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Journal of African Earth Sciences 44 (2006) 1–20

Geological Society of Africa Presidential Review, No. 10

Evidence for the Snowball Earth hypothesis in the Arabian-NubianShield and the East African Orogen

R.J. Stern a,*, D. Avigad b, N.R. Miller a, M. Beyth c

a Geosciences Department, University of Texas at Dallas, Box 830688, Richardson, TX 75083-0688, USAb Institute of Earth Sciences, Hebrew University of Jerusalem, Jerusalem 91904, Israel

c Geological Survey of Israel, 30 Malkhei Yisrael Street, Jerusalem 95501, Israel

Received 7 April 2005; received in revised form 26 September 2005; accepted 12 October 2005Available online 27 December 2005

Abstract

Formation of the Arabian-Nubian Shield (ANS) and the East African Orogen (EAO) occurred between 870 Ma and the end of thePrecambrian (�542 Ma). ANS crustal growth encompassed a time of dramatic climatic change, articulated as the Snowball Earthhypothesis (SEH). SEH identifies tremendous paleoclimatic oscillations during Neoproterozoic time. Earth’s climate shifted wildly, fromtimes when much of our planet’s surface was frozen to unusually warm episodes and back again. There is evidence for four principalicehouse episodes: �585–582 Ma (Gaskiers), �660–635 Ma (Marinoan), �680–715 Ma (Sturtian), and �735–770 Ma (Kaigas). Evidenceconsistent with the SEH has been found at many locations around the globe but is rarely reported from the ANS, in spite of the fact thatthis may be the largest tract of Neoproterozoic juvenile crust on the planet, and in spite of the fact that Huqf Group sediments in Oman,flanking the ANS, record evidence for Sturtian and Marinoan low-latitude glaciations. This review identifies the most important evidencepreserved in sedimentary rocks elsewhere for SEH: diamictites, dropstones, cap carbonates, and banded iron formation (BIF). Expectedmanifestations of SEH are integrated into our understanding of ANS and EAO tectonic evolution. If Kaigas and Sturtian events wereglobal, sedimentary evidence should be preserved in ANS sequences, because these occurred during an embryonic stage of ANS evolu-tion, when crustal components (island arcs, back-arc basins, and sedimentary basins) were mostly below sea level. Previous SEH inves-tigations have been largely reconnaissance in scope, but potentially diagnostic sedimentary units such as diamictites, marine carbonateswith d13C excursions and banded iron formations are reported from the ANS and are worthy of further investigation. Collision anduplift to form the EAO destroyed most marine sedimentary basins about 630 Ma ago, so evidence of Marinoan and Gaskiers glaciationswill be more difficult to identify. Several post-accretionary Neoproterozoic sedimentary basins in Arabia may preserve sedimentary evi-dence but such evidence has not been documented yet. The Huqf Group of Oman contains sedimentary evidence for the Marinoan gla-ciation but no evidence that the Gaskiers glaciation was significant in this part of the world. Deep erosion at �600 Ma throughout thenorthern ANS and EAO may be related to Marinoan continental glaciation, which may have accomplished much of the cutting of theANS peneplain, but final shaping of the peneplain took place over the next 60 million years.

African geoscientists can contribute to our understanding of Neoproterozoic climate change through careful field studies, and theinternational geoscientific community interested in Neoproterozoic climate change should pay attention to evidence from the ANS.Future investigations should include knowledge of the SEH and its controversial aspects, in addition to the greater plate tectonic settingof the ANS.� 2005 Elsevier Ltd. All rights reserved.

Keywords: Neoproterozoic; Snowball Earth; Arabian-Nubian Shield; East African Orogen

1464-343X/$ - see front matter � 2005 Elsevier Ltd. All rights reserved.

doi:10.1016/j.jafrearsci.2005.10.003

* Corresponding author.E-mail address: [email protected] (R.J. Stern).

mailto:[email protected]

2 R.J. Stern et al. / Journal of African Earth Sciences 44 (2006) 1–20

Contents

1. Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 22. The Snowball Earth hypothesis. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 33. The Arabian-Nubian Shield and the East African Orogen. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 54. Diagnostic evidence for Neoproterozoic glaciation and post-glacial warming . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7

4.1. Dropstones . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 74.2. Diamictites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 84.3. Cap carbonates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 104.4. Banded iron formations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 114.5. Paleomagnetic evidence . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 11

5. Expected manifestations of Neoproterozoic glaciations in the ANS and EAO . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 126. The record of glaciation in the ANS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 12

6.1. Evidence for Kaigas (�735–770 Ma) glaciation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 136.2. Evidence for Sturtian (�680–715 Ma) glaciation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 146.3. Evidence for Marinoan (�635–660 Ma) and Gaskiers (�582–585 Ma) glaciations . . . . . . . . . . . . . . . . . . . . . . . . . . 16

7. Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 17Acknowledgment . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 18References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 18

1. Introduction

There are three grand and intertwined Neoproterozoic(1000–554 Ma) themes. First is the evolution of increas-ingly complex life, whereby a biosphere characterized bysingle-celled organisms at the beginning of Neoproterozoictime evolved to be dominated by more complex multicellu-lar organisms at its end (Knoll, 2003). The second grandtheme is a great Wilson/Supercontinent Cycle, which beganwith the rupture of the Rodinian supercontinent and for-mation of new oceanic realms. As these fragments dis-persed, oceanic realms closed and a new supercontinentwas generated from the shards of Rodinia. The closingocean generated great fringing arcs and oceanic plateaus,and these were swept up in front of the advancing conti-nental fragments and incorporated into the new Gondwanasupercontinent. The third grand Neoproterozoic themeconcerns the tremendous paleoclimatic oscillations thathave become the focus of the ‘‘Snowball Earth hypothe-sis’’. Earth’s climate seems to have shifted wildly, fromtimes when perhaps the entire planet’s surface was frozen,quickly turning to sweltering greenhouses, and back again(Hoffman et al., 1998; Hoffman and Schrag, 2002).

These biological, tectonic, and climatic themes arerelated and present a spectacular example of global changeon the maturing Earth. It is a wonderfully interdisciplinaryeffort that seeks to understand how Neoproterozoic tecto-nics and life affected climate, and how Neoproterozoic tec-tonics and climate influenced biological evolution. A largepart of the research focuses on isotopic proxies of the C-cycle to discern how and why atmospheric concentrationsof greenhouse gasses CO2 and perhaps CH4 varied. Atmo-spheric concentrations of these gasses were important con-

trols of Phanerozoic climate (Royer et al., 2004), andshould also have been important for the Neoproterozoic.

Hypotheses of how Neoproterozoic life and climateinteracted are developing rapidly. One possibility is thatproliferating photosynthetic life increased atmosphericoxygen as it drew down atmospheric CO2, leading to cool-ing and allowed the development of a protective ozonelayer. In turn, global cooling and warming cycles may havestimulated evolution by alternately stressing the biosphereand providing warm, shallow water ecosystems when icemelted and sea level rose. Continental dispersal allowedmultiple ecological environments on different continentalshelves to develop in isolation, and so to stimulate evolu-tion (Valentine and Moores, 1974).

It is less clear how Neoproterozoic tectonics affected cli-mate because so many explanations are possible. Continen-tal configurations during Phanerozoic time exert importantcontrols on climate, with harsher climates during times ofsupercontinent assembly and warmer, more humid climatesduring times when continents were dispersed and sea levelis high (Worsley et al., 1986; Veevers, 1990). Similar con-trols must also have been important during Neoprotero-zoic time, although we are not yet confident that weunderstand continental configurations during this time.The supercontinent cycle exerted other controls on climateas well. Continental fragments produced from the breakupof Rodinia clustered at low latitudes, where in theoryintense chemical weathering associated with a very rainytropical climate absorbed atmospheric CO2, while organicmatter was buried near river deltas (Evans, 2000; Donna-dieu et al., 2004). Weathering of flood basalts that wereerupted in association with the break-up of Rodinia mayhave drawn down atmospheric CO2 to start the first

R.J. Stern et al. / Journal of African Earth Sciences 44 (2006) 1–20 3

glaciation (Godderis et al., 2003). An increase in explosivevolcanism is linked to Pleistocene glaciation in the northernhemisphere (Prueher and Rea, 2001), and an increase in arcvolcanism possibly triggered Neoproterozoic glaciations(Stern, 2005). Massive release of methane as a result ofdestabilizing methane hydrates in the continental shelveshas been suggested as an important contribution to rapidwarming following glaciation (Martin et al., 2001).

The spectacular nature of global climate variations inter-preted from Neoproterozoic lithostratigraphic sequencesand attendant isotope records has resulted in catastrophicexplanations. It is suggested that climatic oscillations mayhave been forced by radical changes in Earth’s pole of rota-tion (Williams, 1975; Hoffman, 1999; Evans, 2003) or as aresult of fundamental changes in Earth’s tectonic style(Stern, 2005). It may be some time before all of the possibleexplanations are advanced and tested, but it is clear thatunderstanding Neoproterozoic tectonics and paleogeogra-phy will be essential for understanding interactionsbetween Neoproterozoic life, climate, and tectonics. Theinterdisciplinary nature of the effort to understand interac-tions between the solid earth, hydrosphere, and biospherewill surely provide us with a more robust understandingof how the Earth system operates.

