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

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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 Main

Ethiopian Rift (MER) preserve details of lithospheric structure in the East African Rift System. Despite observed source heterogeneity, basalts, trachybasalts, and basaltic trachyandesites erupted in the Wonjii Fault Belt (WFB) and the Silti-Debre Zeyit Fault Zone (SDFZ) form coherent fractionation paths dominated by variable removal of observed phenocryst phases. Crustal assimilation is not widespread, though it is observed at the southern end of the WFB where both fault belts merge; farther north, assimilation of cumulate phases related to fractional crystallization of previous magmas is identified. Shallow fractionation conditions (!1 kbar) within the WFB do not change from north to south. In contrast, lavas erupted within the contemporaneous SDFZ fractionate at various crustal depths. These results indicate a better developed magmatic system beneath the WFB where magmas rose quickly before undergoing more significant fractionation at near surface levels and a less developed system beneath the SDFZ. The distribution of magmatism and extant geophysical data indicate thinned crust and a single rift-centered zone of magmatic activity northeast of 8!300N, consistent with a transitional lithosphere between continental and oceanic settings. Southwest of 8!300N, thicker crust and riftmarginal axes of extension suggest lithosphere with continental affinities. The WFB is propagating southward in response to extension within the Red Sea Rift; the northward propagating SDFZ is related to rifting within the East African Rift System. This region records the unification of two rift systems, requiring care in interpreting the MER as simply transitional between continental and oceanic environments. Citation: Rooney, T., T. Furman, I. Bastow, D. Ayalew, and G. Yirgu (2007), Lithospheric modification during crustal extension in the Main Ethiopian Rift, J. Geophys. Res., 112, B10201, doi:10.1029/2006JB004916.

1. Introduction [2] Lithospheric modification during continental rifting is an axiomatic consequence of plate tectonic processes. The processes whereby the continental crust is modified by magmatism during the progressive evolution from continental rifting to seafloor spreading, are however poorly constrained. These ambiguities generate substantial uncertainties in detailing thermal structure, rift-related volcanic hazards, and hydrothermal resources. Rift generated lithospheric heterogeneity also frustrates efforts to construct

1 Department of Geosciences, Pennsylvania State University, University Park, Pennsylvania, USA. 2 Now at Department of Geological Sciences, Michigan State University, East Lansing, Michigan, USA. 3 Department of Geological Sciences, University of South Carolina, Columbia, South Carolina, USA. 4 Now at Department of Earth Sciences, University of Bristol, Bristol, UK. 5 Department of Earth Sciences, Addis Ababa University, Addis Ababa, Ethiopia. 6 Now at Dessie/Kombolcha University, Dessie, Ethiopia.

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

coherent geodynamic and geophysical models of continental rifting in zones of active tectonism, and to interpret areas of ancient rifting along passive margins. The East African Rift System (EARS) stretches for over 3000 km from the Red Sea and Gulf of Aden southward to Mozambique and has been recognized as a major extensional feature for well over 100 years [Gregory, 1896]. It is the classic example of continental rifting, generated by subsidence of fault bounded basins coupled with the uplift of rift flanks [Ebinger et al., 1989]. It comprises two branches flanking the Tanzania craton in central and eastern Africa, and a single arm that traverses the Ethiopian plateau and craton in the north (Main Ethiopian Rift). Ongoing research in the Main Ethiopian Rift (MER) has increasingly pointed toward magmatism as a primary mechanism for extension and associated crustal modification processes [Keranen et al., 2004; Kendall et al., 2005, 2006; Rooney et al., 2005]. This interpretation is consistent with the transition from fault-dominated rift morphology observed in continental rifts toward magma-dominated mid-ocean ridge spreading centers. [3] We present new geochemical data from several key locations along the Main Ethiopian Rift between 8! and 10!N (Figure 1) in order to assess the interaction of magmas

