions and the recent magmatism makes melt the natural explanation. If we make the assumption that melt is present in the lower crust and causes the high observed conductivity, we can constrain the amount present beneath the Ethiopian Plateau using the observed conductivity values [e.g., Schilling et al., 2006] independent of seismic velocity.
 Because of the large difference in the conductivity of basaltic melt (up to 10 S/m) and solid rock (<0.01 S/m) [Roberts and Tyburczy, 1999], electrical conductivity is very sensitive to melt percent. This electrical conductivity can be modeled using the Hashin-Shtrikman bounds [Hashin and Shtrikman, 1962], with the upper bound describing an ideal, connected melt network. This bound will describe the lower limit on the amount of melt present. The lower Hashin-Shtrikman bound describes isolated melt pockets and limits the maximum amount of melt present. If we model the upper Hashin-Shtrikman bounds using melt conductivity values of 1 S/m, 3 S/m, and 10 S/m, using a matrix conductivity of 0.001 S/m, we predict a melt percent between 0.5 percent (melt conductivity of 10 S/m) and 5 percent (melt conductivity of 1 S/m) as the lower bound on melt present (Figure 15). This assumes the ideal case of interconnected melt pathways. If some of the melt is present in isolated pockets the melt percent required to produce the observed bulk conductivity is higher. Archie's Law is a reasonable way to estimate a more complex melt system between the Hashin-Shtrikman end-members of completely isolated and completely connected networks [Hermance, 1979; Merzer and Klemperer, 1992]. If we use this method assuming a complex melt system [Hermance, 1979], we estimate between 6% and 20% melt in the lower crust (Figure 15) rather than 0.5 to 5% estimated for ideally interconnected pathways. The presence of melt in the crust beneath the NW plateau away from the MER is supported by elevated bulk crustal Vp/Vs ratios up to 1.9 [Dugda et al., 2005; Stuart et al., 2006]. A Vp/Vs ratio of 1.9 is commonly taken to represent a contribution from melt [e.g., Menke et al., 1998]. This lower crustal melt that we infer is not only present beneath the MER, it is present beneath the entire Ethiopian Plateau even hundreds of kilometers from the rift valley. The plateau has been the location of extensive off-axis Mio-Pliocene and Quaternary volcanism [Abate et al., 1998; Conticelli et al., 1999], possibly reflecting this continued presence of melt far from the rift axis.
Figure 15. Constraints on the amount of melt present in the lower crust in Ethiopia from observed conductivity values of Whaler and Hautot . Blue curves are calculated using the upper Hashin-Shtrikman bounds, which assumes a connected melt network and provides a lower bound on the amount of melt present. Red curves are calculated using the modified version of Archie's Law for materials containing partial melt [Hermance, 1979] for a matrix conductivity of 0.001 S/m and melt conductivities of 1, 3, and 10 S/m. Between 0.5% and 5% melt is present in the lower crust using the H-S upper bound, and between 6% and 20% melt is present using the modified form of Archie's Law.
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 As a final possibility, a lower crust of felsic granulite across the entire region could cause the lower crust to have a shear wave velocity of ∼3.7 km/s [Christensen, 1996]. Such a composition, however, is unlikely. Mafic lower crustal xenoliths are described for both the Arabian-Nubian Shield to the north [McGuire and Stern, 1993; Al-Mishwat and Nasir, 2004] and along the trend of the Mozambique Belt to the south [Mansur et al., 2006]. These terranes meet in Ethiopia, therefore one might expect that the Ethiopian lower crust is composed of a similar mafic lithology. Vp/Vs ratios provide a stronger constraint on composition than either Vp or Vs alone. Vp/Vs ratios in Ethiopia have previously been estimated for the whole crust using the Zhu and Kanamori receiver function stacking method [Dugda et al., 2005; Stuart et al., 2006], but are estimated here for just the lower crust by using the ratio of the lower crustal shear wave velocity measured in our inversions to the lower crustal compressional wave velocity modeled using wide-angle data [Maguire et al., 2006]. The Vp/Vs ratio for the lower crust in Ethiopia ranges from 1.81 to 1.88. As shown in Figure 16, felsic granulite has a Vp/Vs ratio of 1.79, lower than that observed in Ethiopia, and mafic granulite has a Vp/Vs ratio of 1.82, within our observed range.
Figure 16. Vp/Vs ratio versus Vs (km/s) at room temperature and 800 MPa for different rock types using data from Christensen . Vs as shown is higher than expected in the lower crust since data are not temperature-corrected, but Vp/Vs has little dependence on temperature. MGR, mafic granulite; DIA, diabase; GGR, mafic garnet granulite; FGR, felsic granulite; GAB, gabbro. The range of observed lower crustal Vp/Vs values in Ethiopia, calculated using our joint inversion results for Vs and active source Vp profiles, is shown by the horizontal black lines. The Vp/Vs ratio of felsic granulite is lower than the range of observed values in Ethiopia. The Vp/Vs ratio of mafic granulite, present in the lower crust both to the north in the Arabian Shield and to the south in Tanzania, lies within the observed range.
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 The regional mafic lower crustal xenoliths and the observed Vp/Vs ratios cannot absolutely eliminate a felsic lithology in the lower crust as the cause of the low observed lower crustal shear wave velocity in Ethiopia. However, with strong evidence for (1) high temperature causing the upper mantle low-velocity anomaly (section 6.1), which would conduct into the lower crust to reach thermal equilibrium; (2) strong evidence for the presence of melt in the lower crust from conductivity measurements (section 6.2); and (3) suggestions of a mafic lower crust from Vp/Vs ratios and along-strike lower crustal mafic xenoliths (section 6.2), it is quite unlikely that a felsic lower crust is responsible for the low observed shear wave velocities.