This essay explores the extent to which evidence of Neo-proterozoic Snowball Earth events, especially evidence formarine ice cover and continental glaciation, are preservedin the Arabian-Nubian Shield (ANS) of NE Africa andwestern Arabia. The ANS formed during Neoproterozoictime, and the early part of its evolution was associated withvolcanism and sedimentation below sea level, where sedi-mentary evidence of Snowball Earth episodes should bepreserved. A range of mostly marine environments charac-terized the embryonic ANS, from shallow-water shelves tothe abyssal seafloor, and some deposits should record theextreme climatic variations observed for this time periodelsewhere around the globe. There are extensive tracts ofophiolites in the ANS, crustal relicts of the Neoproterozoicdeep ocean. Sediments deposited on ANS ophiolites shouldrecord how the deep ocean behaved during and betweenSnowball Earth events. In spite of these opportunities,there has been litttle effort to use the ANS to investigateNeoproterozoic climate. We know of only two reports thatexplicitly identify rock sequences potentially pertaining toSnowball Earth episodes in the region (Beyth et al., 2003;Miller et al., 2003).

This review is intended to stimulate research of the SEHin the ANS in three ways. First, we hope to inform geosci-entists studying Neoproterozoic rocks in the ANS, so thatthey can help look for the evidence. Second, we hope todraw the attention of geology students to this exciting areaof cross-disciplinary and international research, especiallystudents in those nations that have ANS outcrops wherecareful field studies allow important contributions to bemade at relatively low cost. Finally, we hope to make theinternational community aware that the ANS is a promis-

ing area for understanding Neoproterozoic global change,and to encourage this community to extend their studiesto the ANS.

The organization of this paper is designed for each ofthese target audiences. In the following sections, we outlinethe Snowball Earth hypothesis and the evolution of theArabian-Nubian Shield. Then, we discuss what sorts of evi-dence should be sought in sedimentary rocks. Finally, weuse what we know about the timing of Neoproterozoic gla-ciations and tectonic evolution of the ANS to discuss whatis already recognized and what is likely or unlikely to bepreserved if the extreme climatic events inferred for the restof the world affected the evolving ANS.

A final caveat to the reader: many aspects of the Snow-ball Earth hypothesis are controversial and there is a deve-loping array of competing models to best account for thevital physical and chemical evidence. Most scientists agreethat Neoproterozoic time was characterized by remarkableclimate variations but details of this are still being resolved,for example whether or not the oceans were completely ice-covered and whether or not glaciations were globally syn-chronous (see for example Young, 2004; Williams, 2004).Those studying ANS exposures should keep an open mindabout what is observed and how this is best interpreted.

2. The Snowball Earth hypothesis

The Snowball Earth hypothesis (SEH) focuses on evi-dence that Earth experienced several cycles of unparalleledclimatic fluctuations during Neoproterozoic time and under-standing why this happened. Conditions alternated rapidlybetween ‘icehouse’ (intense, perhaps global ice cover) and‘greenhouse’ (globally warm) conditions (Evans, 2000).Hot and cold climatic swings may have been brief, perhapsa few millions of years long, and these separated by muchlonger intervals of more temperate climate. It is controver-sial whether or not the entire Earth ever became ice-covered,but it is accepted that Neoproterozoic glaciations were moreextensive than late Cenozoic ‘‘ice ages’’. Paleomagnetic evi-dence indicates that much glacial debris was deposited inlow-latitude settings (Harland, 1964; Evans, 2000; Hoffmanand Schrag, 2000, 2002; Kilner et al., 2005). Glacial episodeswere followed by rapid warming, as evidenced by depositionof thick sequences of ‘cap-carbonates’ above diamictitesdeposited by ice-rafting or by other modes of periglacial sed-imentation. These limestones and dolomites may have beendeposited very rapidly, as the warming ocean became super-saturated in carbonate.

Kirschvink (1992) identified three principal ways to testthe hypothesis. First, glacial units around the globe shouldbe more or less synchronous. Efforts continue to determinethe ages of glacial beds, and it may be several years beforewe know how many glacial episodes there were and theextent to which these were globally synchronous. It is oftendifficult to determine the age of these deposits because unitscontaining datable materials, such as interbedded ash beds


with zircons, are often not present. In this regard, thepersistence of Neoproterozoic volcanic activity in theArabian-Nubian Shield should be advantageous for deter-mining the ages of pertinent units as these are identified inthe ANS. The best modern evaluations indicate four prin-cipal icehouse episodes (Hoffman and Schrag, 2000, 2002;Condon et al., 2005) and we use the glacial episode ter-minology and numeric age constraints compiled by Mac-Gabhann (2005): �582–585 Ma (Gaskiers, also calledVaranger), �635–660 Ma (Marinoan), �680–715 Ma(Sturtian) and �735–770 Ma (Kaigas). The formallydefined base of the Ediacaran Period (�630–542 Ma) islocated at the contact of Marinoan glacial rocks and over-lying Ediacaran cap carbonates in Enorama Creek, Austra-lia (Knoll et al., 2004), thus three of the four glacial eventshappened during the Cryogenian Period (850–630 Ma) andone during the Ediacaran Period.

Geochronological studies are rapidly refining our under-standing of when major glaciations occurred, althoughpresently the Gaskiers and Marinoan glaciations are moretightly constrained than are the older the Sturtian and Kai-gas glaciations. Marinoan diamictites in Namibia are datedby U–Pb zircon techniques (ash interbedded at the top ofthe Ghaub diamictite) at 635.5 ± 1.2 Ma (Hoffmannet al., 2004). This age for the Marinoan event is supported

Fig. 1. Secular variation in carbon (A) and strontium (B) isotopic compositioNubian Shield (ANS) basement rocks (shaded) to Neoproterozoic ‘‘Snowball ESturtian glaciations. Figure modified from Miller et al. (2003).

by two U–Pb zircon ages for ash beds from just above theNantuo Tillite (2.3 m above: 635.2 ± 0.6 Ma; 9.5 m above:632.5 ± 0.5 Ma; Condon et al., 2005). In contrast, the ageof the Sturtian glaciation may have taken longer or con-sisted of multiple episodes, from �670 to �725 Ma. Glacialdeposits in Idaho, USA, are constrained using SHRIMPU–Pb zircon techniques to have occurred between709 ± 5 Ma and 667 ± 5 Ma (Fanning and Link, 2004),whereas in Oman ash beds within the Ghubrah diamictiteyielded a U–Pb zircon age of 711.8 ± 1.6 Ma (Allenet al., 2002).

Kirschvink (1992) also suggested that if the SEH wasbroadly correct, then global icehouse/greenhouse eventsshould have produced similar deposits around the globe.This is often found, in particular, where unusual carbonateunits abruptly overlie glacial successions. These are the‘‘cap carbonates’’ discussed later.

Carbon-isotopic compositions of especially carbonaterocks are crucial for characterizing and correlating thesedeposits, particularly in locales where there is no zircongeochronology. Carbon-isotope stratigraphy is uniquelypowerful for correlating Neoproterozoic calcareous sedi-ments because the greatest changes in C-isotopic variationsin Earth history occurred in Neoproterozoic time (Fig. 1),and icehouse–greenhouse deposits are accompanied by

n of shallow marine carbonates, showing the relative timing of Arabian-vents’’. Carbon-isotope excursions correspond to Gaskiers, Marinoan, and


unusual swings in the carbon-isotopic composition of car-bonate sediments. Carbonate sediments proxy for the iso-topic composition of seawater, and provide a robustrecord of the changing C-isotopic composition of the Neo-proterozoic atmosphere and hydrosphere. Neoproterozoiccarbon-isotope cycles are thought to mark when earth’s cli-mate changed from ‘icehouse’ to ‘greenhouse’ conditions,and the number of Neoproterozoic C-isotope excursionssuggest that there may be six Snowball Earth events(Fig. 1).

