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1 Gondwana Research 52 (2017) Contents lists available at ScienceDirect Gondwana Research journal homepage: Depositional ages and provenance of the Neoproterozoic Damara Supergroup (northwest Namibia): Implications for the Angola-Congo and Kalahari cratons connection Débora B. Nascimento a,b,,renatas.schmitt b, André Ribeiro b, Rudolph A.J. Trouw b, Cees W. Passchier c,miguela.s.basei d a Programa de Pós graduação em Geologia, Universidade Federal do Rio de Janeiro (PPGL/UFRJ), CEP Rio de Janeiro, Brazil b Instituto de Geociências, Universidade Federal do Rio de Janeiro (IGEO/UFRJ), CEP Rio de Janeiro, Brazil c Institut für Geowissenschaften, Johannes Gutenberg-Universität Mainz, Mainz, Germany d Instituto de Geociências, Universidade de São Paulo, CEP São Paulo, Brazil article info abstract Article history: Received 25 November 2016 Received in revised form 4 July 2017 Accepted 24 September 2017 Available online 27 September 2017 Handling Editor: A.S. Collins Keywords: Neoproterozoic Damara Supergroup U-Pb geochronology Outjo Basin evolution Rodinia The Damara Orogen is composed of the Damara, Kaoko and Gariep belts developed during the Neoproterozoic Pan-African Orogeny. The Damara Belt contains Neoproterozoic siliciclastic and carbonate successions of the Damara Supergroup that record rift to proto-ocean depositional phases during the Rodinia supercontinent break up. There are two conflicting interpretations of the geotectonic framework of the Damara Supergroup basin: i) as one major basin, composed of the Outjo and Khomas basins, related to rifting in the Angola-Congo- Kalahari paleocontinent or, ii) as two independent passive margin basins, one related to the Angola-Congo and the other to the Kalahari proto-cratons. Detrital zircon provenance studies linked to field geology were used to solve this controversy. U-Pb zircon age data were analyzed in order to characterize depositional ages and provenance of the sediments and evolution of the succession in the northern part of the Outjo Basin. The basal Nabis Formation (Nosib Group) and the base of the Chuos Formation were deposited between ca. 870 Ma and 760 Ma. The upper Chuos, Berg Aukas, Gauss, Auros and lower Brak River formations formed between ca. 760 Ma and 635 Ma. It also includes the time span recorded by the unconformity between the Auros and lower Brak River formations. The Ghaub, upper Brak River, Karibib and Kuiseb formations were deposited between 663 Ma and 590 Ma. The geochronological data indicate that the main source areas are related to: i) the Angola-Congo Craton, ii) rift-related intrabasinal igneous rocks of the Naauwpoort Formation, iii) an intrabasinal basement structural high (Abbabis High), and iv) the Coastal Terrane of the Kaoko Belt. The Kalahari Craton units apparently did not constitute a main source area for the studied succession. This is possibly due to the position of the succession in the northern part of the Outjo Basin, at the southern margin of the Congo Craton. Comparison of the obtained geochronological data with those from the literature shows that the Abbabis High forms part of the Kalahari proto-craton and that Angola-Congo and Kalahari cratons were part of the same paleocontinent in Rodinia times International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved. 1. Introduction The Damara Supergroup is a thick Neoproterozoic siliciclastic and carbonate succession related to the Angola-Congo and Kalahari cratons (e.g., Porada, 1979, 1989; Miller, 1983, 2008). The succession constitutes the Pan-African Kaoko and Damara belts (Fig. 1A; e.g., Hanson, 2003). In Damaraland, northwest Namibia, the Damara succession encompasses deposits that have been traditionally interpreted as related to rift- Corresponding author at: Programa de Pós graduação em Geologia, Universidade Federal do Rio de Janeiro (PPGL/UFRJ), CEP Rio de Janeiro, Brazil. address: debora@geologia.ufrj.br (D.B. Nascimento). (Nosib Group), rift to passive margin (Otavi/Swakop groups) and foreland basins between the Angola-Congo and Kalahari cratons (Mulden Group; Miller, 1983, 2008; Porada, 1983; Paciullo et al., 2007; Miller et al., 2009a, 2009b; Nascimento et al., 2016). However, the limited number of isotopic data hampers interpretations of depositional ages and provenance of the deposits. It is equally uncertain whether the rift to passive margin succession is related to fragmentation of a single continental block or to two distinct paleocontinents. There are two main hypotheses about location of the Angola-Congo and Kalahari protocratons during Rodinia times and, consequently, different propositions for the evolution of the Neoproterozoic basins related to these cratons. The first hypothesis considers the proto-cratons connected during X/ 2017 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved.

2 154 D.B. Nascimento et al. / Gondwana Research 52 (2017) Fig. 1. A) NE-trending Damara, Lufilian and Zambezi belts separating the Angola-Congo and Kalahari cratons, and N-S-trending West Congo, Kaoko, Gariep and Saldania belts in west, and Mozambique Belt in east Africa. Location of the Irumide Belt in the Congo Craton and the Rehoboth Inlier, in the Kalahari Craton. Simplified from Hanson (2003). B)Simplified tectonic map of the Damara and Kaoko belts showing the distribution of the tectonostratigraphic zones of Miller (2008), the main lineaments and the location of the Outjo and Khomas seas. Location of samples analyzed by Foster et al. (2015) is indicated. Box - study area shown on Fig. 2. Rodinia times, generating an intracratonic extension evolving to a proto-oceanic, Red Sea-like basin, during rifting (e.g., Martin and Porada, 1977; Miller, 1983, 2008; Hanson, 2003). The other hypothesis considers the Angola-Congo and Kalahari proto-cratons as distinct continental masses that might not have been part of the Rodinia supercontinent (e.g., Burke et al., 1977, 2003; Stanistreet et al., 1991; John et al., 2003; Tohver et al., 2006; Gray et al., 2008). U-Pb analyses combined with detailed field geology provided accurate data on depositional age and provenance of these clastic successions. They also allow interpretation of the tectonic setting, framework and evolution of depositional basins (e.g., Cawood et al., 2012). We present 569 concordant U-Pb detrital zircon ages constraining the depositional ages and source areas of six formations of the Damara succession in Damaraland, northwestern Namibia. The geochronological and field data together with data from the literature constrain the position of the proto-cratons and the geotectonic setting of the Damara basin succession. 2. Geological setting The assemblage of the Gondwana supercontinent occurred during the Pan-African/Brasiliano orogenic events which generated several mobile belts in Africa and South America (e.g., Porada, 1979, 1989; Barnes and Sawyer, 1980; Miller, 1983; Kukla and Stanistreet, 1991; Prave, 1996; Goscombe et al., 2003a, 2003b; Hanson, 2003; Frimmel, 2009). In southern Africa, these Neoproterozoic orogenic belts constitute segments that surround and separate the Angola-Congo and Kalahari cratons (Fig. 1A). The N-S-trending belts are the Mozambique Belt in eastern Africa (Fig. 1A; e.g., Grantham et al., 2003, Hanson, 2003), and, from north to south, the West Congo, Kaoko, Gariep and Saldania belts (Fig. 1A; e.g., Miller, 1983; Porada, 1989; Hanson, 2003) in southwestern Africa. The NE-trending Damara, Lufilian and Zambezi belts separate the Angola-Congo and Kalahari cratons (Fig. 1A; e.g., Miller, 1983; Hanson, 2003). Although Phanerozoic cover prevents direct verification, the Damara Belt is envisaged as linked to the Lufilian and Zambezi belts (e.g., Goscombe et al., 2000; Hanson, 2003). The Kaoko and Gariep belts are considered to have been connected, before Gondwana break up (b200 Ma; e.g., Trompette, 1994; Hanson, 2003), to the Saldania Belt in southern Africa, and to the Dom Feliciano and Ribeira belts in South America (e.g., Frimmel et al., 2013). The intracontinental Damara Belt, where the study area is located, and the coastal Kaoko and Gariep belts (Fig. 1A), crop out mainly in Namibia and together constitute the Damara Orogen (Porada, 1979;

