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Precambrian Research 240 (2014) 108– 125

Contents lists available at ScienceDirect

Precambrian Research

journa l h om epa ge : www.elsev ier .com/ locate /precamres

he Neoproterozoic evolution of the central-eastern Bayuda DesertSudan)

avid Evuka,∗, Gerhard Franza, Dirk Freib, Friedrich Lucassenc

Department of Mineralogy–Petrology, Technical University Berlin, Ackerstrasse 76, D-13355 Berlin, GermanyStellenbosch University, Department of Earth Sciences, Private Bag X1, 7602 Matieland, South AfricaDepartment of Geosciences and MARUM – Center for Marine Environmental Sciences, University of Bremen, D-28334 Bremen, Germany

r t i c l e i n f o

rticle history:eceived 23 April 2013eceived in revised form 21 October 2013ccepted 26 October 2013vailable online 21 November 2013

eywords:ayuda Terraneeoproterozoic pulses

n situ dating

a b s t r a c t

In the central-eastern Bayuda Desert (Sudan), a border area between the Saharan Metacraton and juve-nile Pan-African crust of the Nubian Shield, the Rahaba-Absol and the Abu Harik-Kurmut Terranes wereamalgamated along suture and shear zones. U–Pb zircon ages by laser-ablation dating of metamor-phosed granitoid rocks reveals ages interpreted as magmatic and metamorphic activities throughout theNeoproterozoic. The meta-granitoids are calc-alkaline porphyroblastic varieties of biotite metagranite(969 ± 5 Ma), quartzfeldspathic metagranite (914 ± 6 Ma), biotite–muscovite metagranite (912 ± 4 Ma)and coarse-grained meta-monzodiorite (909 ± 9 Ma, 818 ± 19 Ma, 669 ± 13 Ma) associated with theRahaba-Absol Terrane, and medium-grained metagranites (813 ± 4 Ma, 808 ± 5 Ma, 799 ± 16 Ma), a meta-quartz–monzonite (810 ± 10 Ma), biotite metagranites (794 ± 15 Ma, 783 ± 13 Ma, 700 ± 7 Ma), alkalimetagranite (645 ± 5 Ma), and porphyritic meta-quartz-monzonite (630 ± 4 Ma) associated with the AbuHarik-Kurmut Terranes. Th/U ratios of the zircon were used to distinguish between intrusion and meta-morphic ages.

The Neoproterozoic ages are interpreted as geodynamic pulses: a pre-Pan-African event 1000–900 Macharacterized by metamorphism and magmatism ending with the known Bayudian Event (920–900 Ma)in the Rahaba-Absol Terrane; a cycle between 850 and 775 Ma of subduction-related magmatism andmetamorphism with a peak around 825–800 Ma; the final collision of the Rahaba-Absol and AbuHarik-Kurmut Terranes began before 714 Ma and ended probably before 645–630 Ma, documentedby discordant E-W trending alkali metagranites and NE trending porphyritic meta-quartz–monzoniterespectively. Theses ages of discordant intrusions constrain the post-collisional horizontal movementalong the southern Keraf Shear Zone between 630 and ∼590 Ma. The existence of inherited zircon grains

in some samples of Pan-African age indicated sedimentary activity contemporaneous with the island-arcmagmatism and metamorphism. This was probably a period of extensive erosion of the earlier island arcmaterial and later stacked crustal material after terminal collisional between the Abu Harik-Kurmut arcand the Rahaba-Absol continental terrane. Pb isotope data (feldspar, whole rock) and Sr isotopes cor-roborate the division into two terranes, the Rahaba-Absol Terrane as an old cratonic part of the SaharanMetacraton, and the Abu Harik-Kurmut Terrane as juvenile island-arc additions during the Pan-African.

. Introduction

The Neoproterozoic is characterized by a world-wide reorga-ization of continental crust, which resulted from the breakupf the supercontinent Rodinia and Greater Gondwana assem-

ly. Many parts of the basement in Africa, South America, India,ntarctica and Australia show remobilization of older cratons asell as juvenile growth of continental crust, regionally distributed

∗ Corresponding author at: Huttenstrasse 5, D-10553 Berlin, Germany.el.: +49 30 314 72222; mobile: +49 17636137622.

E-mail addresses: evuk 74@yahoo.com, dvdevuk@googlemail.com (D. Evuk).

301-9268/$ – see front matter © 2013 Elsevier B.V. All rights reserved.ttp://dx.doi.org/10.1016/j.precamres.2013.10.015

© 2013 Elsevier B.V. All rights reserved.

e.g. in the Pan-African and in South America in the Brasilianobelts. Major new insights came predominantly from in situ radio-metric age determinations, which allow distinguishing in detaildifferent episodes in the Neoproterozoic. It is believed that theTonian (1000–850 Ma) witnessed little crust formation while theCryogenian (850–635 Ma) experienced intense magmatic activitybeginning with Rodinia breakup (Li et al., 1999; Torsvik, 2003)with episodic plume events at about 825 Ma, 780 Ma and 750 Ma(Li et al., 2008; Stern, 2008). The earliest arc magmatism related

to West Gondwana assembly was found at ∼900–800 Ma in bothSouth America and Africa, coeval with the earliest stages of Rodiniabreak-up (Cordani et al., 2001; Kampunzu, 2001; Yoshida et al.,2003). Ages of 930 Ma for early oceanic arcs related to Gondwana

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ssembly have also been reported from South America (Yoshidat al., 2003).

The term Pan-African was introduced by Kennedy (1964) toescribe the tectono-thermal events in the non-cratonic terrain infrica with dominance of 650–540 Ma K–Ar ages. This was because

here was no sign of classic orogenesis or Barrovian regional meta-orphism (Gass, 1977). Kroner (1984) redefined Pan-African as

nvolving a protracted orogenic cycle from 950 to 450 Ma. Stern1994) noted that the time span of 500 Ma is much longer than anyhanerozoic orogeny, so reference to a Pan-African orogenic cycleust suffice until the timing and regional extent of discrete tec-

onic event is better constrained. New information from isotopicnd geochronological studies in North Africa, Arabia, Eastern andouthern Africa down to Antarctica, Australia and South America,xtended the concept to Greater Gondwana and defines now thean-African Orogeny (∼870 to ∼550 Ma) that includes the geo-ynamic activity of a complete Wilson Cycle (Kröner and Stern,004). Because of its large geographical and temporal extent, thean-African is considered to be a protracted orogenic cycle rep-esenting opening and closing of large oceanic tracts as well asmalgamation of crustal blocks by accretion and collision (Krönernd Stern, 2004). The discussion on the limitation of the usage ofhe term Pan-African is ongoing (Stern, 2008). However, Liégeois1998) and Küster and Liégeois (2001) defined the term Pan-Africanrogeny for correlating events at Bayuda Desert as the Neopro-erozoic orogeny that led to the Pangaea Supercontinent, includingubduction, collision and post-collision periods, excluding eventsuch as continent break-up or passive margin sedimentation. Here,eference to Pan-African orogeny encompasses events related toodinia break-up starting in Early Neoproterozoic (post-Bayudianvent) and ending in the Early Paleozoic (the age of the Amaki Series540 Ma).

Based on zircon geochronology and Sr–Nd–Pb isotope geo-hemistry of granitoids, Küster et al. (2008) identified a Bayudianetamorphic and magmatic event around 920–900 Ma that they

onsidered pre-dated the Pan-African orogeny and concluded thathe Pan-African event (860–590) in the Bayuda Desert was causedy collision between the Rahaba-Absol Terrane (continental) andhe Kurmut (arc) Terrane (Fig. 2). Major horizontal movementsccurred at the Keraf Shear Zone, east of the Bayuda Desert, after theerminal collision between east and west Gondwana (Abdelsalamt al., 1998; Fig. 1). We will refer to events after the Bayudian<900 Ma) event as related to the onset of Pan-African orogeny inhe central-eastern Bayuda Desert to confirm with previous usagef the term in this poorly studied region. Küster and Liégeois (2001)lready reported Nd model ages of 899 and 888 Ma (Table 1) fromam El Tor and Kurmut gneisses with less depleted mantle signa-

ure of εNd +5.3 and +5.0 respectively.In NE Africa, two crustal domains are distinguished: the Saha-

an Metacraton (Abdelsalam et al., 2002) and the Nubian ShieldYoshida et al., 2003), which resulted from tectonic, magmaticnd metamorphic activity of the Neoproterozoic to Early Paleozoican-African Orogeny (Fig. 1). The study area in the central-easternayuda Desert, Northern Sudan, is located west of the River Nilend west of the Keraf Zone, on the eastern boundary of the Saha-an Metacraton (Figs. 1b and 2). The Bayuda Terrane was alwaysonsidered as a remobilized cratonic area in transition to theolcano-sedimentary island-arc and back-arc assemblages to theast (Vail, 1983, 1985; Vail et al., 1984; Meinhold, 1979, 1983;röner, 1985; Ries et al., 1985; Abdelsalam and Stern, 1996a). How-ver, the Dam El Tor and Abu Hamed fold-and-thrust belts withinhe Bayuda Terrane (Fig. 1b) were later interpreted as a re-entrant

f oceanic origin to the west of the perceived boundary between theemobilized western cratonic area and the Arabian-Nubian ShieldAbdelsalam et al., 1998). Based on the Sr, Nd isotopic characteris-ics and geochemistry of the high-grade metamorphic basement,

rch 240 (2014) 108– 125 109

Küster and Liégeois (2001) proposed a boundary even furtherwest probably along the Zalingei fold zone (Fig. 1). Abdelsalamet al. (2002) considered the Bayuda Terrane as part of the SaharanMetacraton, a pre-Neoproterozoic highly remobilized continentalcrust during Neoproterozoic as a result of collision between Eastand West Gondwana.