The basic principles that relate C-isotope excursions andSnowball Earth events reflect differing amounts of effec-tively buried dead biosphere and different contributionsfrom a homogeneous mantle. The isotopic compositionof carbon in the atmosphere and the ocean in which car-bonate forms is controlled by equilibrium between reser-voirs of inorganic and organic C, and how much of thelatter is effectively buried. The carbon-isotope balance ispreserved as relative variations between 13C and 12Cretained in carbonate sedimentary rocks, measured asd13C relative to an isotopic standard; typically PDB Creta-ceous belemnite carbonate; an arbitrary seawater proxy isset at 0&). Metabolic processes most effectively integratelight carbon into biomass, so proliferating life depletesthe ocean in 12C (and has negative d13C). This enrichesthe CO2 and bicarbonate in seawater in heavier 13C (withpositive d13C). Fractionation of C-isotopes in the atmo-sphere and hydrosphere (and thus in carbonate rocks)can become extreme if a significant proportion of thisorganic matter is buried as organisms die and are removedfrom the C-cycle, similar to what has been noted on a smal-ler scale for Cretaceous ‘black-shale’ events (e.g., Kuyperset al., 2002). Extreme fractionation is also possible forstratified oceans (mentioned later) or changes in amountsof biomass produced via photosynthetic versus chemoauto-tropic metabolic pathways (Hayes et al., 1999).

There is very active research to resolve the Neoprotero-zoic carbon isotope record and whether or not fluctuationscorrespond to glacial episodes (e.g., Halverson et al., 2005).There are several explanations for the causes of these vari-ations, particularly how these may be related to changes inthe biosphere and hydrosphere. A detailed review of thesehypotheses is beyond the scope of this paper but the inter-ested reader will find a good overview of carbon and otherisotopic systematics in Ohmoto (2004).

According to the most popular accounts (Kirschvink,1992; Hoffman et al., 1998; Hoffman and Schrag, 2002) glo-bal glaciation is thought to have decimated biologicalactivity, freeing much light carbon. This would have beenrecorded as lower d13C in carbonate during times of dimin-ished biological activity. Weakening of the biosphere isthus manifested as lighter C-isotopic composition of theatmosphere and hydrosphere, approaching the isotopiccarbon of carbon escaping from Earth’s mantle due to vol-canic activity (d13C � �6&). Global glaciation is thoughtto have ended when atmospheric CO2 increased sufficientlythat warming due to this ‘greenhouse gas’ overcame the

effect of cooling due to the high albedo of an ice-coveredworld. Destabilization of methane hydrates held in shelfsediments and massive release of methane—another green-house gas—may have also been important for ending Neo-proterozoic glacial episodes (Kennedy et al., 2001). Globalwarming led to rapid deglaciation, accelerated biologicalactivity and renewed burial of isotopically light carbon,and seawater returned to normal, heavier carbon-isotopiccompositions.

Finally, Kirschvink (1992) suggested that deepwaterdeposits of the Neoproterozoic Ocean should also recordthe extreme climatic events, particularly in the form ofbanded iron formations (BIFs). Neoproterozoic BIFs mayreflect re-oxygenation of the oceans following anoxia causedby a global ice sheet, as discussed in Section 4.4 below. Neo-proterozoic carbonate sediments were likely deposited onlyin shallow-water shelf environments because calcareousplankton did not diversify until the middle Mesozoic(Ridgewell et al., 2003). Thus, snowball event carbonates tellus little about the deep ocean. BIF may better record howNeoproterozoic climate change affected the deep ocean.

3. The Arabian-Nubian Shield and the East African

Orogen

The Arabian-Nubian Shield (ANS) outcrops around theRed Sea in NE Africa and W. Arabia as a result of upliftand erosion on the flanks of the Red Sea in Oligoceneand younger times (Fig. 2A). The ANS may be the largesttract of juvenile continental crust of Neoproterozoic age onEarth (Patchett and Chase, 2002). ANS evolution can besimplified into four stages, as shown in Fig. 3. This accom-panied a supercontinent cycle that defined Neoproterozoictectonics, beginning with the breakup of the end-Mesopro-terozoic supercontinent Rodinia in the early Neoprotero-zoic (Hoffman, 1999). ANS juvenile crust was generatedaround and within the Mozambique Ocean (Stern, 1994).Arcs and oceanic plateaux were swept up as the Mozam-bique Ocean closed. The tectonic cycle culminated in aprotracted collision between what has come to be knownas East and West Gondwana (each of which may havebeen only partially consolidated, e.g. Alkmim et al.,2001; Collins and Pisaversky, 2005), resulting in the EastAfrican Orogen (EAO) and a supercontinent ‘GreaterGondwana’ or ‘Pannotia’ at the end of Neoproterozoictime (Fig. 2B). In reconstructed Gondwana, the EAOextends from the Mediterranean (Tethys) southward alongthe eastern margin of Africa and across East Antarctica(Stern, 1994; Jacobs et al., 2003).

Exposed crust of the EAO changes dramatically alongits length. The northern EAO, the ANS, is dominated byexposures of juvenile Neoproterozoic crust, especiallygreenschist-facies supracrustal and abundant intrusiverocks. Fig. 4 presents the Nd-model age summary ofStern (2002) for the EAO and ANS. This provides anisotopic proxy for Neoproterozoic paleogeography, withthose regions characterized by Neoproterozoic model ages

Fig. 2. (A) The Arabian-Nubian Shield. Stars denote regions with evidence for Snowball Earth deposits. Modified after Miller et al. (2003). Location ofFig. 10 is shown as stars labeled BIF. (B) and (C) Paleogeographic reconstructions of the Arabian-Nubian Shield as part of the End-Neoproterozoicsupercontinent at �580 Ma, from Meert and Torsvik (2004). (B) The high-latitude Laurentia option places the present-day eastern margin of Laurentia atthe south pole adjacent to the Amazonian and Rio Plata cratons at 580 Ma. Baltica has rifted from NE-Laurentia opening the Iapetus Ocean. (C)Configuration in (B) is rotated to show an alternative configuration for the final stages of Gondwana assembly and closure of the Mawson Sea betweenAustralo-Antarctica and the rest of Gondwana.


largely forming in oceanic realms, with sedimentation andvolcanism in shallow to abyssal submarine environments,compared to pre-Neoproterozoic Nd-model age regionscharacterized by continental environments. This is an over-simplification, but Fig. 4 does emphasize the point thatthe ANS is largely juvenile Neoproterozoic crust whereasthe southern EAO mostly formed from older continentalcrust. The ANS largely escaped high-grade metamor-phism because terminal collision occurred in the south,allowing the EAO to escape northward (Bonavia andChorowicz, 1992; Abdelsalam and Stern, 1996). In con-

trast, the southern EAO (Tanzania and Madagascar) wasmore intensely deformed and metamorphosed and containsabundant granulite-facies rocks, many with pre-Neoprote-rozoic protolith ages (Kroner et al., 2003). These rocks rep-resent the intensely overprinted margins of the collidingcontinents and testify to greater thickening of the crust inthe south and correspondingly deeper erosion. The ANSand EAO evolved together, especially in their later stages,but because its early development mostly took place belowsea level, the ANS should contain a better sedimentaryrecord of events predating terminal collision.

Fig. 3. Stages in the tectonic evolution of the Arabian-Nubian Shield and the East African Orogen, modified after Stern (1994).


Formation of the ANS began as Rodinia began to disin-tegrate between 900 and 800 Ma, as inferred from the oldest(�870 Ma) juvenile Neoproterozoic rocks in the ANS(Stern, 1994) and from events in eastern Gondwana(Cawood, 2005). ANS crust was largely generated at cir-cum-Mozambique Ocean intraoceanic arc systems (Tadesseet al., 1999; Woldehaimanot, 2000). Oceanic plateaux mayalso have formed above mantle plumes within the greatocean; these would have been accreted and added to themix of juvenile crust (Stein, 2003). Juvenile arc and plateauterranes collided and were welded into larger tracts of juve-nile crust as theMozambique Ocean closed, forming arc–arcsutures, composite terranes, and, ultimately, the ANS(Johnson and Woldehaimanot, 2003). ANS juvenile crustwas trapped as the ocean closed between fragments of Eastand West Gondwanaland, ultimately nestling within the�630 Ma terminal collision zone of the EAO (Meert,2003). Convergence between fragments of E andWGondw-ana continued and theEAOwas further deformed during thelast �80 million years of the Precambrian (Veevers, 2003).Deformation included strike-slip shear zones and tectoniccollapse structures in the northern EAO (Egypt, Sudan,and northern Arabia), formation of N-trending upright tightfolds and shear zones in the central EAO (Ethiopia, Eritrea,and southern Arabia), and formation and uplift of high-grade gneisses and granulites in the southern EAO (Abdelsa-lam and Stern, 1996). The most intense collision was in thesouthern EAO, whichmust have had the thickest crust, high-est mountains, and the deepest erosion. Pannotia or GreaterGondwana began to break up almost as soon as it formed atthe end of Neoproterozoic time, The supercontinent con-tinued to shed microcontinents into especially Asia allthrough Paleozoic and early Mesozoic time, with the coreof Gondwana finally rupturing in Late Jurassic time.