3 D.B. Nascimento et al. / Gondwana Research 52 (2017) Miller, 1983; Prave, 1996). These belts record the closure of an ocean/sea in a collisional triple junction (e.g., Porada, 1979, 1989; Coward, 1981, 1983; Miller, 1983; Hartmann et al., 1983; Hoffmann, 1994; Trompette, 1994; Prave, 1996; Passchier et al., 2016) between the Angola-Congo and Kalahari cratons in southern Africa and the Rio de la Plata Craton in South America (e.g., Hartnady et al., 1985; Porada, 1989; Stanistreet et al., 1991; Frimmel and Frank, 1998). However, these belts display different deformational styles, crustal architecture, and tectonic histories. The Damara Belt spatially links with the Kaoko Belt along a structurally complex zone that is well-exposed southwest of Khorixas, in Namibia (Fig. 1B; Miller, 1983, 2008; Passchier et al., 2002). According to Frimmel (2009), the evidence of a connection zone between the Damara and the Gariep belts are scarce outcrops of Neoproterozoic metasedimentary rocks in the coastal sand dune field of Namibia (Fig. 1B). The sedimentary successions of the Damara, Kaoko, and Gariep belts have been interpreted as deposited in a continental rift to passive continental margin environment fringing the Angola-Congo and Kalahari cratons (e.g., Miller, 1983; Porada, 1989; Frimmel et al., 2002; Paciullo et al., 2007; Frimmel and Miller, 2009; Nascimento et al., 2016). Alkaline volcanic rocks in the Damara, Kaoko and Gariep belts are interpreted as syn-rift magmatism (Fig. 2; e.g., Miller, 1980; Hoffman et al., 1996; Nascimento et al., 2016). According to Frimmel et al. (2001), intracontinental rifting began between ca. 771 Ma and 750 Ma in the Gariep Belt. In the Kaoko Belt metavolcanics of ca. 740 Ma to 710 Ma are interpreted as products of syn-rift magmatism (Konopásek et al., 2014). Neoproterozoic rift basins are related to Rodinia break up on global scale. The successions deposited in these basins were deformed and metamorphosed during the Pan-African tectonic events. The sequence for the closure of the Damara (generating the Damara Belt) and Adamastor oceans (with the formation of Kaoko and Gariep belts) is considered as follows. Stanistreet et al. (1991), Lehmann et al. (2016) and Passchier et al. (2016) proposed that the Damara Ocean started to close before the closure of the southern part of the Adamastor Ocean. Prave (1996), Gray et al. (2008) and Schmitt et al. (2012) proposed that the ending of the northern Adamastor Ocean occurred from 570 to 545 Ma (Kaoko and Gariep belts), before the final closure of the Damara Ocean, at ca. 520 Ma (Damara Belt). Miller (1983, 2008) divided the Kaoko Belt from east to west into four tectonostratigraphic zones (Fig. 1B): (i) Eastern Kaoko (EK), (ii) Central Kaoko (CK), (iii) Western Kaoko (WK) subdivided into Orogen Core (OC) and Coastal Terrane (CT), and (iv) Southern Kaoko Zone (SKZ). The Eastern, Central and Southern Kaoko zones and the Orogen Core are interpreted as the reworked passive margin of the Congo-Angola Craton (e.g., Miller, 1983, 2008). The Coastal Terrane is interpreted as an exotic terrane, accreted to this margin at ca Ma (e.g., Goscombe et al., 2005; Goscombe and Gray, 2007; Heilbron et al., 2008) Damara Belt The Damara Belt was formed during the Neoproterozoic to early Paleozoic Pan-African Orogeny (e.g., Hoffmann, 1990; Kukla and Stanistreet, 1991; Kukla, 1992; Blanco et al., 2011). Miller (1983, 2008) subdivided the belt into seven NE-SW trending tectonic zones, from north to south (Fig. 1B): (i) Northern Platform of Fig. 2. Geological map of the Vrede, Bethanis, Austerlitz and Toekoms areas, Damaraland, northwest Namibia, with location of the samples collected for geochronology. Modified from Nascimento et al. (2016).

4 156 D.B. Nascimento et al. / Gondwana Research 52 (2017) the Congo Craton (NP); (ii) Northern Margin Zone (NMZ); (iii) Northern Zone (NZ); (iv) Central Zone (CZ), subdivided into Northern and Southern parts (ncz and scz, respectively); (v) Southern Zone (SZ), including the Okahandja Lineament Zone (OLZ) and the Deep-level Southern Zone (DLSZ); (vi) Southern Margin Zone (SMZ); and (vii) Southern Foreland of the Kalahari Craton (SF). The Southern Zone is also referred to as the Khomas Trough (e.g., Martin, 1965; Martin and Porada, 1977). The protoliths of the Damaran metamorphic rocks are considered either related to the rifted margin of a single continent (Angola-Congo- Kalahari; e.g., Porada, 1979; Miller, 1983, 2008), or to two distinct margins associated to the Angola-Congo and Kalahari proto-cratons (e.g., Gray et al., 2008; Foster et al., 2015). (1) In the first hypothesis, two parallel northeast-trending paleo-rift basins were recognized in the Damara Belt (e.g., Porada, 1979; Miller, 1983, 2008; Eyles and Januszczak, 2007; Miller et al., 2009a, 2009b; Frimmel and Miller, 2009): the Northern Rift (Outjo Sea/Basin) and Southern Rift (Khomas Sea/Basin). The sedimentary fill of the former is recorded in the Northern and Northern Central Zone, while the sedimentary record of the second rift crops out in the Southern Zone (Fig. 1B). The Khorixas- Gaseneirob Thrust (KGT, Fig. 1B; Miller, 1983, 2008) separates the Outjo Basin to the south from the Northern Margin Zone and Northern Platform to the north. It corresponds to an ancient normal fault system, associated with Outjo Basin opening and reactivated later as a thrust fault during the Damara Orogeny (e.g., Miller, 1983, 2008; Nascimento et al., 2016). To the south, the Outjo Basin is limited by an intrabasinal basement structural high, the Abbabis High, mainly constituted of rocks of the Abbabis Complex (e.g., Smith, 1965; Sawyer, 1981; Brandt, 1987; Kröner et al., 1991; Longridge, 2012), with the Omaruru Lineament interpreted as its northern boundary (Fig. 1B). In this context, the Southern Central Zone (scz) would represent a former shelf environment on the Abbabis High. According to Kasch (1983), Miller (1983) and Kukla (1992), the scz became an active continental margin due to subduction to the NW and subsequent advance of the Kalahari Craton during the closure of the Khomas Sea. The subduction implied the formation of an accretionary prism in the SZ, which includes the Matchless amphibolite (Fig. 1B) interpreted as part of an ocean floor unit (e.g., Kukla, 1992; Meneghini et al., 2014). (2) In the second hypothesis, the Damara Belt successions record deposits in two distinct passive margins, related to the disconnected Angola-Congo and Kalahari paleocontinents (e.g., Gray et al., 2008; Foster et al., 2015). Intracratonic rifting volcanism occurred in both disconnected paleocratonic fragments in the same period of Ma (Gray et al., 2008). The structural and metamorphic history of the Damara Belt is well documented in the literature (e.g., Miller, 1979; Downing and Coward, 1981; Kasch, 1983; Kukla and Stanistreet, 1990, 1991; Kukla, 1992; Masberg et al., 1992; Passchier et al., 2002, 2016). The most important tectonic phase relates to intensive ductile deformation that coincides with high-t metamorphic conditions in the Central Zone and low-t in the Southern Zone, both recording medium P conditions (Miller, 1979; Kasch, 1983; Kukla and Stanistreet, 1991). The last deformational stages of the Damara Orogen are mainly related to coaxial NW-SE shortening (e.g., Meneghini et al., 2014; Passchier et al., 2016) and syn to latetectonic magmatism (e.g., Donkerhuk Granite; Miller, 1979; Clemens et al., 2017). The Damara Supergroup or Damara Sequence (SACS, 1980; Kröner, 1981; i.e., the Pan-African deposits of Miller, 1983) consists of a thick Neoproterozoic siliciclastic carbonate succession with interlayered volcanic rocks. The basal unit is the Nosib Group, interpreted as a record of the initial rifting phase (e.g., Porada, 1983; Borg, 2000; Miller, 2008; Nascimento et al., 2016). In the NMZ and NZ of Miller (1983, 2008); Fig. 1B), this group includes felsic and mafic volcanic rocks of the Naauwpoort Formation which represent bimodal magmatism attributed to the syn-rift (e.g., Miller, 1980; Hoffman et al., 1996; Jung et al., 2007; Nascimento et al., 2016) to transitional evolution stage of the basin (e.g., Nascimento et al., 2016). These igneous rocks yielded U-Pb (zircon) ages between 759 and 746 Ma (Hoffman et al., 1996; Nascimento et al., 2016) defining the approximate age for the basal succession in the Outjo Basin. This stage was followed by open marine basin conditions (between 740 and 590 Ma; Milani et al., 2015) that resulted in the deposition of the Otavi Group sediments, as well as the correlated Swakop, Zerrissene, Hakos, and Witvlei groups. 3. Local geology and previous geochronological data The study region is located in the northern part of the Northern Rift (Outjo Sea/Basin) in the NMZ and NZ (Miller, 1983, 2008; Fig. 1B) and comprises the Vrede (719), Bethanis (514), Austerlitz (515) and Toekoms (508) farms (Fig. 2; Nascimento et al., 2016). In this area, 15 mappable units (at scale 1:25,000) were recognized, correlating with formations of the Nosib, Otavi/Swakop and Mulden groups (Nascimento et al., 2016). The Neoproterozoic successions of the Nosib and Otavi/Swakop groups cover the Paleo/Mesoproterozoic basement unconformably. The Neoproterozoic Mulden Group foreland deposits cover the basement and the Nosib and Otavi/Swakop groups with an angular unconformity (e.g., Hoffman and Halverson, 2008). Nosib Group sedimentary successions are interpreted as alluvial fan deposits; while those of the Otavi and Swakop groups are understood as slope to basin deposits. Volcanic and subvolcanic rocks (Naauwpoort Formation) constitute bimodal rift-related magmatism. The successions were deformed and metamorphosed under greenschist facies (biotite zone) conditions during the Pan-African Orogeny; however, the prefix meta- which should precede the name of the rocks is omitted for the sake of brevity. In the southeast of the study area, two syn- to post-collisional Salem type granitic bodies of the Omangambo Pluton cut the Otavi Group upper unit (i.e., Kuiseb Formation; Fig. 2). Below we describe briefly the Damara Supergroup units according to Nascimento et al. (2016) and references therein, with the available geochronological data Nosib Group The Nosib Group includes granitic breccia and arkose of the Nabis Formation and felsic and mafic volcanic and subvolcanic rocks of the Naauwpoort Formation. The sedimentary succession is interpreted as alluvial fan deposits derived from the local basement (i.e., the Kamanjab Inlier) during the early stages of rifting that originated the Outjo Basin. The Naauwpoort igneous rocks record syn-rift magmatism interlayered in the Nabis Formation and in the Chuos Formation lowermost layers (Nascimento et al., 2016; and references in therein). U-Pb LA-ICPMS data from zircon grains of a pegmatite indicate a minimum depositional age of 763 ± 5 Ma for the base of the Nabis Formation in the Toekoms area (Table 1; McGee et al., 2012). To the east, in the Welwitschia Inlier, Hoffman et al. (1996) obtained a U-Pb (TIMS) zircon age of 756 ± 2 Ma for a syenite that intrudes the base of the Nosib Group, confirming the minimum age for the base of this unit (Table 1). U-Pb SHRIMP data from zircon crystals of a dacite layer (Naauwpoort Formation) interlayered close to the base of the Chuos Formation yielded an age of 757 ± 5 Ma. This age constrains the minimum age for the base of the Chuos Formation and the Nabis Formation in the Austerlitz area (Table 1; Fig. 3; Nascimento et al., 2016). These data are consistent with U-Pb (TIMS) zircon ages of 747 ± 2 Ma and 746 ± 2 Ma (Table 1) obtained for a rhyolite and a tuff layer, respectively, from the Naauwpoort Formation in the Summas Mountains (Hoffman et al., 1996), to the east of the study area. Maximum ages of