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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 crustal structure. This approach focuses on geochemical indicators of assimilation and fractional crystallization during the residence of magma within the crust. We integrate studies of more primitive rocks [Rooney et al., 2005; Furman et al., 2006] from the same eruptive locations and time periods to chart the evolution of magmas at crustal levels. Specifically, we focus on evolved (65% SiO2) from the Tulu Moye region are marked by filled symbols. basaltic magmas. Sample N-3 exhibits substantially heterogeneous populations of clinopyroxene and feldspars that are clearly xenocrystic. These compositions are observed in felsic 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/Pb that 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 crustal xenoliths 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; Sifeta et 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

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[29] We 206 204

observe no correlation between Ce/Pb and Pb/ Pb, weak correlation between Ce/Pb and 87Sr/86Sr, and no correlation between 207Pb/206Pb and 87Sr/86Sr (Figure 10), suggesting any potential contaminant had isotopic ratios very similar to those of the primary basalts. Unfortunately, the lack of isotopic data for the Ethiopian pan-African crust makes it difficult to rule out this component as a potential contaminant for Quaternary lavas.

Figure 9. Variation of P2O5 with Hf/Sm ratio. Small quantities of apatite will significantly lower the Hf/Sm ratio while elevating P2O5. Samples identified with high Ce/Pb values show deflection toward apatite assimilation. The apatite is an accessory mineral in the fractionating assemblage required to generate evolved MER lavas from primitive basalts [Trua et al., 1999].

samples have La/Nb and Ce/Pb values that plot within the range observed in mantle-derived basalts [Hofmann et al., 1986]. Samples with high Ce/Pb cannot be the result of Ethiopian crustal assimilation (Figure 8). Evolved products of the southern WFB have low Ce/Pb and moderate to low La/Nb values [Trua et al., 1999; Peccerillo et al., 2003], suggesting they are an appropriate assimilant for our samples with anomalously low Ce/Pb. Basalts erupted within the Gedemsa caldera which show petrographic evidence of interaction with intrusive sialic rocks form an array of decreasing 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 are co-genetic or with the Ethiopian crust. [28] Samples with elevated Ce/Pb values (1028, 1029, 1024, N-19, 1017, 1021) are spatially restricted to the region between Fantale and Dofan. These samples are characterized by unusually high CaO/Al2O 3 and low MgO, and exhibit small positive Eu anomalies consistent with feldspar assimilation. The unidentified feldspar-bearing assimilant also contains apatite, suggested by anomalously high P2O5 and Ce/Pb, low Hf/Sm (Figure 9). Apatite is found in silicic rocks in the region and the fractionation of a plagioclase-rich assemblage with accessory apatite has been invoked to produce silicic rocks from mafic magmas in the MER [Trua et al., 1999]. These data suggest that the unidentified assimilant is a cumulate related to the fractionation of more primitive lavas. This interpretation supports a model whereby at least some of the silicic products of the MER are derived through fractional crystallization of primitive magma [e.g., Ayalew et al., 1999] consistent with ponding and fractionation of plume-derived melts at the base of the crust as predicted from coupled petrologicnumerical modeling [Farnetani et al., 1996], and geophysical evidence of high velocity lower crust [MacKenzie et al., 2005].

Figure 10. Trace element and isotopic ratios sensitive to crustal contamination. A weak correlation is observed between Ce/Pb and Sr isotopes, but no such correlation is observed with Pb isotopes. Trace element data for the SDFZ are from Rooney et al. [2005]. Isotopic data for WFB are from Furman et al. [2006] and Rooney [2006].