4. Diagnostic evidence for Neoproterozoic glaciation and

post-glacial warming

4.1. Dropstones

A dropstone is an isolated, oversized clast in laminatedsediments that depresses the underlying laminae. Thedropstone may be draped by sediments. Most dropstonesform by debris falling from ice rafts (Fig. 5A), althoughsome Phanerozoic dropstones are transported as kelpholdfasts or by floating tree roots. Such explanations can-not explain Neoproterozoic dropstones. Dropstones couldalso potentially result from coarse material produced by ameteorite impact falling back to Earth, but such clasts arelikely to be highly shocked. Volcanic bombs could bethrown out a few kilometers from a violent eruptionand land on laminated sediments. Other than these cave-ats, Neoproterozoic dropstones indicate glaciation, andthat the ice mass carrying coarse rocks floated on water(Fig. 5A). When they are not deformed, dropstones areusually angular and marked by glacial scratches andgrooves, and they deform the underlying sedimentwhereas the overlying sediment drapes the dropstone(Fig. 5B).

Recognition of dropstones in sedimentary successions ismost convincing where there has been little deformation.Deformation in the ANS renders the identification of drop-stones much more difficult, and we know of no convincingoccurrences here. Dropstones are reported from the HuqfGroup in NE Oman, where Allen et al. (2004) note thatthe Fiq Formation contains four horizons of proximaland distal marine glacial deposits with dropstones andother evidence of ‘rainout’ of debris from icebergs in a mar-ine environment.

Fig. 4. Pre-Jurassic configuration of the East African Orogen in Africaand surrounding regions, modified from Stern (2002). Regions referred toin text: Egypt (Eg); Sudan (Su); Sinai–Israel–Jordan (SIJ); Afif terrane,Arabia (Aa); Rest of Arabian Shield (Ar); Eritrea and northern Ethiopia(En); Southern Ethiopia (Es); Eastern Ethiopia, Somalia, and Yemen (Ee);Kenya (K); Tanzania (T); Madagascar (M). Numbers in italics beneatheach region letter are the Nd-model ages. Regions of juvenile crust haveNd model ages of �1.0 Ga or less; these regions likely existed below sealevel during Kaigas (�770–740 Ma) and Sturtian (�750–700 Ma) glacia-tions. Regions with Nd model ages >1.0 Ga may or may not have beenbelow sea level at these times. The entire EAO was topographicallyelevated following collision, beginning about �630 Ma. The EAO wasprobably subject to extreme continental glaciation during Marinoan(�630–600 Ma) and glaciation. Suture labeled ‘BN’ is the Bir Umq–Nakasib suture.


4.2. Diamictites

The correct identification and interpretation of diamic-tites is critical for evaluating the record of ancient glacia-tions. Flint et al. (1960) introduced the term diamictitefor lithified, poorly-sorted, non-calcareous terrigenous sed-imentary rocks, from the Greek diamignymi meaning ‘tomingle thoroughly’. Diamictites are poorly sorted polymictconglomerates and breccias and contain a wide range ofclast sizes and shapes. These can form in many ways—for example as debris flows and as ejecta blankets from

meteorite impacts, as well as due to the actions of glaciers(Eyles and Januszcak, 2004). Tectonic and volcanic activitythat formed the ANS provided many opportunities to pro-duce diamictites without glaciers. Thus, the identificationof a sedimentary unit as a diamictite does not require theinterpretation of glacial activity, but it does focus attentionon and hopefully result in more careful scrutiny of the unit.Diamictites that result from glacial activity encompass avariety of peri- and subglacial environments, including ter-minal and lateral moraines, deposited in both marine andsubaerial environments. An abundance of angular clastshapes supports an interpretation of glacial origin, butthese may be common only in marine sedimentary environ-ments. Rounded cobbles can also result from terrestrialglaciation, because rock fragments deposited by glaciersmust be transported by melt streams from upland morainesto lowstanding basins. Diamictites deposited in marineenvironments are more likely to be preserved, simplybecause these can be more deeply buried and protectedfrom erosion than those deposited above sea level. Themost unequivocal evidence for a glacial origin of diamictiteis the identification of scratch marks or striations on clastsor recognition of dropstones, but these criteria are compli-cated in the ANS because deformation has obscured manyprimary sedimentary structures. Clast lithology and agemay be better ways of determining whether or not ANSdiamicties are glaciogenic. Those that may be glacial in ori-gin should contain a large diversity of lithologies, shapes,and sizes (polymict conglomerate or heterolithologicbreccias, depending on clast shape). Because of uncertaintysurrounding the origin of ANS diamictites and the obliter-ation of delicate sedimentary structures by deformation,breccias that contain only a single clast type (volcanic, plu-tonic, or sedimentary) are difficult to demonstrate to haveformed glacial origin. Some monomict diamictites mayhave been deposited by glacial action, but the recognitionof limited provenance makes it more likely that such depos-its formed by non-glacial debris flows. ANS diamictiteswith mostly volcanic clasts in particular are not convincingevidence of glacial activity. Identification of clasts thathave no local provenance is a strong argument for long dis-tance transport by floating ice. A potential example is theAtud Conglomerate of E. Egypt, where some granitic cob-bles yield pre-Neoproterozoic U–Pb zircon ages and areinferred to have been transported hundreds of kilometersfrom sources that now lie west of the Nile (Dixon, 1981).No breccias inferred to have formed by meteorite impactare yet reported from Neoproterozoic units of the ANS.

Neoproterozoic sediments of the Huqf Supergroup in N.Oman (J. Akhdar area) provide a valuable lesson in glacialactivity and diamictite sedimentation in a region that isnow near the ANS. The Abu Maarah Group containsare two important diamictite horizons that are associatedwith distinct glacial episodes (Le Guerroue et al., 2005a):the Ghubrah Formation (with a diamictite ash date of711 ± 6 Ma; Allen et al., 2002)), and the overlying FiqFormation (constrained between 712 and 544 Ma but not

Fig. 5. Formation of dropstones in a near-glacial aqueous environment, modified after Hladil (1991). (A) Continental glacier erodes rock and transportsthis to a lake or to the sea. Icebergs form as glacier calves and drifts with the current, carrying embedded rocks. Melting iceberg eventually releasesembedded rocks, which fall to lake or sea floor, impacting sediments and forming dropstones. (B) Results of dropstone experiments where theconsolidation of the substrate is varied, from poorly to very lithified. Note that the intensity and depth of deformation of substrate increases withdecreasing substrate lithification, as a function of the diameter of the dropstone (D). In general, the best diagnostic feature to identify a dropstone is thepenetration of underlying lamina by the clast and simultaneous lack or insignificance of compactional deflection of lamina above the clast.


yet radiometrically dated; Kilner et al., 2005). Ghubrahand Fiq formations are separated by an angular unconfor-mity and the older Ghubrah Formation is highly deformed.The Ghubrah Formation is broadly Sturtian and the FiqFormation is widely regarded as Marinoan in age (Burnsand Matter, 1993; Brasier et al., 2000; Le Guerroue et al.,2005a) , although it could be older (Kilner et al., 2005).The Ghubrah Formation is dominated by several hundredmeters of diamictite characterized by: (1) poor stratifica-tion; (2) unsorted randomly dispersed clasts of diverse size(up to 1 m; usually <10 cm) and lithology (crystalline andmetamorphic rocks, mafic and felsic volcanics and sedi-mentary rocks) that comprise about 15% of the diamictite;and (3) unsorted silty-shaly or sandy matrix. The Ghubrahis not marked by cyclical sedimentation, and lacks the shal-low water elements and thick turbidites characteristic ofthe younger Fiq (P. Allen, pers. comm., 2005). Some clastsare striated and some are interpreted as dropstones. Thediamictite is interpreted by Le Guerroue et al. (2005a,b)as due to ice-rafting. Siltstone units (up to 10 m thick)are thought to have been deposited in a marine environ-

ment during a time of reduced influence from ice-rafting(Le Guerroue et al., 2005a).

The Fiq Formation is �1.5 km thick. Diamictites prob-ably comprise 20% or less of the succession. The rest of theFiq Formation is turbidites, shales, slumped silts andshales, debris-flows, and wave-ripped sandstones. The FiqFormation contains diamictite horizons up to 30 m thick,which are somewhat stratified (clast concentrations ormatrix grain size) when proximal. The Fiq Formation isdivided into two facies associations (Leather et al., 2002;Allen et al., 2004): (1) proximal and distal glaciomarine;and (2) non-glacial gravity flow and shallow marine. Clastsizes range from 1 cm to 2 m, and some are facetted or stri-ated. Leather et al. (2002) identified seven stratigraphiccycles in the Fiq Formation and interpreted these to indi-cate when ice sheets advanced and retreated. Recently iden-tified magnetic reversals in the Fiq (and overlying Hadashcap dolomite) indicate that the glacial to interglacialclimatic transition took place over an extended timeperiod (perhaps >105–106 yr; Kilner et al., 2005). Of fur-ther climatic significance, paleolatitude estimates from


paleomagnetic measurements place the Oman region at13�, indicating tropical glaciers at sea level.