5 D.B. Nascimento et al. / Gondwana Research 52 (2017) Table 1 Geochronological data obtained from the Nosib and Otavi/Swakop groups in Damara Belt and interpretation of the meaning of the ages. Formation Rock type Area Methods Age (Ma) Reference Maximum Depositional Minimum Nosib Group Nabis Pegmatite Toekoms U-Pb LA-ICPMS 763 ± 5 McGee et al. (2012) Syenite Welwitschia U-Pb TIMS 756 ± 2 Hoffman et al. (1996) Arkose Vrede U-Pb LA-ICPMS 1484 ± 38 Nascimento et al. (this paper) Arkose Toekoms U-Pb LA-ICPMS 1033 ± 9 Nascimento et al. (this paper) Naauwpoort Dacite Austerlitz U-Pb SHRIMP 757 ± 5 Nascimento et al. (2016) Rhyolite Summas Mountains U-Pb TIMS 747 ± 2 Hoffman et al. (1996) Tuff Summas Mountains U-Pb TIMS 746 ± 2 Hoffman et al. (1996) Etusis Metasedimentary rock Central Zone U-Pb LA-ICPMS 872 ± 36 Foster et al. (2015) Khan Metasedimentary rock Central Zone U-Pb LA-ICPMS 915 ± 28 Foster et al. (2015) Otavi/Swakop groups Naauwpoort Dacite Austerlitz U-Pb SHRIMP 757 ± 5 Nascimento et al. (2016) Devede (Chuos) Tuff Summas Mountains U-Pb TIMS 760 ± 1 Halverson et al. (2005) Chuos Subarkose Vrede U-Pb LA-ICPMS 743 ± 10 Nascimento et al. (this paper) Subarkose Toekoms U-Pb LA-ICPMS 1016 ± 9 Nascimento et al. (this paper) Volcaniclastic wacke Toekoms U-Pb LA-ICPMS 949 McGee et al. (2012) Ghaub Ash Southern Central Zone U-Pb TIMS 635 ± 1 Hoffmann et al. (2004) Breccia Bethanis U-Pb LA-ICPMS 663 ± 5 Nascimento et al. (this paper) Brak River Feldspathic quartzite Bethanis U-Pb LA-ICPMS 1547 ± 48 Nascimento et al. (this paper) Arkose Bethanis U-Pb LA-ICPMS 677 ± 5 Nascimento et al. (this paper) Karibib Metasedimentary rock Central Zone U-Pb LA-ICPMS 629 ± 21 Foster et al. (2015) Kuiseb Metasedimentary rock Northern Zone U-Pb LA-ICPMS 606 ± 24 Foster et al. (2015) Metasedimentary rock Northern Zone U-Pb LA-ICPMS 619 ± 16 Foster et al. (2015) Metasedimentary rock Okahandja Zone U-Pb LA-ICPMS 587 ± 9 Foster et al. (2015) Metasedimentary rock Southern Zone U-Pb LA-ICPMS 601 ± 20 Foster et al. (2015) Metasedimentary rock Southern Zone U-Pb LA-ICPMS 607 ± 20 Foster et al. (2015) Omangambo Pluton Granite Northern Zone U-Pb LA-ICPMS 527 ± 7 Milani et al. (2015) Granite Northern Zone U-Pb 495 ± 4 Miller (2008) 872 ± 36 Ma and 915 ± 28 Ma (Table 1) were obtained from detrital grains (U-Pb LA-ICPMS) of the Nosib Group (Etusis and Khan formations) in the Central Zone area (Fig. 1B) by Foster et al. (2015) Otavi/Swakop groups The correlated Otavi and Swakop groups (Nascimento et al., 2016; and references therein) include siliciclastic, carbonate and mixed successions that correspond to eight Damara Supergroup formations in the study area (Figs. 2 and 3; Nascimento et al., 2016). The Chuos Formation contains siliciclastic rudite, sandstone, lutite and minor carbonate rocks; the Gauss Formation is composed of siliciclastic sandstone and lutite. Carbonate rocks predominate in the Berg Aukas and Auros formations. These units are unconformably covered by the Brak River, Ghaub, Karibib and Kuiseb formations (Figs. 2 and 3). The Brak River Formation is subdivided in the lower feldspathic quartzite and the upper arkose-lutite turbidite. The Ghaub Formation is constituted mainly of rudstone grading laterally to the upper turbidite of the Brak River Formation (Fig. 3). These units are covered by the carbonate unit of the Karibib Formation, and on top of this, by fine arkose-lutite turbidite, lutite and carbonate rocks of the Kuiseb Formation (see Nascimento et al., 2016 for more detail). These formations are interpreted to represent slope to basin gravitational deposits and were grouped in two tectonic sequences developed during the opening of the Outjo Basin (Fig. 3; Nascimento et al., 2016). The lower sequence which includes the four basal units (i.e., Chuos, Berg Aukas, Gauss and Auros formations) is interpreted as containing slope and proximal submarine fan successions. However, autochthonous Chuos Formation shallow water deposits (i.e., carbonate deposits containing stromatolites) occur in the Bethanis area, leading to the interpretation that this area was a structural high during deposition of, at least, part of this formation (Nascimento et al., 2016). The Brak River, Ghaub, Karibib and Kuiseb formations constitute the upper sequence interpreted as slope (Ghaub) and medium to distal fan deposits (e.g., Swart, 1992; Paciullo et al., 2007; Nascimento et al., 2016). Carbonate lenses interpreted as slide blocks derived from the failure of carbonate slopes (i.e., olistoliths) occur in both tectonic sequences, encased in the Berg Aukas Formation and the upper unit of the Brak River Formation (Nascimento et al., 2016). In the southeast of the study area, within-plate granite bodies including the Omangambo Pluton (e.g., Miller and Frimmel, 2009) intrude the Kuiseb Formation. As shown above, U-Pb zircon data of 757 ± 5 Ma constrain the minimum age of the base of the Chuos Formation in the Austerlitz area (Table 1; Nascimento et al., 2016). A depositional age of 760 ± 1 Ma (Table 1) was obtained from zircon grains (U-Pb TIMS) in a tuff of the Devede Formation, Ombombo Subgroup (Chuos Formation in this paper) in the Summas Mountains (Halverson et al., 2005). A maximum age of 949 Ma for the Chuos Formation in the Toekoms area was obtained from detrital zircon grains (U-Pb LA-ICPMS) by McGee et al. (2012). A depositional age of 635 ± 1 Ma (U-Pb TIMS) for the Ghaub Formation is based on zircon grains of an ash layer in the Karibib-Usakos area, in the Southern Central Zone (Fig. 1B; Hoffmann et al., 2004). A younger detrital zircon age (U-Pb LA-ICPMS) of 629 ± 21 Ma for the Karibib Formation constrains the maximum age for deposition of this unit in the Central Zone area (Fig. 1B; Foster et al., 2015). Younger U-Pb (LA- ICPMS) detrital zircon ages of 606 ± 24 Ma and 619 ± 16 Ma were obtained from two samples of the Kuiseb Formation in the Northern Zone, to the east of the study area. This same unit was sampled from other regions, such as the Okahandja Zone, and Southern Zone (Fig. 1B), and the youngest zircon ages obtained are 587 ± 9 Ma, 601 ± 20 Ma and 607 ± 20 Ma (Foster et al., 2015). These ages are close to 40 Ar- 39 Ar metamorphic ages of 590 Ma obtained in the Kuiseb Formation in the Austerlitz area (Lehmann et al., 2016). This metamorphic age constrains the end of deposition, and probably the beginning of convergent tectonics. These units are intruded by the syn-collisional granitic magmatism that produced the Omangambo Pluton. Rocks of this pluton yielded U- Pb zircon ages of 527 ± 7 Ma (WITS-RTX project, in Milani et al., 2015) and495±4ma(fig. 3; Miller, 2008) Mulden Group Rudstone, grainstone, dolomite, arkose and lutite are the main lithologies of the Mulden Group in the study area (Figs. 2 and 3). The minor

