Aberystwyth University

Glacitectonic deformation in the Chuos Formation of northern Busfield, Marie E.; Le Heron, Daniel P.

Published in: Proceedings of the Geologists' Association DOI: 10.1016/j.pgeola.2012.10.005 Publication date: 2013 Citation for published version (APA): Busfield, M. E., & Le Heron, D. P. (2013). Glacitectonic deformation in the Chuos Formation of northern Namibia: Implications for neoproterozoic ice dynamics. Proceedings of the Geologists' Association, 124(5), 778- 789. https://doi.org/10.1016/j.pgeola.2012.10.005

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Download date: 26. Sep. 2021 Elsevier Editorial System(tm) for Proceedings of the Geologists' Association Manuscript Draft

Manuscript Number: PGEOLA-D-12-00028R1

Title: Glacitectonic deformation in the Chuos Formation of northern Namibia: implications for Neoproterozoic ice dynamics

Article Type: Special Issue: Glacitectonics

Keywords: Neoproterozoic; glacitectonism; ductile deformation; Snowball Earth; Otavi Mountainland

Corresponding Author: Miss Marie Elen Busfield,

Corresponding Author's Institution: Royal Holloway

First Author: Marie Elen Busfield

Order of Authors: Marie Elen Busfield; Daniel P Le Heron

Abstract: The Chuos Formation is a diamictite-dominated succession of Cryogenian age, variously interpreted as the product of glaciomarine deposition, glacially-related mass movement, or -related sediment remobilisation in a non-glacial environment. These interpretations have wide ranging implications for the extent of ice cover during the supposedly pan-global Neoproterozoic icehouse. In the Otavi Mountainland, northern Namibia, detailed analysis of soft sediment deformation structures on the macro- and micro-scale support glacitectonic derivation in response to overriding ice from the south/south-east. Overall, the upward increase in strain intensity, predominance of ductile deformation features (e.g. asymmetric folds, rotational turbates and necking structures, clast boudinage, unistrial plasmic fabrics) and pervasive glacitectonic lamination support subglacial deformation under high and sustained porewater pressures. In contrast, soft sediment structures indicative of mass movements, including flow noses, tile structures, and basal shear zones, are not present. The close association of subglacial deformation, abundant ice-rafted debris and ice-contact fan deposits indicate subaqueous deposition in an ice-proximal setting, subject to secondary subglacial deformation during oscillation of the ice margin. These structures thus reveal evidence of dynamic grounded ice sheets in the Neoproterozoic, demonstrating their key palaeoclimatic significance within ancient sedimentary successions. *Manuscript Click here to view linked References