The Shuram Formation of the overlying Nafun Grouphas a major negative carbon isotope excursion that iscorrelated with the Gaskiers (Varanger) glaciation (LeGuerroue et al., 2005a; Halverson et al., 2005), althoughno glacial sediments or evidence of glacial erosion are rec-ognized. Le Guerroue et al. (2005b) now think there is nocorrelation at all between the Shuram C isotope shift andthe Gaskiers glaciation. They think that the Shuram shiftlasted much longer than any glaciation, starting at about600 Ma and ending around 550 Ma. Regardless of theduration of the Shuram C-isotope shift, the apparentabsence of an unconformity or glacial sediments sug-gests that the Gaskiers glaciation was relatively mild nearOman.

The two main glacial deposits in Oman seem to corre-spond to Sturtian and Marinoan global episodes. TheFiq (Marinoan) glaciation in Oman was characterized byrepeated advances and recessions, whereas Sturtian depos-its were less variable (P. Allen, pers. comm. 2005). If therelative thickness of diamictite beds correlates with theintensity of the glaciation that formed them, it may be thatthe Ghubrah (Sturtian) glaciation was more intense thanthe Fiq (Marinoan) glaciation in the region around Oman.At present, however, diamictite thickness is not related sim-ply to glacial intensity or duration.

4.3. Cap carbonates

Many Neoproterozoic glacial deposits are capped bylayers of pure dolostone and limestone, known as ‘cap car-bonates’ (Hoffman and Schrag, 2002). Cap carbonates aretypically thicker than underlying glacial beds, and thesedifferences may reflect different sedimentation rates. Gla-cial sediments may have been deposited over millions ofyears. Cap carbonates are traditionally thought to havebeen deposited in only a few tens of thousands of years,but recent discovery of paleomagnetic reversals in somecap carbonate sequences (Oman, Kilner et al., 2005;Amazon craton, Trindade et al., 2003) suggests a longerduration (perhaps >105–106 yr ). Cap carbonates are espe-cially paradoxical because they indicate an abrupt changefrom glacial to apparently tropical conditions, and thereis a general perception that cap carbonates reflect a globalgreenhouse climate. Cap carbonates are typically depletedin 13C in the lower beds and rebound in younger cap car-bonate beds to positive d13C values indicating biologicalfractionation becomes increasingly important.

Several physical characteristics of cap carbonates arecommonly noted. As noted by Shields (2005), cap carbon-ates are usually thin (typically <5 m, but up to 27 m thick)and uniform deposits of pale pink to buff microcrystallinedolomite, with minor siliciclastic content. They are oftenlaminated on a cm-scale, seldom preserving primary cal-citic textures, and may show graded (reverse and normal)bedding. Sheet cracks, doming and brecciation associated

with isopachous dolomite cementation are common, asare associations with high energy deposits (i.e., hummockycross-stratification and giant wave ripples; formerly inter-preted as ‘‘pseudo-tepee’’ structures). This dolostone baseis usually laterally extensive, but sometimes discontinuous,and may be overlain by transgressive shales, siltstones orthick post-cap limestones. Many localities preserve evi-dence of post-cap dolostone seafloor precipitation as sea-floor aragonite fans and barite. Stromatolitic carbonatesare often noted in the post-cap dolostone sequence. Itshould be noted that Marinoan cap carbonates are partic-ularly renowned for these attributes and there is a substan-tially smaller body of data (Prave, 1999; Hoffman andSchrag, 2002) concerning Kaigas and Sturtian capcarbonates.

The origin and environmental conditions regulatingphysical and chemical characteristics of cap carbonatesare among the more contentious aspects of competingSEH models. All models associate deglaciation withextreme increases in the alkalinity of seawater. Currentlythere are four competing models: (1) overturn of a redox-stratified ocean (Knoll et al., 1986; Grotzinger and Knoll,1995; Canfield, 1998; James et al., 2001); (2) extremechemical weathering due to supergreenhouse conditions(Kirschvink, 1992; Hoffman et al., 1998; Hoffman andSchrag, 2002; Higgins and Schrag, 2003); (3) massive oxi-dation of destabilized methane hydrates (Kennedy et al.,2001); and (4) sudden formation and gradual dissipationof a global meltwater plume that stimulated microbialmediation of carbonate precipitation (plumeworld hypoth-esis; Shields, 2005). It is beyond the scope of this review tosystematically evaluate these, but each varies significantly inthe interpretation of carbon cycle dynamics from C-isotopedata.

The diagnostic association of cap carbonates aboveglaciogenic sediments has not yet been reported from theANS, but there have been few deliberate searches thusfar. In Oman, where several examples of glacial diamictitesare documented, the older Ghubrah diamictites (Sturtian)do not have a cap carbonate, but this may have beenremoved by erosion (P. Allen, pers. comm.). As mentionedearlier, only the uppermost diamictite of the possiblyMarinoan-aged Fiq Formation has a cap carbonate. Thisis the dolostone of the <15 m thick Hadash Formation(Allen et al., 2004). Another candidate is a thick sequenceof laminated dolomites and stromatolitic limestones ofthe Tambien Group, N. Ethiopia (Beyth et al., 2003; Milleret al., 2003). These have clear petrographic, textural, andisotopic affinities with cap carbonates but have not beenfound to overly glaciogenic diamictites, as discussed in alater section.

Carbonate sedimentary rocks are not common in theANS, so when these are found investigators should exam-ine the basal contact interval for previously recognizedcap carbonate features (e.g., Shields, 2005) as well asunderlying strata for evidence of glaciation. However,investigators should also recognize the comparative lack


of cap carbonate documentation for pre-Marinoansequences that may be more prominent in the ANS.

4.4. Banded iron formations

BIF was commonly deposited during Late Archean andPaleoproterozoic time but disappeared as oxygen concen-trations in the atmosphere and oceans increased and dis-solved Fe was removed from seawater (Rouxel et al.,2005). Rising oxygen concentrations oxidized Fe2+ dis-solved in seawater into insoluble Fe3+, which precipitatedand accumulated on the seafloor to form BIF. The oxida-tion of seawater Fe2+ was completed in Paleoproterozoictime, so that BIFs are missing from the Mesoproterozoicrecord. BIF reappeared in Neoproterozoic times in associ-ation with Snowball Earth events. Most Neoproterozoicoccurrences formed during the Sturtian ice age; only onecase is documented from a possible Marinoan-age glacio-genic sequence (Shields, 2005; Proust and Deynoux, 1994).

There are two general categories of BIF: Superior-typeand Algoma-type. Superior-type BIF is associated withshelf sediments (quartzite, marble, etc.) whereas Algoma-type BIF is associated with volcanic rocks and immaturesediments. Superior-type BIF require a global or at leastregional change in water chemistry to precipitate Fe,whereas Algoma-type BIF may reflect more local oceano-graphic conditions and sources of Fe. Archean BIFs aremostly Algoma-type, whereas Paleoproterozoic BIFs aremostly Superior-type. Neoproterozoic BIF can have affini-ties to either Algoma- or Superior-type. A more detailedreview of BIF can be found in Trendall (2002).

Neoproterozoic BIF are an important argument forSEH but are controversial. Kirschvink (1992) suggested

Fig. 6. Model for formation of Neoproterozoic banded iron formations (BIFseawater from mixing with atmosphere, cutting off the source of oxygen. Oxidresult that seawater becomes anoxic and reducing. Iron introduced as Fe2+ at mof Fe2+ in seawater. (B) Deglaciation: ocean ventilation. Melting of ice allowoxygen concentrations in seawater oxidizes Fe2+ dissolved in seawater to Fe3+

Modified from Fig. 8 in http://www-eps.harvard.edu/people/faculty/hoffman/s

that covering the oceans with ice could isolate the deepoceans from the atmosphere and thus lead to anoxia inthe deep ocean. Deep waters would become reducing, sothat Fe supplied from seafloor hydrothermal ventsremained as Fe2+ in solution. When the ice sheets melted,the supply of oxygenated water to the deep sea resumedand Fe2+ in solution oxidized to insoluble Fe3+, which pre-cipitated out as BIF. This model is shown schematically inFig. 6. A second explanation for Neoproterozoic BIF callson glaciation of Red Sea rift-type basins. Deep, Fe-chargedanoxic brines in such basins would have precipitated Feoxides on being mixed with ‘‘normal’’ seawater as a resultof glacially driven thermal overturn (Young, 2002). Thiskind of argument is especially convincing for Neoprotero-zoic Superior-type BIF and is less convincing for Neopro-terozoic Algoma-type BIF, which is the kind of BIF mostlikely found in the ANS. Nevertheless, a link with large-scale glaciation seems required if evidence for glaciation(e.g., diamictite, dropstone) is found in association withNeoproterozoic BIF of either category.

Trendall and Blockley (2004, p. 421) warn: ‘‘The Snow-ball Earth hypothesis is at an early stage of testing, and theemphasis placed by some authors. . .on the relationshipbetween rift-related mafic volcanism and some Neoprote-rozoic [BIFs] indicates that the evidence for a purely cli-matic control of their deposition is not yet definitive.’’