In summary, the localization of the Bayuda Desert and the sev-eral events from the Early to the Late Neoproterozoic makes theunderstanding of the Neoproterozoic pulses in the central-easternBayuda Desert a challenging task. New detailed field work, pet-rographic, geochemical and isotopic characterization, and in situU–Pb dating of zircons of meta-igneous rocks contribute to theunderstanding of the complex pre-Pan-African and Pan-African his-tory at the transition between the cratonic and the juvenile crust.The main objective was to determine the emplacement ages for themeta-igneous rocks and check for inherited ages or resetting andto reconstruct the Neoproterozoic evolution of the area.

2. Geological setting

The basement of the Bayuda terrane is subdivided into twomajor compositional, tectonic and age domains, the Saharan-Metacraton and the Nubian Shield sector of the Arabian-NubianShield (Fig. 2). The Saharan Metacraton (Abdelsalam et al., 2002)extends from the Tuareg Shield in the west to the Nubian Shield inthe east, and from the Congo Craton in the south to the Phanero-zoic cover in the north (Fig. 1). It is part of West Gondwanaand comprises predominantly medium- to high-grade old Pre-cambrian gneisses, metasediments, migmatites and rare granulitesintruded by voluminous granitoids, and was remobilized in theNeoproterozoic by collision, delamination of the sub-continentalmantle lithosphere, extension, and assembling of the metacratonfrom exotic terranes with further post-collisional dismembering byhorizontal shearing. The Nubian Shield east of the Saharan Metacra-ton (together with its continuation onto the Arabian Peninsula)consists to a large degree of juvenile crustal material accretedduring the period 900–700 Ma (Greenwood et al., 1980; Kröneret al., 1987a, 1992a). It is the northern part of the East AfricanOrogenic belt, which is a structurally complex assemblage of island-arc, back-arc and ophiolitic rocks formed during the Middle toLate Neoproterozoic-Early Paleozoic and presumably controlledthe amalgamation of East and West Gondwana (Teklay et al., 1998;Yibas et al., 2002).

In the central-eastern Bayuda Desert, the series of the Rahaba-Absol Terrane (Küster et al., 2008) as part of the Saharan metacratoncomprises porphyroblastic metagranitoids, gneisses and metased-iments (Fig. 2). The lithostratigraphy, based on age data from theliterature and the new age determinations of this study, is listedin Table 1. The Rahaba Series is made up of muscovite gneiss,muscovite-biotite gneiss, and biotite gneiss, the Absol Series com-prises quartz–mica schist, kyanite–staurolite–garnet–mica schist,tourmaline–mica schist, graphitic schist, manganiferous schist, fer-rogenous quartzite, amphibolites and hornblende gneiss (Barth andMeinhold, 1979). It extends from north of Dam El Tor fold-and-thrust belt (FTB; Fig. 2) to south of Abu Hamed FTB (Fig. 1b). TheAbsol Series interfingers with gneissic quartz–feldspathic RahabaSeries forming the core of a syncline.

The Abu Harik-Kurmut Terrane is part of the Nubian Shield.The Abu Harik Series (Barth and Meinhold, 1979) is made upof biotite gneisses, biotite–hornblende gneisses and amphibo-lites. An ultrabasic ophiolitic complex of olivine–gabbronorite,

olivine–websterite to Iherzolite and serpentinite was reportedfrom the Bayuda Terrane at a locality just adjacent to the centralDam El Tor FTB (Fig. 1) by Barth and Meinhold (1979). Pyrox-enites, metagabbro, and metachert have been reported at Fadlab

110 D. Evuk et al. / Precambrian Research 240 (2014) 108– 125

Fig. 1. (a) Distribution of rocks of the Saharan Metacraton, Neoproterozoic juvenile rocks and pre-Neoproterozoic crust in Northern Africa. Dashed line is approximateb ion an2 ); redy ◦ N/33

(Fwro1KaTavanq

oundary of the Nubian Shield in Sudan and Southern Sudan. (b) Terrane distribut002; Johnson and Woldehaimanot, 2003; Küster et al., 2008; Liégeois et al., 2013ielded information about the Recent lower crust (Lucassen et al., 2011) lies at ∼18

Esam, 2005), and east of the Nile (Lissan, 2003; Evuk, 2007;ig. 1). We found small outcrops of serpentinite and pillow basaltest of and in the central part of Dam El Tor FTB. These rocks

epresent suture zones and are interpreted as dismembered ophi-lites (Abdel-Rahman, 1993; Abdelsalam, 1996c; Abdelsalam et al.,998 and references therein). Low-grade rocks associated with theurmut Series include chlorite schist, chlorite–actinolite schist,lbite–sericite schist, muscovite schist, calc–silicates and marble.hese rock units are dominantly found near the Nile concentratinglong the Keraf Shear Zone (Figs. 1 and 2). The volcanic protoliths

ary from basalt to rhyodacite, the sediments from graywacke,rkose, sandstone to shale (Küster and Liégeois, 2001). We mappedew occurrences of metamorphosed gabbro, diorite, granodiorite,uartz-monzonite, granitoids and volcano-sedimentary sequences

d major tectonic elements of north-east Sudan (compiled from Abdelsalam et al., box shows study area. The Bayuda Volcanic Field, where lower crustal xenoliths◦ E.

in the Abu Harik-Kurmut Terrane. The Amaki Series (El Rabaa, 1976)is restricted to the eastern and western side of the River Nile justsouth of the Abu Hamed FTB (Fig. 1). It is made up of undeformed,unmetamorphosed to slightly metamorphosed molasse-type sedi-ments (conglomerate, sandstone, quartzitic sandstone and marlyshale) marking the last sedimentation in a pull-apart basin (ElRabaa, 1976; Dawoud, 1980; Abdel Rahman, 1993; Abdelsalamet al., 1998) at the end of the Pan-African. Post-collisional plu-tons consist of grey to reddish biotite granite and of subordinatebiotite–hornblende granite and granodiorite (Küster et al., 2008).

The primary targets of this study are the widely distributedmeta-igneous rocks (Fig. 2), because zircons in such rocks bearage information on their crystallization and possible metamorphicoverprint. Their elemental and isotopic signatures provide evidence

D. Evuk et al. / Precambrian Research 240 (2014) 108– 125 111

Table 1Lithostratigraphy of dated rocks from the Bayuda Desert. The respective mineral used in the Rb/Sr method is indicated in brackets after the series and rock unit.

Series and Rock unit Age (Ma) Method Reference

Post-collisional granitoidsEast Abu Nahal (Gneiss/granite-biotite) 555 Rb/Sr Barth and Meinhold (1979)Nabati granite 573 ± 70 Rb/Sr Barth et al. (1983)Nabati granite 573 ± 72 Rb/Sr Meinhold (1979)Khor Rahaba 2 (Pegmatite-muscovite) 594 ± 24 Rb/Sr Vail (1971)An Ithnein granite 597 ± 4 Rb/Sr Küster et al. (2008)

Amaki Late orogenicAbu Harik and KurmutPorphyritic meta-quartz-monzonite (397/1) 630 ± 4 U/Pb Current workWadi Absol (Gneiss/granite-biotite) 635 Rb/Sr Barth and Meinhold (1979)Rahaba 1b (Pegmatite-muscovite) 641 ± 26 Rb/Sr Vail (1971)Rahaba IIb (Pegmatite-muscovite) 645 ± 26 Rb/Sr Vail (1971)Alkali metagranite (293/1) 645.3 ± 5.4 U/Pb Current workWadi Absol (Gneiss/granite-muscovite) 690 Rb/Sr Barth and Meinhold (1979)Biotite metagranite (324/1) 700 ± 7 U/Pb Current workKurmut gneiss 757 ± 29 Rb/Sr Meinhold (1979)Biotite metagranite (376/1) resetting at (714 ± 16) U/Pb Current workBiotite metagranite (376/1) 783 ± 13 U/Pb Current workKurmut Metasediment 787 Nd model Küster and Liégeois (2001)Biotite metagranite (150/1) 794 ± 15 U/Pb Current workMedium-grained metagranite (313/1) 799 ± 16 U/Pb Current workDam ElTor Amphibolite 806 ± 19 Sm/Nd Küster and Liégeois (2001)Medium-grained metagranite (068/1) 808 ± 5 U/Pb Current workMeta-quartz-monzonite (073/3) 810 ± 10 U/Pb Current workMedium-grained metagranite (135/1) 813 ± 4 U/Pb Current workDam ElTor epidote biotite gneiss 858 ± 7 U/Pb Küster et al. (2008)

Rahaba and AbsolRahaba gneiss 874 ± 33 Rb/Sr Meinhold (1979)Kurmut gneiss 888 Nd model Küster and Liégeois (2001)Abu Harik granite 898 ± 51 Rb/Sr Ries et al. (1985)Dam ElTor gneiss 899 Nd model Küster and Liégeois (2001)Absol granodiorite 900 ± 9 U/Pb Küster et al. (2008)Dam ElTor Amphibolite 903 Nd model Küster and Liégeois (2001)Meta-monzodiorite (371/2) resetting at (669 ± 13) U/Pb Current workMeta-monzodiorite (371/2) resetting at (755 ± 15) U/Pb Current workMeta-monzodiorite (371/2) resetting at (818 ± 19) U/Pb Current workMeta-monzodiorite (371/2) 909 ± 9 U/Pb Current workPorphyroblastic bt-msc metagranite (059/1) 912 ± 4 U/Pb Current workPorphyroblastic qtz-feldspathic metagranite (015/1) 914.1 ± 5.5 U/Pb Current workEl Melagi msc-bt gneiss 921 ± 10 U/Pb Küster et al. (2008)El Melagi msc-bt gneiss 917 ± 14 U/Pb Küster et al. (2008)Porphyroblastic biotite metagranite (193/1) 969 ± 5 U/Pb Current work

A c are f

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ges in brackets are resetting ages; ages in bold are from current work, ages in itali

n their origin from juvenile mantle sources or recycling of compo-itionally distinct old craton. Field work and sampling was focusedn selected locations near the presumable transition between theetacraton and the Nubian Shield. Dioritic to gabbroic rocks (with

lagioclase, green-amphibole, secondary epidote minerals, quartz,ccessory oxides and rare zircon) were not considered for isotopicracing and zircon age determination, because they restrict to minorccurrences in tectonically deformed dykes intercalated with theetagranitoids.