1 Glacitectonic deformation in the Chuos Formation of northern Namibia: 1 2 3 2 implications for Neoproterozoic ice dynamics 4 5 6 3 7 8 9 1 1 10 4 Marie E Busfield * & Daniel P Le Heron 11 12 13 5 1Department of Earth Sciences, Royal Holloway, University of London, Egham, Surrey, 14 15 16 6 TW20 0EX 17 18 19 7 *Corresponding author. E-mail: [email protected] 20 21 22 8 Abstract 23 24 25 9 The Chuos Formation is a diamictite-dominated succession of Cryogenian age, variously 26 27 10 interpreted as the product of glaciomarine deposition, glacially-related mass movement, or 28 29 11 rift-related sediment remobilisation in a non-glacial environment. These interpretations have 30 12 wide ranging implications for the extent of ice cover during the supposedly pan-global 31 32 13 Neoproterozoic icehouse. In the Otavi Mountainland, northern Namibia, detailed analysis of 33 34 14 soft sediment deformation structures on the macro- and micro-scale support glacitectonic 35 36 15 derivation in response to overriding ice from the south/south-east. Overall, the upward 37 38 16 increase in strain intensity, predominance of ductile deformation features (e.g. asymmetric 39 40 17 folds, rotational turbates and necking structures, clast boudinage, unistrial plasmic fabrics) 41 18 and pervasive glacitectonic lamination support subglacial deformation under high and 42 43 19 sustained porewater pressures. In contrast, soft sediment structures indicative of mass 44 45 20 movements, including flow noses, tile structures, and basal shear zones, are not present. The 46 47 21 close association of subglacial deformation, abundant ice-rafted debris and ice-contact fan 48 49 22 deposits indicate subaqueous deposition in an ice-proximal setting, subject to secondary 50 23 subglacial deformation during oscillation of the ice margin. These structures thus reveal 51 52 24 evidence of dynamic grounded ice sheets in the Neoproterozoic, demonstrating their key 53 54 25 palaeoclimatic significance within ancient sedimentary successions. 55 56 57 26 58 59 60 1 61 62 63 64 65 27 1. Introduction 1 2 3 28 The concept of a Neoproterozoic icehouse has remained contentious since its inception in the 4 5 th 6 29 early 19 century (‗die Eiszeit‘ of Agassiz, 1837: cf. Allen and Etienne, 2008), with renewed 7 8 30 deliberation in recent years following proposal of the ‗snowball Earth‘ hypothesis 9 10 11 31 (Kirschvink, 1992; Hoffman et al., 1998; Hoffman and Schrag, 2002). This hypothesis has 12 13 32 centred on the recognition of broadly age-equivalent diamictite-dominated successions on 14 15 33 each continent, which are argued to be glaciogenic in origin (Hoffman et al., 1998; Hoffman 16 17 18 34 and Schrag, 2002). Many diamictites are sharply overlain by dolomotized carbonates, 19 20 35 interpreted as the record of rapid post-glacial climatic recovery (Shields, 2005). Compared to 21 22 23 36 younger icehouse intervals, diagnostic glacial indicators, including striated and faceted clasts, 24 25 37 subglacially striated pavements and extrabasinal clast assemblages, are notably scarce in the 26 27 28 38 Neoproterozoic (Etienne et al., 2007), and rarely occur together in any one glacial succession. 29 30 39 Consequently, Neoproterozoic diamictites have been argued to represent non-glacial, syn- 31 32 40 tectonic sediment gravity flows (e.g. Eyles and Januszczak, 2004, 2007), associated with 33 34 35 41 widespread rift activity during break-up of the supercontinent. 36 37 38 42 In Quaternary studies, detailed analysis of soft-sediment deformation structures has received 39 40 41 43 significant credence in discriminating between glacial and non-glacial successions (e.g. 42 43 44 Lachniet et al., 2001; Menzies and Zaniewski, 2003; van der Meer and Menzies, 2011). 44 45 46 45 However, with a few notable exceptions (e.g. Benn and Prave, 2006; Arnaud, 2008 , 2012), 47 48 46 such analyses are scarcely applied to Neoproterozoic deposits. To redress this, we present a 49 50 47 new macro- and micro-scale structural analysis of the Chuos Formation of Cryogenian age in 51 52 53 48 the Otavi Mountainland, northern Namibia (Fig. 1). The study will utilise standard 54 55 49 sedimentological and structural analysis of the diamictite succession in order to determine the 56 57 58 59 60 2 61 62 63 64 65 50 genetic origin of the Chuos Formation, and assess the significance of soft-sediment 1 2 51 deformation structures as palaeoclimate proxies during Neoproterozoic glaciation. 