4.5. Paleomagnetic evidence

For the Snowball Earth hypothesis it is not only crucialto demonstrate which deposits are best interpreted as gla-cial, but also to show evidence that such deposits wereformed at low to equatorial latitudes (Evans, 2000). This

). (A) Snowball Earth: anoxic ocean. Ice covering ocean surface isolatesation of organic matter consumes oxygen dissolved in seawater, with theid-ocean ridge hydrothermal vents remains in solution, causing a buildup

s mixing of atmosphere and seawater, re-oxygenating seawater. Increased. This forms insoluble iron-oxides and is deposited as BIF on the seafloor.nowball_paper.html.

http://www-eps.harvard.edu/people/faculty/hoffman/snowball_paper.html


is best demonstrated with careful paleomagnetic measure-ments. Unfortunately, there have not yet been any paleo-magnetic studies of potential Snowball Earth deposits inthe ANS.

5. Expected manifestations of Neoproterozoic glaciations inthe ANS and EAO

Our understanding of the tectonic evolution of the EAOand ANS (Fig. 3) has implications for how the waxing andwaning of Neoproterozoic Snowball Earth episodes shouldbe preserved. Uncertainty about the number and timing ofSnowball Earth episodes has already been noted, but thereis nevertheless agreement that two major episodes occurredprior to �700 Ma and two after �630 Ma. The older epi-sodes (Kaigas and Sturtian) occurred while ANS crustwas still forming as arcs, back-arc basins, and oceanic pla-teaus around and within the Mozambique Ocean (Fig. 7).There should have been many submarine sedimentarybasins available to collect the distinctive debris producedby Kaigas and Sturtian glaciations and subsequent warm-ing and reoxygenation episodes. Some of these may haveformed in shallow water of a few hundred meters depth,on continental shelves, atop oceanic plateaus, and aroundisland arcs. Other deposits should have formed at abyssaldepths, as deep as the 2500–5000 m characteristic of themodern seafloor, associated with backarc basins and intra-oceanic forearcs, and on the floor of the MozambiqueOcean itself. If the Snowball Earth episodes prior to�700 Ma affected the Mozambique Ocean and its periph-ery, evidence should be preserved in the ANS. These depos-its would have been deformed during later accretion andcollision events, but distinctive ‘Snowball Earth’ sedimen-tary deposits should still be recognizable in parts of theANS.

Fig. 7. Expected interactions of Snowball Earth episodes with evolving ANSfrom Hoffman and Schrag (2002). Column on left refers to events discussed in

It is less likely to find sedimentary deposits of Marinoanand Gaskiers Snowball Earth events. The period after�650 Ma very likely witnessed increasing relief in theANS and EAO, as collisions between various fragmentsin the Mozambique Ocean occurred, culminating in termi-nal collision between E. and W. Gondwana (Fig. 7). Mostof the ANS was probably above sea level by �630 Ma, withthe highest relief in the southern EAO. At the end of theNeoproterozoic, the EAO may have rested near the southpole (Fig. 2B), so if there were global glaciations, the regionis likely to have been covered with a thick continental icesheet. It is possible that glacial deposits of the Gaskiersand Marinoan episodes could be preserved in deep grabenaround the margins of the ANS, such as that preserving theHuqf Supergroup in Oman. Marinoan and Gaskiers depo-sits could also be preserved in continental shelf deposits onthe northern flank of the End-Neoproterozoic superconti-nent, such as may exist beneath Israel.

6. The record of glaciation in the ANS

Diamictites and other likely examples of SnowballEarth deposits preserved in the ANS are discussed below.These are presented in terms of the four episodes iden-tified by MacGabhann (2005): 735–770 Ma (Kaigas),680–715 Ma (Sturtian), 635–660 Ma (Marinoan), and582–585 Ma (Gaskiers). This grouping and the age assign-ments are likely to change as studies advance aroundthe globe, but these episodes provide a useful frameworkfor the following observations. This discussion drawsheavily our understanding of the Huqf Group in Omanfor indications of how glacial episodes are likely to bemanifested in the ANS. There is also a rich record inNW Africa, summarized by Evans (2000) that is alsoinstructive.

and EAO. Column on right refers to Snowball Earth episodes generalizedthe text.


6.1. Evidence for Kaigas (�735–770 Ma) glaciation

The best evidence for the earliest glacial sedimentationin the ANS is found in deposits on the southern marginof the Bi’r Umq–Nakasib Suture Zone of Sudan and Ara-bia (Fig. 4). This suture zone can be traced for more than600 km, from the central Arabian Shield almost to the Nileand is the major suture separating the northern and south-ern ANS (Johnson et al., 2003). Diamictites are found onthe southern flank of the suture at two widely separatedlocations. In Arabia, Johnson et al. (2003) report that thebase of the �770 Ma Mahd Group unconformably overliesthe 816 ± 3 Ma Dhukhr batholith, indicating a significant

Fig. 8. Photographs of outcrops in the ANS with evidence for Kaigas (A,B) aMahd Group basal tillites on 806 Ma Dhukhar batholith, Saudi Arabia. Fingerof granitic rocks in dark matrix. (C) Tambien Group tillite, Negash synformEgypt). Angular dropstone is composed of quartz porphyry. (E) Atud diamiimmature clastics and BIF and consists of blocks up to 2 m of quartzite, graniteScientists point at three of these blocks. (F) BIF in Wadi Dabbagh, Egypt (hacarbonate bands.

episode of erosion between �770 and �816 Ma (Fig. 8A).This is the oldest unconformity documented within theANS. The Mahd Group rests directly on this unconformityand while dominated by volcanic rocks, its base is definedby a 1–5 m thick diamictite. The diamictite is matrixsupported, with a dark-grey, immature, arkosic matrixthat contains abundant, angular to sub-angular clasts (upto 30 cm wide) of granitic and felsic volcanic rocks(Fig. 8B). Johnson et al. (2002) noted that this diamictitewas ‘‘. . .conceivably deposited during a Neoproterozoicglacial event.’’

A diamictite of similar age is found in the Meritri Groupin the Sudanese sector of the Bi’r Umq–Nakasib Suture, a

nd Sturtian (C–E) Snowball Earth events. (A) Unconformity of �770 Mapoints to unconformity. (B) Mahd Group basal tillite, note angular clasts, N. Ethiopia. (D) Atud conglomerate dropstone, Wadi Khuda area (SEctite at Wadi Kareim, Egypt. The diamictite sits stratigraphically below, granodiorite, metavolcanics, and pebble conglomerate in schistose matrix.nd lens for scale). Dark, hematite-rich layers are interbedded with thinner


few km west of Port Sudan. Abdelsalam and Stern (1993)infer that sediments along the SE side of the suture mani-fest a deformed passive margin sequence. This sequencebegins with the Arbaat volcanic Group, which is succeededstratigraphically and structurally sediments of the Salatiband Meritri groups. The Arbaat volcanic Group yields aU–Pb zircon age of 790 ± 2 Ma. Following suture-relateddeformation, the Arbaat, Salatib, and Meritri groups wereintruded by granitic plutons as old as 754 ± 3 Ma (Sternand Abdelsalam, 1998). This constrains the age of theSalatib and Meritri sediments to younger than 790 Maand older than 754 Ma.

The Salatib Group consists of intercalated rhyolite, con-glomerate, mudstone, wacke, quartzite, and carbonate sed-iments. The Meritri Group consists of (from oldest toyoungest): conglomerate, lithic wacke, and interbeddedlimestone, red sandstone, and felsic tuff. Abdelsalam andStern (1993) infer an original thickness of �2 km. The con-glomerate is polymict and matrix supported, and is betteridentified as diamictite. Clasts vary greatly in size, from afew cm up to a meter or more, and in composition. About50% of the clasts are plutonic (granite, granodiorite, dio-rite), �35% are volcanic (rhyolite and ignimbrite), and�15% are sedimentary (carbonates and subordinate clasticrocks). This diamictite is succeeded by a lithic wacke withsedimentary structures indicating transport from SE to

Fig. 9. Location of Tambien Group exposures in Ethio

NW. Given the similarity of ages of Meritri and MahdGroup diamictites, these may be correlatable and provideevidence for the Kaigas glaciation in the ANS. The MeritriGroup diamictite should be investigated by a sedimentolo-gist with appropriate expertise.