. Petrography and sample material

The studied gneisses (sample locations see Fig. 2) are situatedithin the Rahaba-Absol (Saharan Metacraton) and the Abu Harik

nd Kurmut Terranes (Nubian Shield). Most of the outcrops areow-lying at the erosion surface with some exceptions of pinkish

igmatite and whitish granites that can rise between 10 and 20 mbove the surrounding rocks. The samples for Pb–Sr isotope anal-sis and U–Pb zircon dating were selected from a large suite of

amples excluding altered/weathered fragments. Based on texturend field occurrence they are dominantly granitoid meta-igneousocks. In the Rahaba-Absol Terrane they are structurally and tex-urally distinct from those of the Abu Harik–Kurmut Terranes. In

Nd model Küster and Liégeois (2001)Nd model Küster and Liégeois (2001)

rom literature.

the north, they are gneissic, medium and coarse-grained with K-feldspar porphyroblasts (details are shown in Appendix B Fig. 1a,b) with a foliation marked by biotite trending generally NNE to NE.The medium to coarse-grained groundmass is composed of alkali-feldspar, quartz and plagioclase. Interlocking sub-hedral to euhe-dral magmatic texture is locally preserved. Muscovite is present inthese rocks either as a primary mineral or formed after biotite. Gar-net is rare. Accessory minerals are zircon, ilmenite and rutile. Threetypes of porphyroblastic gneiss are distinguished in the field (Fig. 2),biotite gneiss, biotite–muscovite gneiss, and quartzfeldspaticgneiss. We selected representative samples of each type for dating(Table 1). A meta-monzodiorite (sample 371/2), composed dom-inantly of plagioclase forming euhedral interlocking crystals andaffected by epidotization, was also selected for dating.

In the Kurmut Terrane the pinkish to pinkish-grey metagran-itoids are multiply folded and sheared. NE trending foliation isdominant around Dam El Tor FTB, NS to NNW trending foliationis dominant along Keraf Shear Zone, while areas encompassedbetween southern Dam El Tor and western Keraf Shear Zone have

multiple foliation trends because of the multiple deformationphases. Their textures vary from gneissic, mylonitic, porphyriticto pegmatitic (Appendix B Fig. 1c–f). These rocks appear mainly asformer plutons, dykes and stocks that have been stacked together

112 D. Evuk et al. / Precambrian Research 240 (2014) 108– 125

F Thruss b and

bsaagcgmsi0

ig. 2. Location of the analyzed meta-igneous rocks with respect to Dam ElTor Fold/tudy are bold, see Table 1 for the sample numbers. Numbered samples have only P

y deformation thereby forming interleaved lenses and bands withheared boundaries. Only a porphyritic meta-quartz-monzonitend an alkali-metagranite (samples 397/1 and 293/1; Table 1)ppear discordant to the multiple foliation directions. Biotite meta-ranites (samples 071/1, 150/1, 324/1 and 376/1) are dominantlyomposed of subhedral to anhedral crystals of alkali-feldspar, pla-ioclase and quartz (Appendix B Fig. 2c–f). Biotite is the dominant

afic mineral and locally altered to muscovite and chlorite. Some

amples (071/1 and 376/1) and a granodiorite (498/1) have signif-cant amounts of epidote but no garnet, whereas others (samples80/1, 097/2 and 177/1) have significant garnet and minor epidote.

t Belt and the Keraf Shear Zone. Ages from literature are in italics. Ages from current Sr isotope determined. See Table 1 for further explanations.

Samples 006/1 and 150/1 have neither epidote nor garnet. Theaccessory minerals include iron oxide, zircon, titanite, apatite andsometimes calcite as secondary mineral filling fractures.

4. Analytical techniques

A total of 61 samples were selected and prepared into thin-

sections for petrographic studies under the polarized microscopeand XRF analyses at the Department of Mineralogy and Petrology,Technical University of Berlin. For elemental and isotopic analysessamples were first crushed using a hammer and, after removing

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D. Evuk et al. / Precambrian

f parts with visible alteration or with xenoliths, passed through aaw crusher. The crushed samples were cleaned in distilled watern order to remove the dust fraction. Approximately 100 g of frag-

ents (∼10–5 mm) were selected under the binocular microscope,hen ground in a tungsten-carbide chatterbox for 5 min. The lossn ignition was determined on ca. 2 g of dry (110 ◦C) powderedample at 1000 ◦C. Whole rock samples were analyzed for majorlements by XRF (PHILIPS WD-RFA PW 2400) on fused glass diskssing Oxiquant and X-40 Philips software. Reproducibility of the

n-house standard is (n = 90; wt% ± 1sd): SiO2 60.1 ± 0.75; Al2O39.0 ± 0.27; Fe2O3 5.8 ± 0.11; MnO 0.13 ± 0.004; MgO 1.1 ± 0.13;aO 5.1 ± 0.06; Na2O 4.6 ± 0.23; K2O 2.2 ± 0.04; TiO2 0.56 ± 0.012;2O5 0.20 ± 0.008.

Sr and Pb isotope ratios of whole rock samples and Pb isotopeatios of potassic feldspar from selected samples were analyzedy thermal ionization mass spectrometry (TIMS) on a Triton plus

nstrument (Thermo Scientific) at the isotope chemistry laboratory,arum/Department of Geosciences, Universität Bremen. Approx-

mately 20–50 mg rock powder was dissolved in a 5:1 mixture ofouble distilled HF and HNO3, dried, and re-dissolved in 600 �l of 2olar HNO3 for chemical separation. Hand-picked multigrain sep-

rates (10–25 mg) of potassic feldspars were leached in HNO3 andiluted HF (e.g. Housh and Bowring, 1991) before final dissolution

n order to remove potential radiogenic Pb precipitates on grainoundaries and fractures. Strontium and Pb were isolated fromhe matrix elements by using miniaturized columns with 70 �lf Sr- and Pb-specific resin (Sr.spec 50–100 �m particle size byrisKem® International/France). The separation of Sr and Pb fromilicate samples with Sr.specTM in a single stage column run wasreviously described by Deniel and Pin (2001) and was adaptedo the smaller resin-volumes used in this study. Procedure blanksor Sr and Pb are <40 pg. No blank corrections have been appliedo the analyzed ratios because blank contributions were insignif-cant in comparison with the amount of the respective elementsn the sample. Strontium samples were loaded with Ta-emitter one filaments and analyzed by TIMS in the static multi-collectionode. Isotope ratios (Table 2) were normalized to 86Sr/88Sr of

.1194. The external reproducibility according to the NIST 987tandard material is 87Sr/86Sr 0.710247 ± 14 (2sd, n = 50; period:anuary 2012–September 2012). Lead samples were loaded withilica-emitter on Re filaments and Pb isotope ratios (Table 2) werenalyzed at controlled temperature of ∼1250 ◦C in the static multi-ollection mode. Instrumental mass-fractionation has been cor-ected using 0.1% per atomic mass unit. This factor was derived fromepeated analyses of NIST 981 and comparison of analyzed and cer-ified Pb isotope ratios. The reproducibility including the correctionor mass-fractionation is better than 0.1% of the respective ratio.

Thirteen meta-igneous samples were selected for dating zircon.hese samples have >100 ppm Zr and show many zircon grainsn thin-section. They were broken into smaller fragments using

hammer and a jaw crusher, which were grinded to ≤1 mm sizesing a mill. The powder plus the rock fragments were individuallyieved through sieve tubes of 1000 �m, 500 �m, 355 �m, 250 �mnd 125 �m sizes, rinsed with water and dried at ∼120 ◦C forn hour. The grain size fractions <125 �m, 125–250 �m and50–355 �m were separated through a magnetic separator fromagnetic fractions, then separated with heavy liquid bromoform

f density 2.82 g/cm3 at GFZ in Potsdam, again with magneticeparator under a higher current of 1.6 A and diiodomethane withensity of 3.24–3.31 g/cm3 to separate pure zircon grains fromemaining apatite, biotite and opaque minerals. The zircon grainsere hand picked under the binocular onto a sticker tape at Cen-

ral Analytical Facility, Stellenbosch University/South Africa. Therains were mounted in resin and then dried at 40 ◦C. The mountas polished, gold coated in an S150 Ampere SPUTTER COATER

nd wrapped with carbon tape to facilitate current conduction.

rch 240 (2014) 108– 125 113

The cathodoluminscence images were acquired for each zircongrain using CL, LEO 1430 VP, under 10 nA, with 15 mm workingdistance.