6 158 D.B. Nascimento et al. / Gondwana Research 52 (2017) Fig. 3. Stratigraphic column for the study area showing the main geochronological data: in red (present work) and blue (literature) utilized to define depositional age intervals. Thickness of the units not to scale. Modified from Nascimento et al. (2016). (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.) carbonate rock layers probably correlate with the Braklaagte Formation; the siliciclastic rocks may correlate to the Renosterberg Formation (Hoffman and Halverson, 2008), both formations of the Welkom Subgroup. The Mulden Group is interpreted as a molasse succession at the margins of the Damara and Kaoko belts (Hoffman and Halverson, 2008). The sediments were deposited after the second deformational phase of the Damara Orogen (Passchier et al., 2016; Lehmann et al., 2016) that corresponds to collision in the Kaoko Belt (~570 Ma), which constitutes its main sedimentary source (e.g., Germs et al., 2009; Miller, 2013; Lehmann et al., 2016). According to 40 Ar- 39 Ar and K-Ar thermochronological data, the age of deposition of the Mulden Group sediments is between 580 and 550 Ma (Clauer and Kröner, 1979; Gray et al., 2006; Lehmann et al., 2016). 4. Analytical procedures 4.1. Analyzed samples For provenance studies, eight samples, from the Nabis (2 samples), Chuos (2), lower (1) and upper (1) Brak River, Ghaub (1) and Renosterberg (1) formations (Supplementary Table 1; Figs. 2 and 3) were analyzed. The Nabis Formation (Nosib Group) samples are from northeast of the Vrede area and from the Toekoms area (samples N and N , respectively; Fig. 2). Both samples are from the arkosic matrix of a granitic breccia (Fig. 4A and B). Two subarkose samples of the Chuos Formation are from the Vrede and the Toekoms areas (samples N and N ; Figs. 2 and 4C). Locally, in the Bethanis

7 D.B. Nascimento et al. / Gondwana Research 52 (2017) area, the Ghaub Formation has a siliciclastic matrix that was also sampled (sample N14-2-7; Figs. 2 and 4D). A feldspathic quartzite and an arkose from the lower and upper units of the Brak River Formation are localized in the Austerlitz area (N and N11-5-2; Figs. 2 and 4E). The sample from the Renosterberg Formation (N ) is a subarkose from the Toekoms area (Figs. 2 and 4F) ; Figs. 5 and 6). RL was done with a binocular microscope and CL obtained using a Quanta 250 FEG electron microscope equipped with Mono CL3 + cathodoluminescence spectroscope (Centaurus). These steps were carried out in the Geochronological Research Center, University of São Paulo (CPGeo-USP) U Pb analysis 4.2. Mineral separation The selected samples (3 5 kg) were crushed in a jaw crusher, and disk grinder and then sieved to b500 μm grain size (Supplementary Table 1). Heavy minerals were separated by manual panning. These steps were performed in the laboratory of the Geological Survey of Namibia, Windhoek. The heavy minerals were separated using conventional heavy liquid (bromoform) and Frantz Isodynamic Magnetic Separator at the Laboratory of Geological Samples of the Federal University of Rio de Janeiro (UFRJ). Using a binocular microscope zircon grains were hand-picked from the heavy mineral fractions and, avoiding bias introduced by selection, no visual morphological or color differentiation was done. The zircon grains (150 to 300 grains from each sample) were mounted into a circular epoxy resin that was sanded and polished exposing a cross section of each zircon grain at the surface of the disc. Zircon grains were then photographed in reflected light (RL) and imaged using cathodoluminescence (CL) to identify fractures that may interfere in the analysis, as well as to show the internal structure (e.g., Corfu et al., The U-Pb analyses were performed using a pulsed excimer laser ablation (LA) microprobe (λ = 193 nm) coupled to an ICP-MS (Neptune) also at the CPGeo-USP. The ablation was done with a spot size of 32 μm, at a frequency of 6 Hz, intensity of 6 mj and ablation time of 40 s. Results from our samples were intercalated with measurements from the international standard (NIST and GJ) as well as with background (blank) measurements following the order: 5 standards (2 NIST and 3 GJ), two blanks, 13 unknown samples, four standards (2 NIST and 2GJ) and two blanks. This raw data was reduced using a lab internal spreadsheet and corrections for background, instrumental mass bias drift and common Pb were applied. The GJ standard has shown good repeatability for the Pb/U and Pb/Pb ratios but not for the Pb/Th ones, therefore, Pb/Th ages are not considered in this paper. The LAICP-MS data were reduced using the SQUID 1.02 program (Ludwig, 2001). According to the standards of the laboratory, the 206Pb/238U ages were used for zircon grains younger than 1300 Ma while 207 Pb/206Pb ages were used for older samples. Only U Pb zircon ages Fig. 4. Sampled outcrops. A) Nabis Formation (Nosib Group) granitic breccia north of Vrede, and B) Toekoms area. C) Chuos Formation subarkose from the Toekoms area. D) Feldspathic quartzite from the lower Brak River Formation, Austerlitz area. E) Carbonatic fragments and subarkose matrix, Ghaub Formation, Bethanis area. F) Subarkose from the Renosterberg Formation (Mulden Group), Toekoms area. The location of these outcrops is marked on Fig. 2.