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Table 2. Xenolith and Host Lava Components With Temperature and Pressure Estimatesa Sample

Wo

En

Fs

Ac

P, Depth, kbar km

N-20 (C)b N-20 N-22 N-22 N-22 N-22 N-22 N-21 N-14 N-1 N-24 (R)b N-7 N-17 N-17 N-16 N-22 N-22 N-22 N-1 N-1 N-1 N-1

46.45 45.57 47.85 46.64 46.30 48.34 46.82 43.87 44.64 46.35 45.62

41.46 41.99 39.24 42.76 40.57 40.48 36.26 42.74 41.66 42.07 36.89

10.68 11.03 11.49 9.46 11.65 9.91 13.77 11.93 12.07 10.41 15.94

1.41 1.41 1.41 1.14 1.48 1.26 3.15 1.46 1.63 1.18 1.55

1.4 0.3 1.1 0.2 0.9 2.0 1.0 0.4 1.5 1.4 0.1

Fo

Fa

Temp

52.1 47.4 64.8 73.9 71.0 81.2 73.9 77.5 81.8 80.5 78.4

46.6 51.6 34.7 25.7 28.4 18.5 25.6 22.1 17.9 19.2 21.2

977 1034 1171 1201 970 1064 1111 1080 1156 1089 1100

5.0 1.1 4.0 0.6 3.1 6.9 3.7 1.5 5.2 5.0 0.4

a Pressure calculated using Nimis and Ulmer [1998] with a standard error of 1.7 kbar. Temperatures calculated using geothermometer (standard error is ±6!C; analytical uncertainty is generally less than ±35!) of Loucks [1996]. b Values are not systematically distributed (e.g., core and rim) except where marked core (‘‘C’’) and rim (‘‘R’’).

[30] An important conclusion from these data is that the majority of WFB lavas do not show clear evidence of interaction with the continental crust. This interpretation contrasts with observations on older lava series (low-Ti 30 Ma flood basalts [Pik et al., 1999]; 11-6 Ma pre-rift lavas [Chernet and Hart, 1999]). This comparison suggests either that (1) crustal modification by basaltic intrusives within the WFB has replaced the pre-rift continental crust with more recent rift-associated magmas and/or (2) crustal residence time is substantially less for Quaternary lavas than for the older flood and fissure basalts [e.g., Chernet and Hart, 1999]. Geophysical evidence supports the hypothesis of substantial intrusions and crustal modification within the MER [e.g., Keranen et al., 2004; Dugda et al., 2005]. [31] To document the impact of magmatism on lithospheric structure, we first determine the conditions of fractionation. The maximum possible storage and fractionation depth for the SDFZ and WFB lavas is limited by the depth of melt generation. Previous studies indicated that these 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 interpreted as partial melt in the depth range 75 – 250 km [Bastow et al., 2005]. The presence of a high velocity lower crustal unit interpreted as dense mafic cumulates [MacKenzie et al., 2005] raises the possibility that magmas may fractionate at the base of the crust (!30– 35 km). Clinopyroxene barometry [Nimis and Ulmer, 1998] indicates that neither the WFB nor the SDFZ show appreciable along-axis variations in the depth of fractionation. However, WFB lavas fractionate at depths of less than 5 km and at modest temperatures (1050 – 1150!C; Table 2) while

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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 some fractionation occurs at deeper levels beneath the SDFZ in comparison to the WFB, consistent with the increased clinopyroxene fractionation recorded in mass balance models. [32] In order to further refine our working model of lowpressure fractionation, we employed a coupled linear least squares 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) as a parental composition. Although this sample has low Ce/Pb (11), other incompatible trace element ratios indicate no crustal contamination (e.g., La/Nb = 0.74) and compatible trace element abundances suggesting only limited removal of olivine from a primary mantle melt (165 ppm Ni; 556 ppm Cr; 33 ppm Sc). Sensitivity tests indicate that to initiate plagioclase growth (as required from the observed phases and mass balance results) requires a water content for N-21 of !0.2 wt.% and pressure of !1 kbar. At lower MgO contents (10 wt.% MgO) may occur at levels deeper than those modeled in our data set. Peccerillo et al. [2003] extend fractionation to lower MgO contents (0.27%) by !40% fractionation of olivine, clinopyroxene, plagioclase and titanomagnetite. The dominance of olivine and plagioclase in the fractionating assemblage is consistent with the low

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An71: 2.70% En40: 6.70% Fo81: 0.80%