The U–Pb age data were obtained at the Central AnalyticalFacility, Stellenbosch University by laser ablation-single collector-magnetic sectorfield-inductively coupled plasma-mass spectrom-etry (LA-SF-ICP-MS) employing a Thermo Finnigan Element2 massspectrometer coupled to a NewWave UP213 laser ablation system.All age data presented here were obtained by single spot analyseswith a spot diameter of 30 �m and a crater depth of approximately15–20 �m. The methods employed for analysis and data processingare described in detail by Gerdes and Zeh (2006) and Frei and Gerdes(2009). For quality control, the Plesovice (Sláma et al., 2008) andM127 (Nasdala et al., 2008; Mattey, 2010) zircon reference mate-rials were analyzed, and the results were consistently in excellentagreement with the published ID-TIMS ages. Full analytical detailsand the results for all quality control materials are reported in Table2 (Appendix A) in the electronic supplementary material. The cal-culation of concordia ages and plotting of concordia diagrams wereperformed using Isoplot/Ex 3.0 (Ludwig, 2003).

5. Geochemistry

5.1. Whole-rock major, minor and trace elements geochemistry

Data plotted for the rocks are shown in the Appendix A Table 1;the plots are shown in Fig. 3; Appendix B Figs. 3–5. The metagran-itoid rocks have SiO2 between 55 and 76 wt%. Other meta-igneousrocks are gabbroic, monzodioritic to monzonitic (the dated sam-ple 371/2). On the total alkali-silica diagram (Middlemost, 1985),most of the samples plot as granite and granodiorite (Fig. 3a).Samples that lie within shear zones have SiO2 > 76 wt% and arein most cases mylonitized, indicating that secondary silica enrich-ment was probably due to shearing. The major and trace elementsare consistent with the interpretation as meta-igneous, dominantlymetagranitoid rocks. The Al2O3 content varies between 11 and∼25 wt%, Fe2O3tot varies, but is generally < 4 wt%. Samples withhigher Fe2O3tot up to 10.5 wt% have also a higher CaO-content upto 6.9 wt%. Samples 015/1, 145/1 and 195/1 show high K2O con-tent between 7.1 and 8.4 wt%. All selected samples for dating arecalc-alkaline (Fig. 3b, AFM diagram after Irvine and Baragar, 1971;ages annotated). With the exception of Na2O, which shows no cor-relation, all the other major and minor oxides are well correlatedwith SiO2 (Fig. 3c and Appendix B Fig. 3a–h). Most of the oxidesdecrease with increasing silica, K2O increases with increasing silicacontent. Except for the monzodiorite-monzonite that is transitionalthe rocks vary from metaluminous to peraluminous (Fig. 3d). Onthe K2O versus SiO2 plot of Peccerillo and Taylor (1976), the meta-granitoids spread through the tholeiitic, calc-alkaline and high-Kcalc-alkaline series (Fig. 3e). On the MALI (modified alkali-limeindex) versus silica plot after Frost et al. (2001), the metagranitoidsare classified as calcic, calc-alkalic and alkali-calcic with one sam-ple plotting as alkalic and three at the boundary with alkali-calcic(Fig. 3f). Based on the Fe-number, the metagranitoids are separatedas magnesian as well as ferroan (Appendix B Fig. 3i). On the Rbversus Y + Nb and Nb versus Y diagram after Pearce et al. (1984),most of the granitoids plot as volcanic arc and syn-collisional gran-itoids with exception of samples 005/1, 063/1, 069/4, 152/1, 209/1,211/1, 293/1, 376/1 and 397/1that always plot as within-plate gran-itoids in both diagrams (Appendix B Fig. 4a, b), probably due toelevated content of Nb. Whalen et al. (1987) employed Ga/Al, var-

ious major element ratios and Y, Ce, Nb and Zr to discriminatebetween A-type granites and most orogenic granites M-, I and S-types. Our samples plot dominantly as I- and S-type granites, someas A-type granite (Appendix B Fig. 5a–i).

114D

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esearch 240 (2014) 108– 125

Table 2Pb and Rb–Sr isotope data.

No. Location 87Sr/86Sr Rba (ppm) Sra (ppm) Ageb (Ma) 87Rb/86Sr 87Sr/86Sr 206Pb/204Pb 207Pb/204Pb 208Pb/204Pb Feldspar

Lat. Long. (UTM) Analyzed 2 Standarderrors

Initial 206Pb/204Pb 207Pb/204Pb 208Pb/204Pb

006/1 (KT) 2,054,204 550,571 0.708310 (i) 0.000013 102 58 630 5.04 (0.6630) 19.17 15.60 37.53015/1 (RAT) 2,061,475 541,333 0.750251 (ii) 0.000004 214 154 914 4.03 (0.6976) 17.97 15.74 38.15 17.98 15.74 38.20018/1 (RAT) 2,063,662 538,388 0.721820 (ii) 0.000007 129 207 912 1.80 (0.6984) 19.98 15.76 38.03022/4 (RAT) 2,073,341 530,266 0.735502 (ii) 0.000018 118 155 912 2.21 0.7066 18.98 15.80 40.95030/1 (RAT) 2,075,891 534,251 0.858587 (ii) 0.000013 232 59 912 11.44 0.7095 19.55 15.83 37.79059/1 (RAT) 2,049,321 516,101 0.734840 (ii) 0.000007 126 195 912 1.86 0.7105 18.83 15.78 39.09063/1 (KT) 2,040,983 562,390 1,042,351 (iii) 0.000012 142 29 808 14.58 0.8741 20.30 15.66 38.54068/1 (KT) 2,023,392 569,901 0.720520 (iii) 0.000012 73 120 808 1.75 (0.7003) 20.56 15.68 38.63 19.25 15.62 37.99071/1 (AHT) 2,014,694 569,503 0.704156 (iv) 0.000004 32 564 645 0.16 0.7027 19.67 15.62 38.08073/3 (AHT) 2,071,278 570,107 0.704960 (iv) 0.000006 16 191 810 0.24 0.7022 20.66 15.60 39.03080/1 (KT) 2,070,278 557,300 0.703277 (iv) 0.000005 16 474 700 0.10 0.7023 17.67 15.48 37.16097/2 (KT) 2,071,002 557,289 0.703541 (iv) 0.000007 20 493 700 0.12 0.7023 17.54 15.46 37.08135/1 (KT) 2,071,665 558,364 0.718828 (iii) 0.000007 41 81 813 1.46 (0.7019) 22.86 15.77 40.06 18.76 15.54 37.62145/1 (RAT) 2,069,809 558,557 0.766229 (ii) 0.000008 199 137 969 4.21 0.7079 17.93 15.75 38.21150/1 (KT) 2,069,734 560,131 0.728135 (iii) 0.000008 74 84 794 2.53 (0.6994) 22.20 15.93 39.62 18.70 15.51 37.50177/1 (KT) 2,066,807 555,977 0.703731 (iv) 0.000005 19 398 700 0.14 0.7023 17.61 15.49 37.22179/1 (KT) 2,065,406 555,379 0.706250 (i) 0.000008 16 91 630 0.52 (0.7016) 18.57 15.51 38.11193/1 (RAT) 2,068,485 552,569 0.734824 (ii) 0.000010 149 126 969 3.42 (0.6875) 18.17 15.74 38.10 17.93 15.72 37.98195/1 (RAT) 2,068,795 551,062 0.775040 (iii) 0.000010 207 128 969 4.70 0.7099 17.69 15.69 37.56291/1 (AHT) 2,030,219 555,396 0.705851 (i) 0.000007 40 313 645 0.37 0.7025 21.85 15.76 40.19293/1 (AHT) 2,026,703 554,703 0.773486 (iii) 0.000016 70 30 645 6.88 0.7101 19.83 15.63 38.12 18.70 15.60 37.89313/1 (KT) 2,060,959 566,811 0.717842 (iii) 0.000050 100 183 799 1.59 (0.6998) 23.72 15.86 41.31 18.70 15.51 37.50324/1 (AHT) 2,072,464 562,128 0.704221 (iv) 0.000008 64 1173 700 0.16 0.7027 19.14 15.63 38.97332/2 (KT) 2,055,299 562,834 0.720599 (iii) 0.000009 48.4 80.3 630 1.74 0.7050 24.46 15.93 42.08352/1 (KT) 2,043,685 558,221 0.706009 (i) 0.000005 24 205 803 0.34 0.7022 34.68 16.57 49.82355/1 (KT) 2,045,247 555,984 0.704786 (iv) 0.000007 22 271 700 0.23 0.7025 19.84 15.65 38.27371/2 (RAT) 2,057,666 542,663 0.703541 (iv) 0.000007 110 2661 909 0.12 0.7020 21.43 15.69 43.90376/1 (KT) 2,052,350 551,311 0.714577 (iii) 0.000009 57 132 783 1.23 (0.7008) 21.47 15.61 40.10397/1 (KT) 2,039,502 540,961 0.706593 (i) 0.000008 131 489 630 0.77 (0.6997) 18.83 15.63 38.04 18.43 15.59 37.64458/1 (KT) 2,050,114 567,748 0.709899 (i) 0.000006 64 246 813 0.75 (0.7012) 22.78 15.82 40.59476/1 (KT) 2,038,166 569,992 0.752891 (iii) 0.000018 109 61 813 5.18 (0.6928) 22.62 15.80 40.32498/1 (AHT) 2,069,200 573,374 0.703291 (iv) 0.000006 14 747 630 0.05 0.7028 18.39 15.56 38.40500/1 (AHT) 2,051,490 578,206 0.703541 (iv) 0.000007 106 274 700 1.12 (0.6924) 19.12 15.61 37.88

Initial 87Sr/86Sr ratios in brackets: correction for in situ decay failed; the lower boundary of data acceptance is the assumed depleted mantle composition at 1 Ga of 87Sr/86Sr ∼0.7020 (e.g. Workman and Hart, 2005).(i) Juvenile intermediate group Sr < 0.71 > 0.705; Rb/Sr < 1 one sample 5; four of six samples: correction for in situ decay failed.(ii) Cratonic group Sr > 0.72; vaiably high Rb/Sr > 1.8; 4 of 7 samples: correction for in situ decay failed.(iii) Juvenile radiogenic Sr > 0.71, variably high Rb/Sr > 1; 6 of 10 samples: correction for in situ decay failed.(iv) Juvenile mantle group < 0.7050, 1 of 10 samples: correction for in situ decay failed (and this sample has elevated Rb/Sr.RAT, KT, AHT: are Rahaba-Absol Terrane, Kurmut Terrane and Abu Harik Terranes respectively.

a Rb and Sr were analyzed by XRF.b Ages in bold are our new age data and ages in plain format are infered based on field relationship and lithological similarities to the dated rocks.