6.2. Evidence for Sturtian (�680–715 Ma) glaciation

Evidence for Sturtian glaciation relatively close to theANS is found in the Huqf Supergroup of SE Oman, wherethe basal Ghubrah Formation contains thick glaciogenicdiamictite. Tuffaceous wackes interbedded with the diamic-tite yielded a U–Pb zircon age of 723 +16/�10 Ma (Brasieret al., 2000). Less well-dated evidence of broadly Sturtian‘Snowball Earth’ deposits are found in Ethiopia, Eritrea,Egypt, and northern Arabia. In Ethiopia, a deformedmetasedimentary unit known as the Tambien Groupcontains evidence of a Sturtian glaciation (Fig. 9). TheTambien Group is mostly carbonate, but in the Negashsynform it consists of a thick section of carbonates cappedby a distinctive polymict diamictite interpreted to be glacialin origin (Miller et al., 2003; Beyth et al., 2003; Fig. 8c). Inthin section, the diamictite contains clasts of felsic volca-nics, fine-grained carbonates, low-grade semipelitic metase-diments. In the field we also saw red granite, blacklimestone, pegmatite quartz and chert clasts up to 5 cm

pia and Eritrea (modified after Beyth et al., 2003).


in greatest dimension. The age of the Tambien Group isconstrained by the fact that it overlies �800 Ma metavolca-nics and syntectonic granitoids and is intruded by�610 Ma granites. This age constraint is consistent withthe inference that Tambien Group sediments weredeformed by the �630 Ma terminal collision to form theEAO. Within the 800–610 Ma window, Miller et al.(2003) found that Sr- and C-isotopic compositions ofTambien Group carbonates are most consistent with anage range of 720–750 Ma, broadly corresponding to theSturtian glaciation. A cap-carbonate has not been foundabove the diamictite. Stratigraphic equivalents of theTambien Group can be expected to exist in southernArabia but these have not yet been identified.

Sedimentary units in the Eastern Desert of Egypt andNW Saudi Arabia may record evidence of the SturtianSnowball Earth event in the northernmost ANS. The evi-dence from this region consists of diamictite and BIF.The diamictite is known as the Atud conglomerate in Egyptand as the Nuwaybah Formation (Zaam Group) in Arabia.BIF is distributed throughout the Central Eastern Desertof Egypt and is also found in the Silasia Formation inNW Arabia (Fig. 10). Atud conglomerate and BIF are partof a metasedimentary succession associated with ophiolites(Stern, 1981; Stern et al., 2004). The ophiolite at WadiGhadir has been dated by zircon evaporation techniquesat 746 ± 19 Ma (Kroner et al., 1992). Ophiolite, BIF, anddiamictite represent an oceanic assemblage that maypreserve evidence of deep marine conditions during theSturtian glaciation. Ophiolite and overlying metasedimentswere similarly deformed and then intruded by syntectonicgranodiorites dated by Rb-Sr whole rock techniques at

Fig. 10. (A) Location of Neoproterozoic BIF in the Central Eastern Desert of E1984); (B) BIFs of the Sawawin District, N. Saudi Arabia. Location and extentsLocation of maps shown as stars labelled ‘BIF’ in Fig. 2.

674 ± 13 Ma (Stern and Hedge, 1985), and these ages con-strain the Atud conglomerate and BIF to broadly belong tothe Sturtian episode.

The Atud conglomerate is only recognized in easternEgypt, where it can be found between �26�N and 22�N.Its clasts are poorly sorted, polymict, and matrix sup-ported. Clasts are generally subrounded and range in sizeup to a meter across. It is a distinctive unit because its clastsare quite different than the ensimatic assemblages thatcharacterize the Eastern Desert, and include grey quartzite,arkose, felsic metavolcanics, granodiorite, and minor darkgrey marble. This is not a formal stratigraphic name, andwe propose that the unit is better referred to as the ‘Atuddiamictite’. Geochronologic data support the inference thatAtud diamictite clasts sample much older rocks than areexposed in the Eastern Desert of Egypt and so must havebeen transported some distance. Two granitic cobbles fromthe NW of Marsa Alum (also referred to as the WadiMobarak medisedimentary unit) yielded highly discordantU–Pb zircon upper intercept ages of 1120 and 2060 Ma(Dixon, 1981). Dixon (1979) obtained a discordant U–Pbzircon upper intercept of 2.3 Ga for a granitic cobble fromAtud conglomerate outcrops west of Quesir. Pre-Neoprote-rozoic basement is unknown in Egypt east of the Nile, andDixon (1981) concluded that these clasts were derived fromolder crust to the west or south, perhaps from the SaharanMetacraton (Abdelsalam et al., 2002). Dixon (1979) sug-gested this material was transported such great distancesby ice rafting, a conclusion that is consistent with occa-sional dropstones (Fig. 8D).

Diamictite is also found in the Nuwaybah Formation(Zaam Group) of NW Arabia (Davies, 1985). This unit

gypt; X’s mark approximate locations of major deposits (Sims and James,of major deposits are shown as dark lines (modified after Goldring, 1990).


has not been studied in detail but appears similar in thefield to and is probably correlative with the Atuddiamictite.

BIF of broadly Sturtian age is found in the Central East-ern Desert of Egypt and in NW Saudi Arabia (Fig. 8E andF). Arabian and Egyptian BIF formed in a single basin andwere separated by the Cenozoic opening of the Red Sea.BIF in Egypt is found in the Central Eastern Desert,between latitudes 25�15 0 and 26�30 0N, where 15 occur-rences are known (Fig. 10A). Egyptian BIF occurs as fairlyregular bands interbedded with metasediments and meta-volcanics in a zone that originally had a stratigraphic thick-ness of 100–200 m, within which the aggregate thickness ofBIF is about 10–20 m (Sims and James, 1984). The BIF-bearing sediments are associated with metavolcanics rocksand are intruded by metadiabase sills. There is controversyregarding how the Egyptian BIFs formed, although theseideas were mostly developed prior to the Snowball Earthhypothesis. Kamel et al. (1977) advocated an effusive-marine sedimentary mode of formation for Wadi Kareimiron ores. Sims and James (1984) suggested that BIFformed as chemical precipitates during lulls in dominantlysubaqueous, calc-alkaline volcanism, apparently within anintraoceanic island-arc environment.

BIFS in the Midian region of NW Saudi Arabia occupya smaller region than do their Egyptian counterparts(Fig. 10B). Arabian BIF occurs within the Silasia Forma-tion, which, like the Egyptian section, is associated withmetavolcanic rocks. The exposed thickness of the SilasiaFormation is estimated to be about 1160 m in the referencearea of Wadi Sawawin. Also similar to the Egyptian sec-tion, the Silasia Formation is intruded by metadiabase sills.It is also intruded by plutonic rocks of the Muwalylih suite,dated by U–Pb zircon techniques at 710–725 Ma (Hedge,1984). Johnson (2004) suggested on this basis that the Sil-asia Formation BIF could have been deposited in associa-tion with Sturtian glaciation.

Both Egyptian and Arabian BIFs are strongly deformedand metamorphosed to the greenschist facies. These oresare similar, mostly oxide facies, interbedded hematite andjasper, and contain 40–46% Fe (Sims and James, 1984;Goldring, 1990).

We infer that ANS BIFs formed about the same time asthe Sturtian glaciation, or shortly afterwards. Goldring(1990) agreed with Sims and James (1984) that the MidianBIFs were Algoma-type deposits, but also suggested thatthe iron was precipitated as a result of oxidation of ferrousiron in water by oxygen evolved during photosynthesis byalgae. The identification of algal fossils in the EgyptianBIFs (El-Habaak and Mahmoud, 1995) supports the inter-pretation that biological activity may have been importantfor multiple episodes where a marine environment that wasrich in Fe2+ was converted to an oxygenated environmentprecipitating Fe3+.

BIF has only been reported from the northernmost partof the ANS, as discussed above. Other components of theANS should have been deep basins below sea level during

the Sturtian glacial episode, so it is puzzling why BIF is notmore common in the ANS if formation was a synchronous,deep sea expression of a ‘‘hard’’ Neoproterozoic SnowballEarth.

6.3. Evidence for Marinoan (�635–660 Ma) and Gaskiers(�582–585 Ma) glaciations

By about 630 Ma, collision had advanced sufficientlythat much of the ANS had probably risen above sea level.Marinoan and Gaskiers Snowball Earth episodes, if pres-ent, are likely to have been manifested as continental glaci-ations, perhaps continental ice sheets. These would havebeen powerful agents of erosion and could have rapidlyreduced relief as the EAO mountains grew.

Garfunkel (1999) identified the ‘Main Erosion Phase’ inthe northern ANS, which he suggested cut 8–14 km deep atabout 600 Ma. This is also consistent with evidence from40Ar/39Ar studies of micas for rapid cooling (and uplift)at �600 Ma (Cosca et al., 1999). Similarly, a major phaseof erosion identified in the NE part of the Arabian Shieldduring the interval 615–585 Ma was inferred to result fromepeirogenic uplift (Cole, 1988). Most explanations for�600 Ma exhumation focus on tectonic unroofing (Al-Husseini, 2000; Blasband et al., 2000). We suggest thatMarinoan glaciation may have also been responsible formuch of this deep erosion. Unroofing farther south, inSudan and southern Egypt and Sudan, may have occurredabout 570 Ma, perhaps related to Gaskiers glaciation(Bailo et al., 2003).