88.90% 0.375

An73: 3.50% En43: 8% Fo85: 7.10%

81.80% 0.112

82.40% 0.06

An65: 8.90% En38: 6.80% Fo74: 2.20%

0.01 $0.05 $0.01 0.02 0 $0.02 0 $0.02 $0.04 $0.03

$0.12 0.06 0.01 $0.06 0 0.07 0.07 0.36 0.43 0.17

0.13 $0.07 $0.11 $0.02 0 $0.04 $0.07 0.05 $0.26 $0.04

71.70% 0.594

An71: 14.10% En40: 7.70% Fo81: 5.70%

$0.22 $0.15 0.09 0.07 $0.01 0.05 0.14 0.27 0.64 0.06

From N-18 to 1026 Differenceb

An54: 23.40% En37: 14% Fo61: 5.30% Fe-Ti: 5.20% 52.20% 0.2

0 $0.28 $0.03 0.11 $0.01 $0.12 0.07 $0.06 $0.08 0.29

From 1026 to N-6 Differenceb

91.10% 0.385

En41: 7% Fo85: 2%

0.25 $0.01 $0.38 $0.11 0 $0.06 $0.11 0.29 0.23 $0.12

From DZ-1009 to DZ-1007 Differenceb

Debre Zeyit

94.60% 0.843

En45: 3.90% Fo83: 1.50%

$0.52 0.1 0.18 0.55 0.02 $0.01 0.34 0.07 $0.3 0.12


88.30% 0.38

En45: 8.9% Fo82: 2.2%

$0.35 0.28 0.17 0.21 0 0.05 0.2 0.21 0.12 0.07


86% 0.145

An71: 5.4% En44: 5.6% Fo86: 2.4%

$0.03 0.31 0.01 $0.09 0 0.04 0 0.07 0.18 $0.06

From BJ-1045 to BJ-1048 Differenceb

Butajira

90% 0.384

En45: 8.1% Fo82: 2.6%

$0.12 $0.18 $0.08 0.38 0.01 $0.1 0.1 $0.13 $0.39 0.01


84.80% 0.237

En34: 8.7% Fo80: 2.7%

$0.24 0.06 0.3 0.13 0 0.03 0.12 $0.2 $0.01 0.09


a The primitive samples used here (N-21; N-18; 1030) are presented by Furman et al. [2006]. The model was used to simulate discrete steps (e.g., N-21 to N-18, N-18 to 1026) and not a continuous fractionation trend. The fractionated phases were clinopyroxene (CPX), olivine (OL), feldspar (PL), and titanomagnetite (MT). b Observed minus calculated, times the weight factor. All oxides are 1 except SiO2 (0.4) and Al2O3 (0.5).

SiO2 TiO2 Al2O3 FeO MnO MgO CaO Na2O K2 O P2O5 Fractionated phases PL CPX OL MT Residual liquid Sr2

From 1030 to N-17 Differenceb

From N-21 to N-18 Differenceb

From N-18 to 1030 Differenceb

Wonjii Belt

Table 3. Mass Balance Fractional Crystallization Modeling of the Wonjii Fault Belt Basaltsa

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1256 1 $6.7 QFM+1 47.16 1.91 14.19 2.54 8.33 0.18 10.51 11.09 2.84 0.67 0.37 0.2*

Fo86: 6.53% Di52,Ce15,He13: 9.51% An75: 3.64%

1172 1 $7.63 QFM+1 47.35 2.31 15.46 2.84 9.02 0.23 7.35 10.46 3.43 0.85 0.47 0.25

Calc

N-18 1177 1 $7.57 QFM+1 46.9 2.27 15.85 2.74 8.54 0.19 7.45 10.93 3.23 1.14 0.51 0.25*

Measured

Fo83: 0.86% Di50, Ce14, He13: 5.61% An75: 3.43%

1163 1 $7.73 QFM+1 46.77 2.48 15.91 2.9 9.14 0.22 6.83 10.1 3.51 1.28 0.58 0.28

Calc

1030 1166 1 $7.69 QFM+1 47.27 2.34 16.43 2.78 9.27 0.2 6.87 10.06 3.12 0.78 0.38 0.5*

Measured 1138 1 $8.03 QFM+1 47.23 2.75 15.28 3.14 10.46 0.24 6.24 9.29 3.38 0.92 0.46 0.6