D. Evuk et al. / Precambrian Research 240 (2014) 108– 125 115

Fig. 3. Major element composition of meta-granitoids. (a) Most rocks are granitic transitional to quartz-monzonitic; one rock is monzodiorite–monzonite (sample 371/2);TAS diagram after Middlemost (1985). (b) All rocks are calc-alkaline with exception of three tholeiittic samples; AFM diagram after Irvine and Baragar (1971). (c) Mostelements are well correlated with SiO2, for example Al2O3, indicting no anomalous rock compositions. (d) Al-saturation index (ASI) and aluminosity (Frost et al., 2001)i aylor (A cles) at n of th

5

ottr

ndicate metaluminous to peraluminous rocks. (e) K2O vs SiO2 after Pecerillo and Tbsol Terrane samples (unfilled circles), Kurmut Terrane samples (orange filled cir

he references to color in this figure legend, the reader is referred to the web versio

.2. Pb and Sr isotope composition

Lead and Sr isotope ratios of whole rock and potassic feldspar

f selected samples are in Table 2. The leached Pb isotope ratios ofhe feldspar are considered to be close to the initial values either atime of magmatic crystallization or metamorphic overprint. Theyeveal systematically less radiogenic Pb isotope ratios than the

1976); (f) MALI vs SiO2 after Frost et al. (2001). Symbols for sample plots: Rahaba-nd Abu Harik Terrane samples (Martian Green filled circles). (For interpretation ofe article.)

related whole rock (Table 2). Two groups of feldspar and wholerock data are clearly distinguished by their 207Pb/204Pb values(Fig. 4a). We also note that a number of whole rock Pb isotope

ratios plot in the vicinity of the respective feldspar data (Fig. 4a),indicating low U/Pb ratios in these samples and hence a retardedgrowth of Pb since the Neoproterozoic. The differences betweenthe two groups in 207Pb/204Pb values relate to the compositional

116 D. Evuk et al. / Precambrian Research 240 (2014) 108– 125

Fig. 4. Pb- and Sr-isotope ratios; uranogenic Pb isotope composition (a) of our samples compared to (b) cratonic and juvenile basement of NE Africa. Two groups aredistinguished, which represent cratonic additions (207Pb/204Pb > 15.7, 206Pb/204Pb < 20) and juvenile material derived from depleted mantle. (a) K-feldspar data (open squares)represent intital Pb isotope ratios and respective whole rocks (dots) are related to radiogenic Pb growth at different � (=238U/204Pb) and age (t) of crystallization or lateradjustment of �. The lines represent possible examples of influence of �- and t-variation on Pb isotopes. (b) Our samples with high 207Pb/204Pb plot above the evolution ofterrestrial Pb (S&K, Stacey and Kramers, 1975) in line with data of Pan-African magmatic and metamorphic rocks from Darfur (Davidson and Wilson, 1989; Lucassen et al.,2008a) and from the Bayuda Desert (Küster et al., 2008), which represent reworked craton. Samples below the S&K line represent material dominated by Pb from the depletedmantle (triangles: present day and initial compositions of ophiolite and magmatic rocks; Brueckner et al., 1988, 1995; Zimmer et al., 1995; Stern and Kröner, 1993; Stern andA ssen ea er, thei

dih2

(mj

t

bdelsalam, 1998); lower crustal xenoliths from juvenile Pan-African sources (Lucare annotated. Low 87Sr/86Sr restrict to samples with juvenile Pb signature, howevsotopes (Fig. 4). See text for discussion.

ifferences in the dominant Pb sources of the rocks. One sources cratonic basement with an old, Archaean to Palaeoproterozoic,igh-� component (� = 238U/204Pb), which results in elevated07Pb/204Pb values compared with average terrestrial evolutionStacey and Kramers, 1975), whereas the other source, the depleted

antle, has lower 207Pb/204Pb values and provides the melts foruvenile crust (Fig. 4b).

The Pb isotope evolution from Pan-African crystallizationo present depends on the initial 238U/204Pb in the rocks and

t al., 2011) plot also within this array. (c) 207Pb/204Pb vs 87Sr/86Sr. Sample numbers87Sr/86Sr show considerable overlap between the two groups distinguished by Pb

secondary mobility of U, which is mobile in aqueous fluids downto temperatures of weathering (for review e.g. Skirrow et al.,2009). The Pb isotope composition in whole rocks and in respec-tive feldspar–whole rock pairs is variable (Table 2 and Fig. 4a)and therefore the 238U/204Pb must have been already variable

during the crystallization stage of the rocks. The fitted � for somefeldspar–whole rock pairs requires mobility of U after the crystal-lization stage (Fig. 4a). High 206Pb/204Pb in some whole rocks, whichplot right of the main juvenile trend (Fig. 4a), are explained by

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D. Evuk et al. / Precambrian

ddition of U late in their history. It has to be emphasized, whateverhe secondary U mobility, the differences between the ‘cratonic’ andjuvenile’ Pb (Fig. 4a) remain because they are based on 207Pb/204Pb.herefore, the uranogenic Pb isotope composition is an excellentnd robust tracer to distinguish between juvenile additions andrust strongly influenced by the craton. The new data underlinehe validity of this method in the regional context (Fig. 4b).

The calculation of initial Sr isotope ratios from whole-rockata failed in many samples (Table 2) or remains ambiguous evenhere the numerical results appear to be feasible, and thereforee restrict to the interpretation of analyzed Sr isotope ratios. In

hese old rocks secondary Rb–Sr mobility, i.e. unknown originalb/Sr, and/or the lack of a precise age for each sample are the prin-ipal causes for the failure (e.g., Küster et al., 2008). However, low7Sr/86Sr in our samples occur exclusively with Pb isotope signa-ure of the juvenile rocks, i.e. are in line with this interpretationAppendix B Fig. 6). These samples have low Rb/Sr ratios and didot evolve to the high radiogenic 87Sr/86Sr. The very high radio-enic 87Sr/86Sr appear to be more frequent in granitoids related tohe cratonic source, but there is also considerable overlap betweenoth groups. The isotope signatures show no relation to the majorlement chemical composition because the rocks evolved not fromommon parent magma and the granitoids likely represent only amall window of the fractionation–contamination history.

.3. Geochronology

Most of the zircon ages are concordant in the Wetherill (1956)oncordia diagram (Fig. 5a–m; details in Appendix A Table 3). Theathodoluminscence (CL) images are presented in Appendix B Fig.(a–m). The Th/U ratios (Appendix B Table 3A–M) were plottedersus 207Pb/206Pb for the zircon population from each rock toiscriminate between magmatic and metamorphic crystallizationAppendix B Fig. 8a–m). In some samples, a few analyses show dis-ordant ages and/or not all concordant ages cluster around the sameean value (Fig. 5d, e, i and j). In these samples, we selected the

ges with a concordance range of 100 ± 5% and calculated weightedverage ages (Appendix B Table 3D, E, I, and J). Upper intercept agesre commonly close to the main cluster of concordant ages, whereasower intercepts are insignificant because they are zero and/or have

large error. The concordia ages as well as the calculated weightedverage concordant ages are interpreted as crystallization ages ofhe zircons either from a melt or during metamorphism, and thushe crystallization age of the respective rocks.

An average Th/U ratio of 0.47 (Ahrens, 1965) has been vieweds typical of magmatic zircon. However, recent studies on felsicgneous rocks show a large range between 0.1 and 1.0 (da Silvat al., 2000; Rubatto, 2002; Schersten et al., 2004). Magmatic zir-ons with Th/U ratios outside this range are also described (Pidgeonnd Compston, 1992; Williams, 2001), possibly because the initialelt also had an anomalous ratio (Grant et al., 2009), and there-

ore the distinction between magmatic and metamorphic must beiewed with care. Th/U values below 0.1 are common when theelt is generated by partial melting and anatexis, and are attributed

o a combination of Th depletion from coeval growth of monazitend U enrichment of the metamorphic fluids that accompany min-ral breakdown reactions (Rubatto et al., 2001; Williams, 2001).ith a few Th/U ratios that scatter outside the magmatic or meta-orphic range of a particular rock, the majority of the samples

how ratios consistent with the respective type either magmaticr metamorphic (Appendix B Fig. 8a–m).

.3.1. The Rahaba-Absol TerraneSamples from the Rahaba-Absol Terrane included porphy-

oblastic biotite metagranite (sample 193/1), porphyroblasticuartzfeldspathic metagranite (sample 015/1), porphyroblastic

rch 240 (2014) 108– 125 117

biotite–muscovite metagranite (sample 059/1) and a meta-monzodiorite (sample 371/2). The CL images of zircon grains fromporphyroblastic biotite metagranite (sample 193/1) show pris-matic elongated as well as stubby zircon grains in cross-section.Some zircons have oscillatory zoning while others have uniformhomogeneous internal texture but with a thin, zoned dark rim(Appendix B Fig. 7a). The zircon grains have a concordia age of969 ± 5 Ma (Fig. 5a; Appendix B Table 3A). One zircon grain hasa dark sub-rounded core of 1001 Ma (Appendix B Fig. 7a). It is over-grown by a brighter oscillatory zoning followed by a grey oscillatoryzoning with an age of 974 Ma. The Th/U ratio of the zircons has alarge variation between 0.00 and 0.95. The rock was not affected bylater Pan-African re-juvenation.