Certainly it is possible that some beveling occurred atthe base of thick continental ice sheets, but this possibilityhas not been widely explored in the literature, largelybecause unequivocal evidence for glaciation of the appro-priate age has not been found. The greatest erosion isexpected to have occurred where relief was highest, in thesouthern EAO, but we do not yet understand when andhow this region was beveled.

Evidence that pertains to the unroofing puzzle may alsobe preserved in post-amalgamation basins of the northernArabian Shield (Johnson, 2003). Rocks of the JurdhawiyahGroup and Hibshi Formation were deposited between 640and 620 Ma in partly fault-controlled basins, which couldhave existed at the time of Marinoan glaciation, but thereis no reported evidence for glacial deposits in these basins.The Jurdhawiyah and Hibshi basins closed and invertedduring subsequent north–south shortening and north-and south-vergent reverse faulting. The Jibalah Groupwas deposited in isolated, pull-apart basins caused bystrike- and dip-slip movements of the Najd fault system.A study of �3 km thick section of conglomerate, limestone,sandstone, and shale in the Jifn Basin (NE Arabian Shield)was reported by Kusky and Matsah (2003). They con-strained its age (by U–Pb zircon techniques) to lie between625 ± 4 Ma and 577 ± 5 Ma, so the Jibalah Group in theJin Basin could have been deposited during Gaskiers glaci-ation. Kusky and Matsah (2003) show a photograph of a


‘‘possible dropstone’’ (their Fig. 7D) but do not explore thesignificance of this in detail. Further studies are needed toestablish whether or not evidence for Snowball Earthevents is preserved in Arabian post-accretionary basins,and especially to constrain their ages as tightly as possible.

There are also sediments in the northern ANS that couldrepresent glacial deposits, or at least fluvial reworking ofsediments from Marinoan glaciers. The Saramuj conglom-erate of Jordan and the Hammamat Group of NE Egyptare both about 600 Ma old (Jarrar et al., 1993; Wilde andYoussef, 2002) and perhaps were deposited as Marinoanglaciation waned. Jarrar et al. (1991) interpreted a highvelocity, braided stream/alluvial fan system. Jibalah Groupsediments could have a similar origin. We speculate that allof these coarse sediments could be periglacial tillites ofMarinoan age, reworked by meltwater streams as Mari-noan glaciers receded but further investigation is neededto confirm or refute this suggestion.

The Zenifim Formation, found only in boreholes fromthe subsurface of Israel, Jordan, and Sinai, may be anothermanifestation of Marinoan or Gaskiers glaciation. It is>2500 m thick and consists of immature arkose-dominatedclastics and conglomerates associated with alkaline volca-nics (Weissbrod and Sneh, 2002). Recanati (1986) reporteda K/Ar age of 606 ± 9 Ma for an igneous intrusion intoZenifim sediments, implying that the Zenifim Formation isolder than this and supporting an interpretation that itwas mostly deposited during Marinoan time. It will be diffi-cult to prove on the basis of drill core that these deposits areor are not sedimentary deposits associated with a Marinoanglacial episode, but the possibility should be considered.

There is less evidence to support an important role forGaskiers (�600–570 Ma) glaciation in the ANS. The ‘MainErosion Phase’ happened before Gaskiers time, and conti-nental sediments deposited �600 Ma, such as the Hamma-mat Formation of Egypt and Saramuj conglomerate ofJordan, have not been removed. This is also consistent withthe record of glaciations preserved in the Huqf Supergroupof Oman, discussed above. Garfunkel (1999) infers modest(1–2 km) erosion of the northern ANS between 600 Maand the beginning of Cambrian time. The �560–540 MaElat Conglomerate of southern Israel may have beendeposited during Gaskiers time, on a deeply dissected reliefthat suggests sea level was quite low, perhaps as a result ofGaskiers glaciation elsewhere. The Elat conglomerate con-tains clasts as large as 1.5 m (Weissbrod and Sneh, 2002),and the possibility of a glacial or periglacial origin is wor-thy of further study.

One important observation is the vast peneplain thatNorth Africa and Arabia, which formed in multiple stagesover �100 million years following terminal collisionbetween E. and W. Gondwana and prior to deposition ofCambrian marine sediments. This represents a continent-scale erosional unconformity, which can be traced fromMorocco in the west to Oman in the east (Avigad et al.,2003). There appears to be a sharply beveled surface belowthe oldest Phanerozoic sediments all across North Africa

and Arabia, except for local monadnocks and where tecto-nism has occurred.

This extraordinary surface must have an extraordinaryorigin. The wearing down of orogenic relief of Arabian-Nubian Shield had to be sufficient to yield a uniform, N-sloping surface that permitted Cambro-Ordovician streamsto flow north across it, as indicated by north-flowing paleo-current directions from the Wajid sandstone in southernArabia (Dabbagh and Rogers, 1983). In order to explain1.1–1.2 Ga detrital zircons in Cambrian quartz arenites inIsrael (Avigad et al., 2003), the headwaters of Cambro-Ordovician drainage may have reached as far south asmodern Tanzania, the northernmost limit of crust of thisage (Kroner et al., 2003). Alternatively, till could have beenglacially transported from southern Africa at least partways to the north and later reworked by streams.

Deep erosion and peneplanation involved at least twoepisodes of erosion. The final cutting of the peneplainoccurred in early Cambrian time, because the peneplainin southern Israel truncates dikes as young as 532 Ma(Beyth and Heimann, 1999). The final cutting of the pene-plain during early Cambrian time was not glacial, but asso-ciated with a warm and humid climate, as indicated bythick laterite immediately below the peneplain.

In conclusion, the geological record may be taken tosupport, if indirectly, an important role for Marinoan gla-cial erosion of the Arabian-Nubian Shield but much lessevidence in support of Gaskiers glaciation. This is consis-tent with the record preserved in Oman sediments, dis-cussed above. How the basal Cambrian peneplain formedhas not been well studied, and the possible role of glacia-tion in its formation needs to be considered further. N.African and Arabian geologists could contribute by initiat-ing field research programs to characterize this unconfor-mity in their regions.

7. Conclusions

The Snowball Earth hypothesis provides a valuable newperspective on the evolution of the Arabian-Nubian Shield,and new opportunities for African scientists to contributeto our understanding of the Earth system. The effects ofNeoproterozoic glaciations can also provide age-diagnosticstratigraphic markers, which are rare in ANS supracrustalsequences. Tectonic evolution must be considered whenconsidering these effects, because—like most orogenicbelts—the ANS evolved as Neoproterozoic time progressedfrom a mostly marine realm to a mountainous continentalenvironment. The transition occurred �630 Ma, about thetime of Marinoan glaciation. Sedimentary deposits of gla-cial episodes older than �630 Ma—Kaigas and Stur-tian—are recognized in the ANS. There is evidence ofdiamictites and BIF, and possible, but yet unsubstantiated,cap carbonates. Kaigas-aged diamictites of apparent gla-cial origin are preserved on the south flank of the BirUmq–Nakasib suture, both in Arabia and in Sudan. Sedi-mentary evidence for Sturtian glaciation appears to be


widespread and is found in diamictites of the TambienGroup in Ethiopia, Atud diamictites in Egypt, Nuwaybahdiamictites in Arabia, and banded iron formations in Egyptand Arabia. This sedimentary evidence for a strong episodeof Sturtian glaciation is also consistent with the presence ofthick glaciogenic diamictites of the Ghubrah Formation inOman.

Evidence for Marinoan and Gaskiers glaciations is lessclear for the ANS, partly because the ANS was above sealevel during this time. The sedimentary record in Omansuggests that Marinoan glaciation was important region-ally but that the Gaskiers glaciation was not. Marinoanglaciation may have caused the deep erosion of the ANSthat occurred about 600 Ma ago, providing a surface thatevolved over the next 60 million years or so into the conti-nental-scale peneplain beneath basal Cambrian strata.

The ANS also provides opportunities for studying theeffects of Neoproterozoic climate change on the deep, openocean. Analogy with modern seafloor indicates that ANSophiolites formed at depths of 2–3 km below sea leveland the overlying pelagic sediments should record chemicalproducts of these deep waters. There is opportunity forgeochronologists and sedimentary geochemists to worktogether to date the ophiolites and interpret the chemicalmessage preserved in overlying sediments.

Acknowledgment

We are grateful for assistance in the field from scientistsat Mekelle University and Ezana Minerals Corporation inMekelle, Ethiopia, especially Solomon Gebresilassie,Kurkura Kabeto, Dirk Kuster, and Kiros Mehari. We alsoappreciate comments from Erwan Le Guerroue and PhillipAllen (ETH, Zurich), Peter Johnson (SGS, Jeddah), andJoe Meert (U Florida). The comments of two anonymousreferees and editor Eriksson are greatly appreciated as well.This work is supported by USA–Israel Binational ScienceFoundation (BSF) grant no. 2002337. This is UTD Geo-sciences contribution number 1068.

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