Calc

N-17

Fo82: 2.62% Di44,Ce16,He16: 2.70% An75: 10.25%

Trend ‘‘A’’

1149 1 QFM+1 46.73 2.72 16.40 3.02 10.16 0.21 6.28 9.06 3.35 0.98 0.50 0.6*

1186 1 $8.46 QFM 46.95 2.27 15.87 1.89 9.3 0.19 7.44 10.95 3.24 1.14 0.51 0.25

Measured Parent

N-18

Fo80: 4.55% Di49,Ce13,He16: 10.42% An73: 12.18%

1138 1 $9.03 QFM 46.71 2.92 15.44 2.26 11.05 0.27 5.44 9.32 3.96 1.56 0.71 0.35

Calc

1026 1160 1 $8.77 QFM 47.71 3.18 15.25 2.13 11.08 0.23 5.48 9.81 3.48 0.66 0.63 0.35

Measured

1106 1 $9.44 QFM 51.29 2.65 13.53 2.21 10.66 0.39 4.05 8.08 4.41 1.06 1.06 0.6

Calc

N-6

Fo74: 2.21% Di36,Ce19,He23: 11.12% An63: 19.60% Mt32,Sp42,Usp58: 5.69%

Trend ‘‘B’’

1118 1 QFM 50.65 3.22 15.10 2.05 9.82 0.25 4.03 7.94 4.39 1.30 0.65 0.6*

Measured

a Ghiorso and Sack [1995]. The primitive samples used here (N-21; N-18; 1030) are presented by Furman et al. [2006]. FeO:Fe2O3 for the samples is calculated on the basis of the relevant oxygen buffer and liquidus conditions. No measured water contents are available, and all water values are modeled. Measured values for samples are normalized to 100% given the modeled water contents and FeO:Fe2O3. The model was used to simulate discrete steps (e.g., N-21 to N-18, N-18 to 1026) and not a continuous fractionation trend. The starting material for each step is the measured value for that sample and not the end point of the previous model step. A QFM + 1 oxygen fugacity dropping to QFM was used, similar to that used by Trua et al. [1999]. Initial water contents and pressure were determined by model sensitivity analysis. The model follows two broad paths corresponding to trends ‘‘A’’ and ‘‘B’’ in Figure 4. An asterisk (*) indicates that these water contents are not measured and are taken from modeled values. Fo, fosterite; Di, diopside; Ce, clinoenstatite; He, hedenbergite; An, anorthite; Mt, magnetite; Sp, spinel; Usp, ulvospinel.

Liquidus, !C Pressure, kbar Log(10) f O2 Buffer SiO2 TiO2 Al2O3 Fe2O3 FeO MnO MgO CaO Na2O K2O P2O5 H2O Fractionated phases OL CPX PL MT

Parent

N-21

Table 4. Thermodynamic Modeling of Wonjii Fault Belt Basalts Using MELTSa

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modeled pressure of !0.5 kbar for samples from Gedemsa, slightly shallower than modeled in this study as expected for lavas erupted within an existing caldera system. 5.2. Spatial Variations in Crustal Structure [35] The data presented here allow for comparison of fractionation depths across the rift (east-west, between the WFB and SDFZ) and along the length of the rift (northeastsouthwest, along the WFB). This three-dimensional framework of fractionation depths within the rift has direct relevance to existing geophysical work that has pointed to the presence of melt within the crust. Analysis of teleseismic receiver functions provides information on bulk crustal 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) and point to melt within the crust, but this method is unable to locate this melt precisely. Kendall et al. [2005, 2006] also suggest the presence of melt within the crust to explain observations of seismic anisotropy beneath the MER. Analyses of magnetotelluric data have suggested melt within the crust beneath the SDFZ and WFB at both shallow (