8 160 D.B. Nascimento et al. / Gondwana Research 52 (2017) Fig. 5. Cathodoluminescence images showing U-Pb spots in representative zircon grains of the Nabis Formation, samples A) N14-1-3, and B) N ; and of the Chuos Formation, samples C) N14-1-6, and D) N Red circle spot analysis (32 μm scale); (xx.x) xx ± x number of spot and age in Ma; [0.13] Th/U. Sample location on Fig. 2. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.) with discordance equal to or b10% were considered in this work, plotted in the concordia diagrams and probability density plots (Figs. 7, 8 and 9) using the ISOPLOT 3.0 software (Ludwig, 2003). Zircon U Pb isotopic data for each sample are presented in Supplementary Table 2. Th/U ratios can be used to help distinguish primary igneous zircon (ratios N0.1) from metamorphic domains (ratio between 0.01 and 0.1; Williams and Claesson, 1987). This ratio was also used in other studies (e.g., Vavra et al., 1996; Hoskin and Black, 2000; Hartmann and Santos, 2004; Cavosie et al., 2004). In order to determine provenance of the zircon grains, the CL images of the crystals that yielded low Th/ U ratio(b0.1) were used to judge the probability of a metamorphic origin. 5. Results A total of 1384 LA-ICP-MS analyses were performed on 1225 detrital zircon grains extracted from eight rock samples (Supplementary Table 1, Figs. 2, 3 and 4). 551 zircon grains yielded 569 concordant 206 Pb/ 238 U and/or 207 Pb/ 206 Pb ages (Supplementary Table 1). The remaining analyses presented high discordance and/or high common Pb. Even with a wide range of analyses, the number of concordant analyses obtained from some samples (e.g. sample N14-2-9, 55 concordant analyses out of 263) is small (Supplementary Table 1), reflecting disturbance of the isotopic system. The largest data set comprises 97 concordant analyses; the smallest, 32. The RL photographs and CL images (Figs. 5 and 6)revealthat,forall eight samples, most of the zircon grains are rounded or sub-rounded, but sub-angular mostly broken grains also occur. They vary in size from 60 to 350 μm in length (Supplementary Table 1), and in color from translucent to brown, most of them being pink. Oscillatory zoning characteristic of igneous growth (e.g., Corfu et al., 2003) is present in most of the populations, but some grains show sector zoning or are homogeneous (Figs. 5 and 6). Predominantly the analyses show 232 Th/ 238 U ratios higher then 0.1, indicating igneous growth, but seven concordant zircon analyses yielded ratios between 0.01 and 0.1 (Supplementary Table 2) and were interpreted as metamorphic Nabis Formation Nosib Group Two samples of arkose from this lithostratigraphic unit were analyzed.

9 D.B. Nascimento et al. / Gondwana Research 52 (2017) Fig. 6. Cathodoluminescence images showing U-Pb spots in representative zircon grains from samples of: A) the lower Brak River Formation (N14-2-9), B) upper Brak River Formation (N11-5-2), C) Ghaub Formation (N14-2-7), and D) Renosterberg Formation (N ). Red circle spot analysis (32 μm scale); (xx.x) xx ± x number of spot and age in Ma; [0.13] Th/U. Sample location on Fig. 2. (For interpretation of the references to r in this figure legend, the reader is referred to the web version of this article.) In the sample collected from the northern Vrede area (sample N14-1-3; Figs. 2 and 4A) 68 analyses were performed on 61 zircon grains (Supplementary Table 1). Most of the crystals are metamictic, and only 32 concordant analyses were obtained from 29 grains (Supplementary Table 1). Only one concordant age was attained in a preserved core of a metamictic grain (zircon grain #1; Fig. 5A). This age, 1720 Ma, is coherent with the major population of the concordant grains analyzed. The probability plot for the sample presents a major peak at 1770 Ma (Fig. 7A). The three oldest ages are 1882 ± 60 Ma, 1877 ± 44 Ma and 1868 ± 53 Ma (#18.1, 50 and 18.2; Fig. 5A). Most of the crystals (27, ~84% of total) yielded ages between 1790 ± 34 Ma and 1710 ± 82 Ma (Figs. 5A and7a). One enigmatic grain yielded a rim of 1770 ± 36 Ma and core of 1593 ± 45 Ma (spots #45.1 and 45.2, with Th/U ratio of 0.73 and 0.49, respectively; Figs. 5A and 7A). The youngest analysis from this sample yielded an age of 1484 ± 38 Ma (#41; Figs. 5Aand 7A); this is the only analysis with Th/U b 0.1 (0.09). In sample N , from the Toekoms area (Figs. 2 and 4B), 152 analyses were performed on 131 zircon grains, with 92 concordant analyses carried out on 89 grains (Supplementary Table 1). The probability plot shows two major peaks at 1992 Ma and 1744 Ma (Fig. 7B). The oldest zircon crystal yielded a 207 Pb/ 206 Pb concordant age of 2575 ± 34 Ma (core analysis, #94; Fig. 5B). One grain produced an age of 2476 ± 35 Ma (#4; Fig. 5B) and two of 2071 ± 39 and 2066 ± 34. The 84 results between 2021 ± 37 Ma and 1489 ± 41 Ma represent the most abundant population (~91%; Fig. 7B). The only crystal with Th/U ratio b 0.1 yielded an age of 1572 ± 49 Ma (zircon grain #39, ratio of 0.06; Fig. 5B). Two ages of 1252 ± 11 and 1215 ± 17 Ma and one of 1033 ± 9 Ma were also obtained. The youngest grain yielded a Neoproterozoic concordant age of 711 ± 8 Ma (spot #2; Fig. 5B), which was discarded (see Section 6.1 for explanation) Otavi/Swakop group Chuos Formation Two subarkose samples, both near the top of the unit, one in the Vrede and the other in the Toekoms area (Fig. 2), were analyzed. In total, 282 analyses were performed on 260 zircon grains. Only 170 analyses within 163 grains are concordant. The sample from the Vrede area (N14-1-6; Fig. 2) yielded 98 concordant analyses performed on 91 grains (Supplementary Table 1). The probability plot shows five peaks at 1986 Ma, 1868 Ma, 1761 Ma, 1287 Ma and 1234 Ma (Fig. 7C). The oldest analyses are 2724 ± 31 Ma, 2298 ± 29 Ma and 2100 ± 31 Ma (#33, 72 and 65, respectively; Figs. 5C and7c). The most abundant population of 61 analyses (~63%) ranges between 2060 ± 32 Ma and 1566 ± 32 Ma (Fig. 7C). Four zircon grains yielded ages between 1477 ± 35 Ma and 1413 ± 38 Ma and 28 are bracketed among 1358 ± 58 Ma and 1221 ± 8 Ma. Three grains with Th/U ratios of 0.03, 0.08 and 0.07, yielded ages of 1566 ± 32 Ma, 1287 ± 9 Ma and 1221 ± 8 Ma, respectively (#53, 86 and 18; Fig. 5C).

10 162 D.B. Nascimento et al. / Gondwana Research 52 (2017) Fig. 7. Probability density diagram and histogram (left) and Concordia diagram (right) for detrital zircon ages of the Nabis Formation, samples A) N14-1-3, and B) N ; and of the Chuos Formation, samples C) N14-1-6, and D) N Values in the left diagrams indicate the peak age modes.

11 D.B. Nascimento et al. / Gondwana Research 52 (2017) Fig. 8. Probability density diagram and histogram (left) and Concordia diagram (right) for detrital zircon ages of: A) the lower Brak River Formation (N14-2-9), B) the upper Brak River Formation (N11-5-2), and C) the Ghaub Formation (N14-2-7). Values in the left diagrams indicate the peak age modes. The youngest concordant zircon grains (#66 and 4; Fig. 5C) have a 206 Pb/ 238 U age of 786 ± 7 Ma and 743 ± 10 Ma (Fig. 7C). The sample from the Toekoms area (N ; Figs. 2 and 4C) yielded 72 concordant analyses out of 163 on 152 analyzed grains (Supplementary Table 1). The probability plot for this sample shows three peaks at 1961 Ma, 1838 Ma and 1754 Ma (Fig. 7D).The oldest zircon is dated at 2738 ± 29 Ma (#85; Figs. 5D and7d) very similar to the Vrede sample; one grain yielded an age of 2203 ± 42 Ma (#86) and five grains from 2074 ± 46 Ma to 2004 ± 47 Ma (#53, 77, 65, 27 and 31; Fig. 5D). Fifty-nine ages between 1974 ± 38 Ma and 1700 ± 43 Ma represent the most abundant population (~82%; Fig. 7D). Two grains present an age of 1.6 Ga (1604 ± 42 Ma and 1604 ± 41 Ma;