The CL images of zircon grains from a porphyroblasticquartzfeldspathic metagranite (sample 015/1), a porphyroblasticbiotite–muscovite metagranite (sample 059/1) and a meta-monzodiorite (sample 371/2) show elongated prismatic zirconswith prominent oscillatory zoning and some with sub-rounded toirregular and sub-hedral inherited cores (Appendix B Fig. 7b–d).The oscillatory zoned grains from sample 015/1 gave a concordiaage of 914 ± 6 Ma (Fig. 5b), interpreted as magmatic because theTh/U ratios are greater than 0.10 (Appendix B Fig. 7b). Cores of zir-cons yielded a Paleoproterozoic age of 1656 Ma and 1928 Ma. Onezircon grain has a Mesoproterozoic rim (1155 Ma) around a younger(probably resetted metamict) core of 926 Ma. Some ages of 842, 843and 851 Ma (Appendix B Table 3B) indicate Pan-African resetting.

Zircon grains from a porphyroblastic biotite–muscovite meta-granite (sample 059/1) have cores and with rims overgrown insimilar orientation. The rims gave a concordia age of 912 ± 4 Ma(Fig. 5c; Appendix B Table 3c), considered as the age of metamor-phism due to the low content of Th/U ratios (Appendix B Fig. 8c) forthe spots around 900 Ma, and a large scatter for the inherited ones,which have Paleoproterozoic ages of 1617–2281 Ma and Mesopro-terozoic ages of 1007–1434 Ma. There is only one zircon with a rimof 852 Ma, indicating minor Pan-African resetting (Fig. 5c).

The meta-monzodiorite (sample 371/2) zircon grains haveinherited cores with different textures and age groups. A groupof inherited bright to light grey cores is irregular to sub-roundedwith Paleoproterozoic ages of 2080–1726 Ma, another group is pris-matic with prominent oscillatory zoning and Mesoproterozoic agesof 1169–1054 Ma. Based on the Th/U ratio, all were probably frommagmatic rocks (Appendix B Figs. 7d and 8d). The other concordantage groups have weighted averages at 909 ± 9, 818 ± 19, 755 ± 15and 669 ± 13 Ma (Fig. 5d; Appendix B Table 3D) considered as meta-morphic ages because of their low Th/U ratios.

5.3.2. Abu Harik-Kurmut TerranesSamples from the Abu Harik-Kurmut Terranes included three

biotite metagranites (Table 1; samples 150/1, 324/1 and 376/1),three medium-grained metagranites (samples 068/1, 135/1 and313/1), an alkali metagranite (sample 293/1), a meta-quartz-monzonite and a porphyritic meta-quartz-monzonite (073/3 and397/1 respectively).

CL images of zircon grains from sample 150/1 reveals some dark,high-U cores and some bright rims indicative of complex dissolv-ing and redeposition of U (Appendix B Fig. 7e). There are brightmarginal bands and amoebic dark zones entirely enclosed in thegrains. According to Corfu et al. (2003) such textures occur in high-grade metamorphic rocks. Hoskin and Black (2000) interpretedsuch textures as due to solid-state recrystallization. Some zircongrains have regular magmatic oscillatory zoning while others havea homogeneous dark grey core with or without darker rims encom-

passing the brighter areas. A weighted average of concordant agesis 794 ± 15 Ma (Appendix B Table 3E) and interpreted as the mag-matic crystallization age for the zircons (Th/U > 0.10; Appendix BFig. 8e). Some zircon grains have inherited Early Neoproterozoic

118 D. Evuk et al. / Precambrian Research 240 (2014) 108– 125

Fig. 5. (a)–(m) Concordia and calculated weighted average of concordant ages. Data-point error ellipses are 2�. The 2� and decay constant errors ignored. The respectiverock units and sample numbers are inserted in the respective figures. Th/U ratios for discordant and concordant ages used for calculating weighted average ages are plottedagainst each ellipse. Red ellipses are concordant or near concordant ages; blue ellipses represent Pb loss from cores in Neoproterozoic (a); blue broken ellipses are inheritedMeso to Paleoproterozoic cores (b) and broken red ellipses are inherited Neoproterozoic cores (e). (For interpretation of the references to color in this figure legend, thereader is referred to the web version of the article.)

D. Evuk et al. / Precambrian Research 240 (2014) 108– 125 119

Fig. 5. (Continued .)

1 Rese

ce7T

tindgcsCib(eFmp

3ccaTtg(t

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20 D. Evuk et al. / Precambrian

ores of 895, 917, 946 and 999 Ma. The age of the metamorphicffect on the zircon grains is uncertain, because ages range from62 to 606 Ma. However, most of the ages in these ranges have lowh/U ratios indicating dominance of metamorphic resetting.

CL images of zircons from two samples (068/1 and 135/1) fromhe medium-grained metagranites show magmatic oscillatory zon-ng texture that are sometimes fractured, dissolved and filled withew material. Zircon grains from sample 135/1 gave a concor-ia age at 813 ± 4 Ma (Fig. 5f; Appendix B Table 3F,). Some zirconrains have relict ages of 840–877 Ma. Sample 068/1 yielded aoncordia age at 808 ± 5 Ma (Fig. 5h; Appendix B Table 3H) withome inherited ages between 840 and 873 Ma (Appendix B Fig. 7h).L images of zircons from the third medium-grained metagran-

te (sample 313/1) have magmatic oscillatory zoning with someroad homogeneous cores. A weighted average of concordant agesFig. 5i; Appendix B Table 3I) of 799 ± 16 Ma for the rock is consid-red as magmatic emplacement age (high Th/U ratios, Appendix Big. 8i). These three concordant ages are interpreted as the age ofagmatism. However, the textures of the zircons indicate contem-

oraneous hydrothermal recrystallization.CL images for zircons from of biotite metagranite (two samples

76/1 and 324/1) show homogeneous zircons with broad light-greyores with some irregular dark rims. There is faint zoning. The zir-on grains of sample 376/1 have weighted averages of concordantges (Fig. 5j; Appendix B Table 3J) at 783 ± 13 Ma and 714 ± 16 Ma.hese ages are interpreted as magmatic and metamorphic respec-ively (high and low Th/U ratios, Appendix B Table 3J). The zirconrains from sample 324/1 yielded a concordia age of 700 ± 7 MaFig. 5k; Appendix B Table 3K). We interpret the concordia age ashe magmatic age of the biotite metagranite (Th/U > 0.10).

Zircon grains of a meta-quartz-monzonite (sample 073/3) havescillatory zoning (Appendix B Fig. 7g). The grains have a con-ordia age at 810 ± 10 Ma (Fig. 5g; Appendix B Table 3G). Zirconrains from a porphyritic meta-quartz-monzonite (sample 397/1),lso with magmatic oscillatory zoning (Appendix B Fig. 7m) gave

concordia age at 630 ± 4 Ma (Fig. 5m; Appendix B Table 3M).e interpret both ages as magmatic emplacement ages (Th/U

atios > 0.10, Appendix B Fig. 8g, m).An alkali metagranite (sample 293/1) has equant stubby zircon

rains mostly with no or faint broad oscillatory zoning (Appendix Big. 7l). The zircon grains gave a concordia age at 645 ± 5 Ma (Fig. 5l;ppendix B Table 3L). The Th/U ratios show one cluster <0.10 and

arge variation of ratios >0.10 (Appendix B Fig. 8l). This age is con-idered as the age of anatexis of the Rahaba-Absol Terrane (Series)ock. The initial 87Sr/86Sr ratio of the alkali metagranite is 0.7101ndicating a remobilized cratonic source signature.

. Discussion

The time frame of magmatic and metamorphic activity at theentral-eastern Bayuda Desert is indicated by the concordant ages

rom this study and published evidence (Fig. 6). The Bayudian event,s the latest Pre-Pan-African activity, was placed by Küster et al.2008) in a short time span between 920 and 900 Ma bracketed by

etamorphism of the El Melagi gneiss at 921 ± 10 Ma and the intru-ion of the Absol-type granodioritic pluton at 900 ± 9 Ma (Fig. 2).hey reported inherited ages of 1.0, 2.0, 2.5 and 2.7 Ga from the Elelagi gneiss in the Bayuda Desert west of our area of study. Our

ge data for magmatism and metamorphism at 969 ± 5 Ma indicatehat the Bayuda Event began earlier, but confirm the peak activityt 920 to 900 Ma with the ages of 914 and 912 Ma for metagran-

tes. Relict ages in these zircon grains go back to 2.28 Ga, indicatinglder rocks in the sources for the granitoids. Samarium–Nd modelges of separation from the depleted mantle of 2.1 to 2.2 Ga foretasediments of the Rahaba-Absol Terrane (Küster and Liegois,

arch 240 (2014) 108– 125

2001) and 2.8 to 1.3 Ga for rocks from the Halfa Terrane further tothe north-west of our study area (Harms et al., 1990, 1994; Sternet al., 1994) are consistent with the inherited zircons. Shang et al.(2012) even reported an Archaean zircon xenolith age of 3.0 Ga froma tonalite-trondhjemite-granodiorite north of Delgo in the Halfaterrane (Fig. 1). Such old Proterozoic and Neoproterozoic ages havenever been observed in the Nubian Shield, and clearly confirm thatthe Rahaba-Absol-Terrane is part of the Saharan Metacraton.