3 4 5 6 52 1.1 Geological background 7 8 9 53 The Otavi Group is a carbonate-dominated succession of Neoproterozoic age housing two 10 11 54 diamictite horizons, the older Chuos and younger Ghaub formations (Fig. 1), each sharply 12 13 14 55 overlain by fine-grained carbonate deposits (Hoffmann & Prave, 1996; Hoffman & 15 16 56 Halverson, 2008; Miller, 2008). These horizons have been dated in turn to <746 ± 2 Ma 17 18 19 57 (Hoffman et al., 1996) and 635.5 ± 1.2 Ma (Hoffmann et al., 2004), through U-Pb zircon ages 20 21 58 of underlying and interbedded volcanic ash beds, leading to correlation with the purportedly 22 23 24 59 global Sturtian and Marinoan glaciations, respectively (Kennedy et al., 1998). In light of the 25 26 60 argued syn-rift derivation of the diamictite assemblages (e.g. Eyles & Januszczak, 2004, 27 28 61 2007), proponents of the glacial hypothesis have focussed largely on the younger Ghaub 29 30 31 62 Formation, considered to have accumulated during the ‗drift‘ stage of post-rift subsidence 32 33 63 (Hoffman & Halverson, 2008). The older Chuos Formation, by comparison, has received less 34 35 36 64 attention. 37 38 39 65 The glacial origin of the Chuos Formation was first proposed by Gevers (1931) due to its 40 41 42 66 lithological similarity to the Late Palaeozoic Dwyka Tillite (cf. Henry et al., 1986), and its 43 44 67 abundance of faceted and extrabasinal clasts. The stratigraphic position of the Chuos between 45 46 68 carbonate successions was used to support a glaciomarine origin (Martin 1965a, b; Hedberg 47 48 49 69 1979), in-keeping with the regional absence of subglacial striated pavements (Kroner & 50 51 70 Rankama, 1973). Alternative studies conversely describe the textural immaturity of the 52 53 54 71 diamictites, abundance of locally-derived erratic lithologies and their spatial and temporal 55 56 72 association with faults as evidence of high energy, rift-related submarine gravity flow 57 58 59 73 deposition (e.g. Hedberg, 1979; Miller 1983; Porada, 1983; Porada & Wittig 1983a, b; Martin 60 3 61 62 63 64 65 74 et al., 1985). Under this scenario, the abundant outsized clasts, frequently cited as evidence of 1 2 75 ice-rafting (e.g. Hoffman et al., 1998), are interpreted as the product of local mass flow 3 4 5 76 ‗rafting‘, or gravitational settling from overlying diamictites (Martin et al., 1985; Eyles & 6 7 77 Januszczak, 2007). Further models propose a compromise between these two hypotheses, 8 9 10 78 wherein both glacigenic deposition and gravitationally driven mass flows interact within an 11 12 79 ice marginal, glaciomarine environment (Hoffmann, 1983; Henry et al. 1986). 13 14 15 80 Three interpretations of the diamictites are thus possible: 1) those generated directly by 16 17 18 81 glacial processes; 2) those of primary glacial origin but re-worked by gravitational mass 19 20 82 transport; and, 3) those generated by mass flow or slope failure without glacial influence. 21 22 23 83 Consequently, criteria to correctly distinguish these environments remain pivotal to the 24 25 84 debate surrounding the origin of the Chuos Formation. These criteria, in turn, have wide 26 27 28 85 implications for interpretations of Neoproterozoic diamictites from a global perspective. 29 30 31 86 1.2. Study area and 32 33 34 87 The Otavi Mountainland in northern Namibia exposes a thick succession of Neoproterozoic 35 36 37 88 strata flanking the southern margin of the Owambo Basin (Fig. 1). The Chuos Formation 38 39 89 exhibits considerable lateral thickness variations across the region, reaching up to 130 m in 40 41 42 90 the central and western sectors, and pinching out towards the south-east (Hoffmann & Prave, 43 44 91 1996). In the study area, on Ghaub and Varianto farms (Fig. 2), it rests with angular 45 46 92 unconformity on and conglomerates of the Nosib Group (here the Nabis 47 48 49 93 Formation; Miller, 2008), and is sharply overlain by fine-grained carbonate mudstone of the 50 51 94 Berg Aukas Formation. The study area is ideally situated in the least deformed northern 52 53 54 95 margin of the Damara Orogenic Belt, characterised by a low shear, fold-thrust zone and sub- 55 56 96 greenschist facies metamorphism (Gray et al., 2008; Miller, 2008). These characteristics 57 58 59 60 4 61 62 63 64 65 97 permit detailed sedimentological and structural analysis of the diamictite horizons as they 1 2 98 have suffered minimal tectonic and metamorphic overprint. 