12 164 D.B. Nascimento et al. / Gondwana Research 52 (2017) Fig. 9. Probability density diagram and histogram (left) and Concordia diagram (right) for detrital zircon ages of the Renosterberg Formation, Mulden Group (N ). Values in the left diagram indicate the peak age modes. #52 and 82; Fig. 5D) and two of 1.5 Ga (1541 ± 49 Ma and 1523 ± 89 Ma; #76 and 39; Fig. 5D). The two youngest grains yielded ages of 1069 ± 7 Ma and 1016 ± 9 Ma (#36 and 26; Figs. 5D and7d) Brak River Formation The feldspathic quartzite sample (N14-2-9; Figs. 2 and 4D) from the lower unit, shows 55 concordant U-Pb ages out of a total of 277 analyses carried out on 263 different grains (Supplementary Table 1). The probability plot shows two peaks at 1971 Ma and 1792 Ma (Fig. 8A).The most abundant population age (78%, 43 zircon grains) is between 2008 ± 29 Ma and 1908 ± 45 Ma (Figs. 6A and 8A). Eleven analyses range between 1812 ± 33 Ma (#14) and 1751 ± 35 Ma (#85), and the youngest grain yielded an age of 1547 ± 48 Ma (#46; Fig. 6A). The arkose sample (N11-5-2; Fig. 2) from the upper unit yielded 70 concordant U-Pb ages out of 217 analyses from 188 grains (Supplementary Table 1). The probability plot shows a single peak at 1968 Ma (Fig. 8B). The oldest zircon presented an age of 2542 ± 94 Ma (#33; Fig. 6B). Five analyses range between 2471 ± 34 Ma and 2379 ± 41 Ma and one yielded 2302 ± 44 Ma. The most abundant population (45 zircon grains; Fig. 8B) falls between 2122 ± 111 Ma and 1905 ± 41 Ma. In addition, eight zircon crystals yielded ages between 1848 ± 135 Ma and 1727 ± 54 Ma, three between 1538 ± 70 Ma and 1507 ± 113 Ma, one at 1130 ± 104 Ma, three at 1.0 Ga (1070 ± 9 Ma, 1023 ± 9 Ma and 1016 ± 8 Ma) and one at 965 ± 8 Ma. The youngest zircon grains yielded ages of 709 ± 40 Ma and 677 ± 5 Ma. The 2542 ± 94 Ma and 709 ± 40 Ma grains had Th/U ratios of 0.1 and 0.01 respectively (#33 and 43; Fig. 6B) Ghaub Formation A subarkose matrix from a rudstone breccia was sampled in the Bethanis area (sample N14-2-7; Figs. 2 and 4E). In total, 210 analyses were performed on 171 grains, but only 61 analyses yielded concordant ages (Supplementary Table 1; Figs. 6C and 8C). The probability plot shows a major peak at 1970 Ma (Fig. 8C).The oldest zircon grain presents an age of 3294 ± 23 Ma (#4; Figs. 6Cand8C). One grain produced a 2631 ± 30 Ma age, and another one an age of 2546 ± 35 Ma (#13 and 87, respectively; Figs. 6C and 8C). Forty grains varied between 2037 ± 46 Ma and 1930 ± 51 Ma (65%) representing the most abundant population (Fig. 8C). Eleven analyses yielded ages in the range of 1858 ± 56 Ma and 1750 ± 32 Ma, one at 1489 ± 59 Ma, one at 1370 ± 55 Ma, one at 1000 ± 7 Ma and one at 731 ± 7 Ma. Younger ages were acquired in only three zircon grains at 674 ± 7 Ma, 663 ± 5 Ma and 549 ± 5 Ma. The youngest crystal was discarded (see Section 6.1) Mulden Group Sample N was taken from a subarkose of the Renosterberg Formation in the Toekoms area (Figs. 2 and 4F). In total 178 analyses were performed on 151 zircon grains, with 89 analyses on 84 grains being concordant (Supplementary Table 1). The probability plot shows six peaks at 2624 Ma, 1856, 1052 Ma, 1011 Ma, 709 Ma and 650 Ma (Fig. 9). The oldest population varies between 2664 ± 27 Ma (#75) and 2556 ± 34 Ma (#32.2; Fig. 6D; 11 grains). Two crystals have ages of ca. 2.0 Ga (2055 ± 33 Ma and 2045 ± 37 Ma), four grains are dated between 1977 ± 29 Ma and 1924 ± 44 Ma and five grains yielded ages between 1873 ± 33 Ma and 1831 ± 38 Ma (Fig. 9). One grain of 1390 ± 63 Ma, two of 1.1 Ga (1171 ± 9 Ma and 1107 ± 12 Ma) and 29 (32%) grains produced ages between 1076 ± 15 Ma and 935 ± 8 Ma. Two grains with Neoproterozoic ages of 842 ± 7 Ma and 773 ± 8 Ma were also identified. The most abundant population age (33 grains; 37%; Fig. 9) ranges between 717 ± 5 Ma and 591 ± 21 Ma (#94; Fig. 6D). 6. Interpretation and discussion The probability density plots for the two samples of the Nabis Formation (Fig. 7A and B) indicate a main source area of ~1760 Ma (Fig. 3). This formation records deposition in alluvial fan setting close to a local basement source, which prevents grains from a more distal source. The distribution of the basement rocks (e.g., orthogneiss, quartz schist and amphibolite; Nascimento et al., 2006) can provide different zircon populations to distinct parts of the basin. The local source can explain the absence of Archean and Neoproterozoic ages in the Vrede area sample. This difference may also be related and/or accentuated by the fact that most of the zircon grains are metamictic and did not generate concordant ages. The age of the main source rocks for the sample of the Chuos Formation collected in the Toekoms area is the same as the Nabis Formation (i.e., ~1760 Ma; Fig. 7D). Also, this sample shows a 1960 Ma minor contribution which is close to the lowest value obtained for the Nabis Formation sampled in the same area (~1990 Ma). The two samples also present Stenian age provenance. The sample of the Chuos Formation from the Vrede area also showed asignificant amount of zircon grains of ca Ma and 1980 Ma