The old age of the Rahaba-Absol Terrane is completely in linewith the Pb-isotope data (Fig. 4a, b). Samples with a cratonic sig-nature occur exclusively in the Rahaba-Absol Terrane. Our sample145/1, a porphyroblastic metagranite, was taken just between twosample points, which yielded younger zircon ages and also juve-nile Pb-signatures (Fig. 2). It is interpreted as a rock slice within theanticlinal structure of the Azuma interference fold. This is likelya small window of Rahaba-Absol Series that was overlane by AbuHarik-Kurmut Series, then exposed by erosion. Our samples withhigh 207Pb/204Pb plot above the evolution of terrestrial Pb (Fig. 4a,b; Stacey and Kramers, 1975) in line with data of Pan-African mag-matic and metamorphic rocks from Darfur (Davidson and Wilson,1989; Lucassen et al., 2008a) and from the Bayuda Desert (Küsteret al., 2008), which represent reworked craton. Samples, whichrepresent material dominated by Pb from the depleted mantle istypically juvenile (values for present day and initial compositionsof ophiolite and magmatic rocks from the area in Brueckner et al.,1988, 1995; Zimmer et al., 1995, Stern and Kröner, 1993; Sternand Abdelsalam, 1998). Lower crustal xenoliths from the CenozoicBayuda Volcanic Field attributed to juvenile Pan-African sources(Lucassen et al., 2011) plot also within this array.

Age resetting is recorded in the zircon population from themeta-monzodiorite (371/2) as weighted averages of concordantage clusters at 818 ± 19, 755 ± 15 and 669 ± 13 Ma. The first twoages are interpreted as newly grown metamorphic zircons or over-growths on inherited igneous crystals. The later one is mainlyovergrowth rim representing the last metamorphic overprinting ofthe zircons (Appendix B Fig. 7d). Corresponding ages are frequentlyobserved in the Dam El Tor FTB and other parts of the KurmutTerrane. There, the magmatic and metamorphic cycles are docu-mented by several ages between 813 ± 4 and 630 ± 4 Ma and suchages have been obtained previously from the Nubian Shield (Fig. 1b:Gebeit-Gabgaba area; Stern and Kröner, 1993; Stern et al., 1989,Klemenic, 1985; Klemenic and Poole, 1988; the Nakasib Suture;Abdelsalam and Stern, 1993a,b; Johnson et al., 2003; Stern andKröner, 1993; Stern and Abdelsalam, 1998; and the Haya area;Kröner et al., 1991; Reischmann et al., 1992; Abdelsalam and Stern,1993a,b, 1996c), but not from the western part. They confirm thatthe Kurmut Terrane and the Dam El Tor FTB are part of the amal-gamation of the Nubian Shield. Küster and Liégois (2001) and Küsteret al. (2008) presented a Sm/Nd isochrone age from amphibolites of806 ± 19 Ma for metamorphism with an intrusion age of 858 ± 7 Ma(Table 1). Our ages of 810 ± 10, 794 ± 15, 808 ± 5, 783 ± 13, 700 ± 7and 630 ± 4 Ma are interpreted as intrusion ages, the other ages813 ± 4, 714 ± 16 and 669 ± 13 Ma as ages of metamorphism or645 ± 5 Ma as the age of anatexis of the Rahaba-Absol Terrane rocksbecause of high intitial Sr ratio (Table 2) for the alkali metagran-ite indicating cratonal contamination or partial melting. From thefrequency distribution (Fig. 6) it is obvious that the peak of themagmatic–metamorphic activity is around 820–780 Ma. The sameage group is very prominent in a statistical study of zircon ages frommodern stream sediments (Iizuka et al., 2012), especially from theNile.

Our samples from the Kurmut-Abu Harik Terrane show a juve-

nile Pb-isotope signature, typical for the Nubian Shield (Fig. 4). Thisis in line with juvenile, mantle-derived characteristics of publishedO- and Hf-isotope characteristics of zircons from this age group(Iizuka et al., 2012). The Rb–Sr isotope system of rocks from the

D. Evuk et al. / Precambrian Research 240 (2014) 108– 125 121

vents

Bpmc2paiSP

snrarS

Fig. 6. Time-frame for geological e

ayuda Desert has long been recognized to be disturbed mainly byost-crystallization Rb mobility, which results in unreliable, com-only overcorrected initial ratios, even in samples with known

rystallization ages (Barth and Meinhold, 1979; Küster and Liégeois,001; Küster et al., 2008). Nevertheless, samples with low Rb/Srreserve near-initial Sr isotope ratios, which allow distinguishing

juvenile Pan-African mantle signature from cratonic sources asn other areas (e.g. Harms et al., 1994; Stern and Kröner, 1993;tern, 2008) and correspond in our sample set with the respectiveb isotope signatures (Fig. 4a, b)

In the time between 780 and 720 Ma, the frequency distributionhows few data, indicating a period of quiescence. However, to theorth of our study area in the Halfa Terrane Shang et al. (2012)

eported an important U–Pb zircon age group of 728–702 Ma,nd also north-east from the Gabgaba area such ages have beeneported in the Allagi-Heiana suture and the Onib-Sol Hameduture (Fig. 1). The final magmatic and metamorphic activity

at central-eastern Bayuda Desert.

peaked between 720 and 630 Ma. Such ages are known in theBayuda Desert since Vail (1971), and can be considered as Pan-African in the strict sense. They are also commonly observed inmany parts of the Saharan Metacraton as well as in the Arabian-Nubian Shield (see Fig. 1).

The crustal cooling in the central-eastern Bayuda Desert likelycontinued from 590 to 550 Ma (Table 1 and Fig. 6). Because alsothe youngest dated rock at 630 ± 4 Ma from our study is a meta-granitoid, there must have been metamorphism ≤630 Ma, butthe post-collisional granitoids reported by Küster at al. (2008) at597 ± 4 Ma are un-deformed plutons. So the latest metamorphism-deformation occurred in this time frame. High heat flow duringpost-collisional extension probably resulted in weak metamor-

phism of sediments in pull-apart basins of the Amaki Series (ElRaaba, 1976). Such basins are believed to have formed between620 and 540 Ma in the Arabian-Nubian Shield (Johnson et al.,2011).

1 Rese

7

(GcAfharcidatcmb(9fmr

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22 D. Evuk et al. / Precambrian

. Geodynamic evolution

In the Late Neoproterozoic during the East African OrogenyStern, 1994; Meert, 2003) the collision between East and Westondwana sandwiched the Arabian-Nubian Shield in between oldratonic blocks. The collision between the western boundary of therabian-Nubian Shield and the Saharan Metacraton (the eastern

oreland of West Gondwana) at the Bayuda Desert is believed toave formed the Keraf Zone, a continental-arc suture (Abdelsalamnd Stern, 1996a). All our dated rocks from the Rahaba-Absol Ter-ane are calc-alkaline peraluminous (Appendix B Fig. 4a, b) inomposition similar to some of the Pan-African island-arc rocksn the central-eastern Bayuda Desert. It is likely that the Bayu-ian Event spans a period from 1000 to 900 Ma with metamorphicnd magmatic events (Fig. 6). It is also probable that extensionalectonics played a greater role during this 100 Ma period in theentral-eastern Bayuda Desert, which led to high-heat flow, mag-atism and metamorphism, and extensive anatexis was reported

y Barth and Meinhold (1979) from the Rahaba Series gneissesour porphyroblastic metagranitoids). Our dated samples with ages69 ± 5, 914.1 ± 5.5, 912 ± 4 and 909 ± 9 Ma with large potassiceldspar porphyroblast represent this period of magmatism and

etamorphism, which probably affected most of the Bayuda Ter-ane.

Likewise, the metagranitoids from the Rahaba-Absol Terraneave inherited cores of detrital zircon grains spanning 2.28 to 1 Gaeriod with their intrusion and metamorphic ages at 969 ± 5 Ma,14.1 ± 5.5 Ma, 912 ± 4 Ma and 909 ± 9 Ma. This implies extensionas probably active prior to the intrusion and metamorphism

f the metagranitoids allowing for sediment transport with thessociated zircon grains. Moreover, the metagranitoids are dom-nantly peraluminous S-type granites (Fig. 3d; appendix B Fig.a–i). Their calc-alkaline nature and island-arc setting supportsormation in a collisional orogen (Appendix B Fig. 4a, b) at thend of the 1000–900 Ma period with probable synchronous exten-ion that can permit the ascendance of aesthenospheric mantleelow the lithosphere leading to high-heat flow, magmatism andetamorphism. Further work is required to further understand

n detail how the Bayuda Terrane behaved during the Rodiniassembly, which is beyond the scope of this research; nevertheless,ucassen et al. (2008) concluded from isotopic studies of mantleenoliths from the Bayuda Desert that the sub-continental litho-pheric mantle of the Saharan Metacraton is an old depleted mantlehich was formed and separated from the convective mantle

efore the Neoproterozoic Pan-African event ∼1 Ga ago. The pro-ess of mantle-lithosphere decoupling and or mantle lithosphereelamination would call for extension. In addition, delamination

s an oft-cited mechanism that allows decompression meltingf upwelling asthenosphere and consequent voluminous pluton-sm, which consequently results in crustal uplift, extension andollapse (Gerbi, 2002). Such mechanisms might have operated dur-ng the Bayuda Event at least during its ending period. Recent

ork on detrital zircons from part of the Arabian-Nubian Shieldnd Saharan Metacraton (Morag et al., 2011; Meinhold et al.,012 and Iizuka et al., 2012) favours an existence of a 1000 to00 Ma Terrane with some island-arc setting in areas occupied byorthern ANS or its western and eastern boundaries. Far-reachingtresses in large orogenic areas can induce some extension in colli-ional tectonics where transpressional movements are replaced byranstensional movements. Of course the assembly of Rodinia goesack to 1300 Ma and the 1000 to 900 Ma period represents onlyhe final stages in a broader framework of its orogeny. For example

etasedimentary rocks of the Peloritani Mountains of NE Sicily areharacterized by the presence of a distinct 1000 to 900 Ma detri-al inherited zircon component with a proposed East African originWilliams et al., 2012). It is likely that the central-eastern Bayuda

arch 240 (2014) 108– 125

Desert magmatism and metamorphism of the metagranitoids wasthe result of a regional event.