13 D.B. Nascimento et al. / Gondwana Research 52 (2017) (Fig. 7C). However, these crystals do not belong to the main 1870 Ma population. Crystals with this age constitute a second major source for the other sample of this formation collected from the Toekoms area (Fig. 7D). The probability density plots for both samples of the Brak River Formation and the Ghaub Formation show a main source area of ca Ma (Fig. 8). A 1790 Ma contribution is also present in the sample of the lower Brak River Formation (Fig. 8A). Samples of the upper Brak River and Ghaub formations show a similar pattern with Neoproterozoic and Neoarchean ages (Fig. 8B and C). Apart from this, the sample of the Ghaub Formation yielded a single Paleoarchean grain (Fig. 8C). The sample of the Renosterberg Formation (Mulden Group) presents an entirely different pattern, with principal source areas at ca. 650 Ma, 709 Ma, 1011 Ma, 1052 Ma, 1856 and 2624 Ma (Fig. 9). The probability density plots for the samples of the Nosib and Otavi/ Swakop groups show aging of the source area towards the top of the succession. This is probably related to periodic uplift of the source area producing erosion of increasingly older rocks Th/U ratio The single zircon crystal obtained from the Nabis Formation in the Vrede area (N14-1-3) that yielded Th/U ratio of 0.09 and age of ~1490 Ma shows oscillatory zoning in cathodoluminescence image (CL) and is interpreted as magmatic. However, even with Th/U ratio of 0.49, the core age of ca Ma, younger than the rim of 1770 Ma, obtained for zircon grain #45 is interpreted as non-igneous (i.e., recrystallized or hydrothermally affected). The ~1770 Ma rim is interpreted as igneous, and the younger age probably reflects a disturbance in the isotopic system. The six crystals that yielded low Th/U ratio are interpreted as probably metamorphic. One crystal is from the Nabis Formation, Toekoms area (N ; Th/U of 0.06 and age ~1570 Ma), three are from the Chuos Formation, Vrede area (N14-1-6; 0.03, 0.08 and 0.07; ~1560 Ma, 1290 Ma and 1220, respectively), and two from the upper Brak River Formation (N11-5-2; 0.1 and 0.01; ~2540 Ma and 710 Ma). CL images show that crystals with ages similar to the seven grains mentioned above with low Th/U ratio can be both igneous and metamorphic. Thus, it is likely that there are both igneous and metamorphic sources, with ages of ca Ma, 1570 Ma, 1480 Ma and 710 Ma, and between ~1290 Ma and 1220 Ma, for the sedimentary protoliths of the Damara Supergroup Depositional age The youngest reliable maximum depositional ages obtained for the Nabis Formation are 1484 ± 38 Ma and 1033 ± 9 Ma from the Vrede and Toekoms areas respectively (Table 1). In the Austerlitz area, the base of the Nabis Formation should be older than 757 ± 5 Ma, a concordant age of Naauwpoort dacite intrusive in this unit (Fig. 3; Nascimento et al., 2016). In the Toekoms area, the base of the formation should be older than 763 ± 5 Ma (McGee et al., 2012). Based on the obtained data, the depositional age for the Nabis Formation lies between ca and 757 Ma. Considering the detrital zircon age obtained by McGee et al. (2012), this interval is constrained to Ma (Fig. 3). In this context, the concordant age of 711 ± 8 Ma obtained from a single detrital zircon crystal of the Toekoms area has to be discarded. The geological map and data for this region do not sustain a younger age for the Nabis Formation. Therefore, we consider this single grain as insufficient to change regional interpretations. The youngest detrital zircon age obtained for the Chuos Formation (Otavi/Swakop groups) is 743 ± 10 Ma, interpreted as the maximum depositional age for the strata near the top of the Chuos Formation in the Vrede area where the sample was taken (Table 1; Fig. 3). The youngest age obtained from the sample collected from the Toekoms area is 1016 ± 9 Ma (Table 1). An intrusive sill of the Naauwpoort Formation in the lower part of the Chuos Formation, with an age of 757 Ma (Nascimento et al., 2016) shows that the base of this formation must be older than 757 Ma. This fits well with the ~760 Ma age obtained for the correlate Devede Formation (Halverson et al., 2005). The age of 1547 ± 48 Ma is the youngest one obtained for detrital zircon from the lower unit of the Brak River Formation (feldspathic quartzite, N14-2-9). However, a single zircon grain from an arkose of the upper unit yielded a concordant age of 677 ± 5 Ma, interpreted as the maximum depositional age for the upper Brak River Formation (Fig. 3; Table 1). The Ghaub Formation was dated by Hoffmann et al. (2004) at 635 ± 1 Ma, based on an ash layer interlayered in this unit in the Southern Central Zone. The three youngest ages we obtained are 674 ± 7 Ma, 663 ± 5 Ma and 549 ± 5 Ma. The youngest of these is not consistent with the published age and our geological map and was therefore discarded. Hence, we interpret the maximum age for the analyzed sample of the Ghaub Formation as ca. 670 Ma, based on the two Cryogenian detrital crystals. The data above show an age interval from 743 to 635 Ma for deposition of the upper succession of the Chuos Formation, as well as the Berg Aukas, Gauss, and Auros formations. This range also includes the unconformity between the lower and upper tectonic sequences and the deposition of the lower Brak River Formation (Fig. 3). The deposition of the upper Brak River and Ghaub formations is constrained between ca. 663 and 590 Ma, an interval limited by the youngest zircon grain of the Ghaub Formation and the 40 Ar/ 39 Ar age obtained from phengites of the Kuiseb Formation formed during metamorphism and deformation (Fig. 3; Lehmann et al., 2016). This range also includes the deposition of the Karibib and Kuiseb formations with maximum depositional ages of ca. 630 Ma and 605 Ma, respectively (Foster et al., 2015). Finally, the youngest age of 591 ± 21 Ma is interpreted as the maximum depositional age for the Renosterberg Formation (Mulden Group) in the study area. According to Lehmann et al. (2016) the rocks of the Mulden Group were deposited between 580 Ma and 550 Ma Provenance and source area Samples of the Nosib and Otavi/Swakop groups show dominance of Paleoproterozoic zircon grains with minor Mesoproterozoic and some Archean and Neoproterozoic crystals. The sample of the Mulden Group is dominated by Neoproterozoic zircon grains with a Mesoproterozoic contribution and minor populations of Paleoproterozoic and Archean crystals (Figs. 7, 8 and 9;Supplementary Table 2). Compilation of geochronological data from the basement of the Angola-Congo, Kalahari and Coastal Terrane can be found in Supplementary Table 3. Despite the occurrence of crystals from the Siderian Period, these ages are not reported from the cratons surrounding the Damara Belt. The main Nabis Formation age spectra range between ca Ma and 1480 Ma. The possible source rocks are: A -Neoarchean - 1) Central Kaoko Zone gneisses (Fig. 1B) with ca Ma to 2580 Ma (Supplementary Table 3; Seth et al., 1998; Franz et al., 1999); 2) rocks of the distant Irumide Belt (Zambia) ranging from ca Ma to 2540 Ma (Key et al., 2001); 3) Orosirian igneous rocks of the Huab (Kamanjab Inlier) and Grootfontein (Otavi Mountains area) complexes with zircon xenocryst of ~2550 Ma to 2500 Ma (Lobo-Guerrero Sanz, 2005). B -Paleoproterozoic 4) gneiss and amphibolite of the Abbabis Complex (Central Zone; Fig. 1B) with ages between ~2060 and 1945 Ma (Kröner et al., 1991; Tack et al., 2002; Longridge, 2012; Longridge et al., 2014); 5) granite and ignimbrite of the Humpata Plateau (Angola block, southern Congo Craton) with ages of ca Ma, 1950 Ma and 1800 Ma (McCourt et al., 2013); 6) orthogneiss and amphibolite of the Epupa Complex (Kaoko

14 166 D.B. Nascimento et al. / Gondwana Research 52 (2017) Belt) with ages between ~2030 Ma and 1960 Ma, and ~1860 Ma and 1690 Ma, as well as of ~1930 Ma (Supplementary Table 3; Seth et al., 1998, 2005; Franz et al., 1999; Kröner et al., 2004, 2010; Goscombe et al., 2005; Luft et al., 2011); 7) metasedimentary rocks of the Epupa Complex with detrital zircon ages between ~1810 Ma and 1630 Ma (Seth et al., 2003); 8) granitic gneiss from NE Namibia with an age of ~2020 Ma (Hoal et al., 2000); 9) igneous rocks of the Grootfontein and Huab complexes with an age of ca Ma and ~1980 Ma to 1730 Ma, respectively (Lobo-Guerrero Sanz, 2005); 10) Rehoboth Inlier (SMZ; Fig. 1B) meta-igneous rocks with ages ranging from 1800 Ma to 1720 Ma (Burger and Coertze, 1978; Ziegler and Stoessel, 1993; Nagel et al., 1996; Becker et al., 1996; Hilken, 1998) and a Mesoproterozoic granite with xenocryst zircon of ca Ma (Lobo-Guerrero Sanz, 2005); 11) igneous rocks from the Irumide Belt with ages ranging from 2060 Ma to 1860 Ma (Key et al., 2001; Rainaud et al., 2002; De Waele, 2005). C -Mesoproterozoic 12) Epupa Complex orthogneiss with ages between ~1520 Ma and 1450 Ma and ca Ma (Supplementary Table 3; Seth et al., 1998; Littmann et al., 2000; Kröner et al., 2004; Luft et al., 2011), and metasedimentary rocks with metamorphic ages between ~1530 and 1510 (Seth et al., 2003); 13) igneous rocks of the Rehoboth Inlier, with ages ranging between ~1230 Ma and 1210 Ma (Burger and Coertze, 1978; Pfurr et al., 1991; Ziegler and Stoessel, 1993; Hoal and Heaman, 1995; Hilken, 1998; Schneider et al., 2004; Becker et al., 2005, 2006); 14) igneous rocks of the Irumide Belt ranging from 1080 to 1000 Ma (Rainaud et al., 2002; De Waele, 2005; De Waele et al., 2003); 15) gneiss from the Abbabis High with ca Ma (Kröner et al., 1991; Foster et al., 2015); 16) igneous rocks of the Namaqua Complex ranging from 1200 Ma to 1000 Ma (e.g., Burger and Coertze, 1975, 1978; Pfurr et al., 1991; Hilken, 1998; Robb et al., 1999; Nagel, 2000; Becker et al., 2005, 2006). Rocks with similar age also crop out in Botswana (Singletary et al., 2003). The Chuos Formation age spectra fall in the range between ca Ma and 1520 Ma. These data show that the main source rocks are the same as those of the Nabis Formation. The main difference, as shown below, is the presence of grains that yielded ages of ca Ma, 2300 Ma, 2200 Ma, 1360 to 1250 Ma and 750 Ma. The possible source rocks for these distinct grains are: 1) igneous rocks of the Irumide Belt with ca Ma (De Waele, 2005). Other possible Neoarchean sources are granitoids in the Congo, Tanzania, Zimbabwe and Kaapvaal cratons (Milani et al., 2015, and references therein); 2) granitoids from NW Namibia that yielded xenocrysts with Paleoproterozoic ages of ~2290 Ma and 2260 Ma (Supplementary Table 3; Sethetal., 1998; McCourtetal.,2013); Mesoproterozoic igneous rocks of the following complexes: 3)theKuneneComplex ranging from ~1390 Ma to 1370 Ma (Drüppel et al., 2000, 2007; Mayer et al., 2004; McCourt et al., 2013); 4) the Epupa Complex with ages from ~1380 Ma to 1310 Ma (Seth et al., 2005); 5) the Mudorib Complex (Western Kaoko Belt; Fig. 1B) with ages from ~1340 Ma to 1290 Ma (Luft et al., 2011); 6) the Rehoboth Inlier with ages ranging from ~1380 Ma to 1340 Ma (Ziegler and Stoessel, 1993; Hoal and Heaman, 1995; Nagel, 2000; Becker et al., 2006); 7) granitoid of the Kalahari Craton with an age of ~1100 Ma and xenocryst zircon of ~1200 Ma (Lobo-Guerrero Sanz, 2005). Neoproterozoic rift-related igneous rocks of the 8) Naauwpoort Formation with ages between ~760 Ma and 730 Ma (Supplementary Table 3; Miller and Burger, 1983; Hoffman et al., 1996; De Kock et al., 2000; Nascimento et al., 2016); 9) Huab Complex with ages between ~765 Ma and 745 Ma (Hoffman et al., 1996; Lobo-Guerrero Sanz, 2005); 10) Kaoko Belt with ages of ca. 740 Ma to 710 Ma (Konopásek et al., 2014). The main Brak River Formation age spectra range between ca Ma and 1720 Ma. This interval fits well with the Nabis and Chuos formations main age spectra confirming that the hinterland still continued to be eroded. The main difference is the presence of Neoproterozoic grains with ages of ca. 710 Ma and 680 Ma. The known sources for these crystals are the granitic gneisses of the Coastal Terrane (Kaoko Belt) that yielded ages of ca. 730 Ma and 690 Ma (Kröner et al., 2004). The main Ghaub Formation age spectra fall in the range between ca Ma and 1750 Ma. The difference is the presence of zircon grains with ages of ca Ma, 675 Ma and 660 Ma. Paleoarchean source rocks occur in the Congo Craton (Kasai Block; whole-rock Rb-Sr dates of between 3490 Ma and 3330 Ma; Cahen et al., 1984) and in the Kalahari Craton (U-Pb ages of ca Ma; Garzanti et al., 2014). The probable Cryogenian source is the magmatism that generated the 635 Ma tuff layer interlayered in the Ghaub Formation in the Southern Central Zone (Hoffmann et al., 2004). The sample of the Renosterberg Formation (Mulden Group) yielded U-Pb ages of zircon grains ranging from Archean to Neoproterozoic. The abundance of Neoproterozoic grains with also Ediacaran ages distinguishes the Mulden from the Nosib and Otavi/Swakop groups (Fig. 10). The probable source rocks for this Ediacaran population are localized in the Kaoko Belt. They include granite with age of ca Ma of the Central Zone (e.g., Johnson et al., 2006; Bergemann et al., 2014; Jung et al., 2015; Stammeier et al., 2015) and of ca. 580 Ma of the Coastal Terrane (Fig. 1B; Konopásek et al., 2008), and rocks with metamorphic ages between ~660 Ma and 570 Ma (Seth et al., 1998; Franz et al., 1999; Kröner et al., 2004; Goscombe et al., 2005; Konopásek et al., 2008) Regional interpretations A comparison of the detrital zircon ages with basement data from the literature (Supplementary Table 3) reveals that the sedimentary protolith of the Nosib and Otavi/Swakop rocks in the study area is most likely situated in the Angola-Congo Craton (Fig. 10), as previously suggested for the Toekoms area (McGee et al., 2012). Some authors also interpret the Kalahari Craton as a possible source area (e.g., McGee et al., 2012; Konopásek et al., 2014); others (e.g., Foster et al., 2015) consider that Congo and Kalahari cratons were not united during Rodinia times. The two samples of the Nabis Formation show main source areas with ages of ca Ma, 1870 Ma and 1750 Ma. This contrasts with those of ca Ma and 970 Ma (Fig. 11) obtained by Foster et al. (2015) for the correlative Etusis and Khan formations exposed in the Abbabis High, southern Central Zone (scz; Figs. 1B, 11 and 12). All these units are interpreted as derived from local sources (e.g., Foster et al., 2015; Nascimento et al., 2016) which is corroborated by the age of 1027 ± 10 Ma obtained from a basement gneiss in the structural high (Foster et al., 2015). According to Miller (2008) and Foster et al. (2015), among others, the Abbabis High was part of the Angola-Congo Craton. In this craton, rocks with ca Ma ages are exposed far away (N1000 km) from the Outjo Basin, in the Irumide Belt (Fig. 1A; e.g. De Waele et al., 2003; De Waele, 2005). However, rocks with ca Ma occur in the nearby Abbabis High and in the Rehoboth Inlier and Namaqua Complex, both within the Kalahari Craton (Fig. 1B; e.g., Burger and Coertze, 1975; Nagel, 2000; Singletary et al., 2003). We therefore interpret the Abbabis High as mainly related to the Kalahari proto-craton. According to this interpretation, and if the Congo and Kalahari cratons were distinct paleocontinents in Rodinia times (e.g., Gray et al., 2008; Foster et al., 2015), the Matchless amphibolite records a minor ocean separating two blocks of the Kalahari paleocontinent: the Abbabis High and the Kalahari Craton. Since there is no register of ocean floor rocks in the Outjo Basin, the two proto-cratons, Angola-Congo and Kalahari, were probably together during Rodinia times.