Moreover, the Earth has witnessed worldwide orogenic eventsbetween 1300 Ma and 900 Ma (Li et al., 2008). In reconstructing thebuilding blocks that formed Rodinia (Li et al., 2008), from 1100 Mato 1000 Ma the Sahara was always west of the Congo-São Franciscocraton and got welded to the position of the Arabian-Nubian Shieldat 900 Ma. The Bayuda area was situated NNW of the position of theSahara. It is not clear how the whole crust of the Saharan Metacra-ton behaved during the complete process of Rodinia assembly.However, Liégeois et al. (2013) concluded that at the end of theNeoproterozoic the pre-existing Saharan craton was involved incollisional events along all its margins, against the Tuareg Shieldbacked by the West African craton in the west, against the Congocraton and intervening Pan-African belts to the south, against theArabian-Nubian Shield to the east and against an unknown conti-nent to the north. With a prevented escape to all sides, the Saharancraton was metacratonized on its margins as well as in its interior asdescribed for the Djanet-Edembo terranes in Eastern Hoggar duringthe Murzukian episode (Fezaa et al., 2010). They further proposedthat the Murzug, Al Kufrah and Chad cratons are remnants of thepre-Neoproterozoic Saharan craton (Fig. 1).

The metagranitoids situated either along or south of Dam ElTor FTB are associated with high-grade metasedimentary rocks andgreenschist rocks of the Arabian-Nubian Shield of the Abu Harik-Kurmut Terranes. They vary from calc-alkaline to high potassiccalc-alkaline series typical for an evolving island-arc environmentto a mature arc. The frequency distribution of the concordantage data is high between 850 and 775 Ma and moderate between700 Ma and 625 Ma (Fig. 6). We interpret the first period as activ-ity related to magmatism and metamorphism in a subduction zone,and the second as related to the collision between the Rahaba-AbsolTerrane and the Abu Harik and Kurmut Terranes.

When the ages of the Abu Harik-Kurmut Terranes are corre-lated with the eastern Gabgaba-Gebeit Terranes (832–700 Ma) ofthe Red Sea Hills of Sudan and the Haya Terrane to the south-east(900 to 800 Ma), it is likely the Haya Terrane extended to the AbuHarik-Kurmut Terrane and together formed the foreland whereoceanic crust was subducted along the zone of Nakashib-Fadlab-Dam El Tor. The subduction related activity at the Rahaba-Absoland the Abu Harik-Kurmut Terranes boundary probably started ear-lier and reached its maximum at ∼800 Ma and collision started at∼714 Ma ending at ∼630 Ma (Fig. 6). Reduction in magmatic andor metamorphic activity between 800 and 700 Ma could be relatedto a changing subduction rate or shifting of subduction or collisionsomewhere else in the Arabian-Nubian Shield, e.g. the NakashibSuture (800–750 Ma) or Atmur-Delgo Suture (750–718 Ma; seeFig. 1b).

Bailo (2000) studied the central part of the Keraf Shear Zone, andhe concluded that high-grade metamorphism occurred at ∼730 Mawhile post-tectonic plutons were emplaced at 710 Ma with the lastthermal event at 560 Ma. Accordingly, Bailo (2000) envisaged thatthe horizontal movement in the central Keraf Shear Zone occurredprior to 710 Ma (the age of a discordant pluton). Abdelsalam andStern (1996c) constrained the sinistral movement at the KerafShear Zone between 700 and 610 Ma. Deformation along the adja-cent Hamisana Shear Zone has been constrained to 660–610 Ma(Abdelsalam and Stern, 1996c). The age of our biotite metagranitethat has been folded along NE trending structure, then truncatedby the NS trending Keraf structures, is 700 ± 7 Ma. Both our datedalkali metagranite and the porphyritic meta-quartz-monzonite arediscordant to multiply folded rocks of the Kurmut Terrane and have

ages of 645 ± 5 Ma and 630 ± 4 Ma respectively. The alkali meta-granite trends E-W at high angle to the NS trend of southern Kerafstructures implying formation under a NS stress regime, while theporphyritic meta-quartz-monzonite trends NE parallel to Dam El

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or FTB, but shows no over-printing of the Keraf structures. It wasikely forcefully emplaced in extensional NE trending structuresblique to the NW-SW shortening. NE-trending structures also con-inue east of southern Keraf Shear Zone similar to Dam El Tor FTBrientation. Therefore, we think that the NS movement, mostlynterpreted as post-collisional horizontal movement along south-rn Keraf Shear Zone, is probably restricted to a narrow zone andccurred between 630 and 590 Ma.

Barth and Meinhold (1979), and Vail (1971) obtained agesy the Rb/Sr method from granite and pegmatite of the Rahaband Absol Series at 645 to 635 Ma. These ages likely representigmatization related to the terminal collision. Rb/Sr-ages from

he post-collisional high-K granites obtained gave 597 to 555 MaBarth et al., 1983; Küster et al., 2008; Meinhold, 1979; Vail, 1971).he period between 605 and 591 Ma has been interpreted by Küstert al. (2008) as the emplacement time for post-collisional granitoidsontemporaneous with late Pan-African metamorphism relatedo post-collisional crustal extension and uplift. A thickened crustrobably existed in the central-eastern Bayuda Desert from 630o 600 Ma after terminal collision between Rahaba-Absol Terranend the Abu Harik-Kurmut Terranes. This gives a ∼30 Ma periodf erosion and denudation of the stacked crustal material beforehe beginning of post-collisional magmatism that was enhancedy horizontal movements a long shear zones.

. Summary

Petrographic, geochemical, isotopic and geochronological studyf meta-igneous rocks (dominantly metagranitoids) revealedxtended magmatic and metamorphic activities in the central-astern Bayuda Desert during the Neoproterozoic between ∼1000nd 630 Ma (Fig. 6).

1) From current age data, a clear age difference between Tonianof the continental Rahaba-Absol Terrane and the Cryogenian ofthe Abu Harik-Kurmut arc terrane exists.

2) The calc-alkaline nature of the porphyroblastic metagranitoidsand the meta-monzodiorite from the Rahaba-Absol Terrane butlacking associated rocks typical for island-arc tectonic settingsuch as ophiolites allows us to propose the hypothesis that theporphyroblastic metagranitoids and the meta-monzodioriteof the Rahaba-Absol Terrane were formed in a continentalarc setting in an extended Bayudian Event between 1000and 900 Ma probably beginning with extension as early as1100 Ma related to sub-continental mantle lithospheric delam-ination and decoupling processes. Sedimentary activity likelyprevailed up to early Neoproterozoic bringing in older zir-con grains. Metamorphism and magmatism occured during theperiod 969 ± 5 Ma and intensified at 914 ± 6 Ma, 912 ± 4 and909 ± 9 Ma. The period at 920–900 Ma (Küster et al., 2008) prob-ably represents the last stage in the heating of the crust byupwelling hot asthenosphere. However, further work on thepressure–temperature–time path is needed to clearly under-stand the type of metamorphism in the early Neoproterozoic.

3) The extension due to the break-up of Rodinia and the beginningof the Pan-African cycle started probably as early as 870 Ma.Subduction related magmatism and metamophism reached itsmaximum between 825 and 800 Ma. The subduction relatedactivity was reduced in intensity between 775 and ∼725 Ma.

4) Collision between Rahaba-Absol and the Abu Harik-KurmutTerranes related to the Pan-African event began around

714 ± 16 Ma and ended at ∼640–630 Ma, the intrusion ages ofE-W trending alkali metagranite (645 ± 5 Ma) and a porphyriticmeta-quartz–monzonite (630 ± 4 Ma) both oblique to the NSshearing trend and NW-SE shortening direction. It is likely

rch 240 (2014) 108– 125 123

that horizontal movement in the southern Keraf Shear Zoneat central-eastern Bayuda Desert occurred between 630 and590 Ma after the terminal collision.

(5) A quiescence period between 630 and 600 Ma existed whereerosion probably dominated because of a thickened crust aftercollision. This period was followed by alkaline magmatismbetween 605 and 591 Ma and later cooling that continued upto 555 Ma.

(6) Pb isotope data indicate two Pb sources of the rocks relatedto the remobilized craton and one related to the juvenile pan-African crust. The Pb isotope confirms the boundary betweenthe Rahaba-Absol Terrane and the Kurmuk Terrane at the north-ern boundary of Dam El Tor FTB. To delineate this boundaryfurther north and south this method seems very promising.

(7) Assuming an initial Sr ratio for juvenile Neoproterozoic mantlemelt of 0.703, we can generally conclude that some of the meta-granitoids are mantle melts with no or little influence of theolder cratonic material while others are either contaminatedor are partial melt products of the remobilized older cratoniccrust.

Acknowledgments

This work is part of the first author’s PhD project at the Techni-cal University of Berlin under the scholarship of DAAD. Therefore,D.E. is grateful to the DAAD for such an opportunity. We are veryobliged to Juba University for sponsoring the field work and to theGeological Research Authority of Sudan for providing the field worklogistics. Many thanks to Dr. Glodny, Dr. Gleißner, Miss. Zecha andMarsiske for help with sample preparations.

Appendix A. Supplementary data

Supplementary material related to this article can be found,in the online version, at http://dx.doi.org/10.1016/j.precamres.2013.10.015.

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