15 D.B. Nascimento et al. / Gondwana Research 52 (2017) Fig. 10. Probability density diagram comparing data of this work with those from the literature for the Angola-Congo (southern) and Kalahari cratons and the Coastal Terrane (Kaoko Belt) (references in Supplementary Table 3). Data obtained from the Irumide Belt are shown in a different color as those of the Angola-Congo Craton. Detrital zircon ages obtained by Foster et al. (2015) for the Kalahari and Angola-Congo Damara sequences are also shown. Vertical axis is not in scale. The Outjo Basin had two faulted edges, the northern one related to the Angola-Congo Craton and the southern one to the Abbabis High. The distal succession filled the center of the basin, while proximal facies were deposited along both margins. Since the studied succession is localized in the Angola-Congo Craton border and deepened towards the south, the main age peaks from rocks of the Kalahari Craton are not present (Fig. 10), discarding it as a main source area. Accordingly, the main source areas for the studied succession had to be to the N, NE, and NW, and therefore situated in the Angola-Congo Craton. According to Goscombe and Gray (2007, 2008), the Coastal Terrane of the Kaoko Belt (Fig. 1B) and the Angola-Congo Craton margin were not separated by an extensive ocean prior to the main Kaoko Belt collision (~570 Ma). This terrane is interpreted here as the main source area for the sediments that composed the Mulden Group (Fig. 10). Foster

16 168 D.B. Nascimento et al. / Gondwana Research 52 (2017) near the Kalahari, south of the Abbabis High (Fig. 1B) as related to Angola-Congo Craton sources. However, the ca Ma grains may have their source both in the Kalahari Craton and the Abbabis High, while the younger grains could be related to erosion of the Coastal Terrane. Volcanic intraclasts derived from the Naauwpoort Formation (Nascimento et al., 2016) reinforce the interpretation that the riftrelated rocks are also the source of ca. 750 Ma zircon grains found in the Chuos Formation and younger units. This is contrary to the interpretation that does not consider rift-related rocks as source of Neoproterozoic zircon grains (e.g. Konopásek et al., 2014) Tectono-sedimentary model for the Damara Supergroup at the Angola- Congo Margin Fig. 11. Probability density diagram for detrital zircon ages of the two samples of the Nabis Formation (blue, this work) and the samples of the correlate Etusis and Khan formations from the southern Central Zone (red, Foster et al., 2015). Values in the diagram indicate the peak age modes. Sample location on Figs. 1B and2. (Forinterpretationofthe references to color in this figure legend, the reader is referred to the web version of this article.) et al. (2015) separate the geochronological data according to sources in Angola-Congo or Kalahari cratons. The authors also considered Neoproterozoic data (ca Ma, and between 700 Ma and 600 Ma) The Damara Supergroup successions record rift to proto-ocean phases in the Angola-Congo-Kalahari basement as previously described by Miller (1983, 2008), Porada (1983, 1989), Stanistreet et al. (1991), Borg (2000) and Eyles and Januszczak (2004), among others. Rifting produced the NE-SW Outjo and Khomas basins separated by the Abbabis High, an intrabasinal basement structural high (Figs. 1B and 12A and C). The records of the initial continental rift phase in the northern Outjo Basin are the alluvial fan granitic breccia and arkose of the Nabis Formation deposited between 872 Ma and 757 Ma (Figs. 3 and 12C and D). The advance of lithospheric stretching led to the sea entrance in the rift and created accommodation space for the thick slope to basin deposits of the Chuos Formation. Contemporaneous 757 Ma bimodal magmatism (Naauwpoort Formation; Fig. 12D) constrains the minimum age for the base of the formation, while a 635 Ma tuff Fig. 12. Evolution of the Damara Supergroup stratigraphy in the northern Outjo Basin. Schematic maps show A) the continental rift successions of the Nabis Formation and correlatives; B) the rift to proto-ocean basins where the Otavi/Swakop successions were deposited. Arrows basin opening sense. Based on Fig. 9 of Porada (1989). Schematic sections showing basin fill related to the C) Nabis Formation, D) lower Brak River Formation and units of the lower tectonic sequence, and E) units of the upper tectonic sequence. Thickness of the units not to scale. KGT Khorixas-Gaseneirob Thrust; OML Omaruru Lineament; OL Okahandja Lineament.