3 4 5 6 99 2. Methodology 7 8 9 100 Sedimentary logging of the Chuos Formation was undertaken at several exposures across the 10 11 101 Ghaub and Varianto Farms at a metre-scale resolution. This process determined upper and 12 13 14 102 lower boundary relationships, enabling thickness changes to be documented across the area, 15 16 103 and the internal architecture of the formation to be characterised in detail. Lithofacies were 17 18 19 104 described on the macro-scale, including clast fabrics, bedding relationships, and the presence 20 21 105 and orientation of deformation structures. 22 23 24 106 For the purpose of this study, micro-scale analysis is restricted to an exposure of highly 25 26 27 107 deformed diamictite, where vertical changes in deformation style could be documented in 28 29 108 detail. Oriented samples were collected at appropriate intervals for thin section analysis, 30 31 32 109 determined where macro-scale changes in sedimentary or structural features were apparent. 33 34 110 Thin section analysis was undertaken using a petrographic microscope at low magnification 35 36 37 111 (1x and 2x), under plane and cross-polarized light, as well as examining high resolution 38 39 112 photographs and digital scans. Micro-scale features were described using standard structural 40 41 42 113 terminology and micromorphological techniques (sensu van der Meer, 1987, 1993; Menzies, 43 44 114 2000; Carr, 2004). 45 46 47 115 3. Sedimentology 48 49 50 51 116 3.1. Description 52 53 54 117 Three principal facies can be recognized comprising: 1) stratified diamictite facies, 2) sheared 55 56 118 and laminated diamictite, and 3) massive diamictite facies. In the south-east of the study area 57 58 59 119 the latter facies dominates (Fig. 2), with intercalated and overlying units of stratified and 60 5 61 62 63 64 65 120 sheared diamictite becoming more abundant northwards. At outcrop scale, diamictite units 1 2 121 typically coarsen upwards, although grading is rarely observed within individual beds. Clast 3 4 5 122 lithologies consist predominantly of well rounded quartzite, with minor sub-angular to 6 7 123 rounded mudstone, mica schist, gneiss and andesite. Striated and faceted clasts were 8 9 10 124 not observed. Evidence of impact-related deformation beneath some of the larger outsized 11 12 125 clasts is typically restricted to the stratified units, expressed through puncturing and 13 14 15 126 downwarping of the underlying laminae (Fig. 3a). This facies also preserves large-scale (1-2 16 17 127 m) eastward-dipping foreset structures (Log 2; Fig. 2), overlain by a series of soft-sediment 18 19 128 striated surfaces, comprising centimetre-scale linear grooves and ridges which trend 20 21 22 129 approximately north-south along exposed bedding planes (Fig. 3c). Detailed description of 23 24 130 deformation structures within the sheared and laminated facies will be discussed in section 4 25 26 27 131 below. 28 29 30 132 4. Deformation structures 31 32 33 133 A spectacularly well exposed vertical section of the Chuos Formation containing 34 35 36 134 approximately 30 m of highly deformed and attenuated diamictite crops out on Varianto 37 38 135 Farm (GR 19°24‘415 S, 17°42‘443 E; Fig. 2), in the central portion of the Otavi 39 40 41 136 Mountainland. The sequence overlies coarse of the Nabis Formation. These 42 43 137 deposits exhibit well developed convolute bedding and soft-sediment fold structures 44 45 46 138 approximately 6 m below the boundary with the Chuos, passing upwards into undeformed, 47 48 139 well bedded sandstone units (Log 3; Fig. 2). Likewise, finely laminated carbonate mudstone 49 50 140 of the Berg Aukas Formation, sharply overlying the Chuos Formation, is undeformed. 51 52 53 54 141 Lateral exposure of the Chuos Formation at this locality is limited, and thus descriptions on 55 56 142 the macro- and micro-scale below will focus on upsection evolution of the deformation 57 58 59 143 regime. For ease of description, the section has been divided into three structural zones, 60 6 61 62 63 64 65 144 shown on Fig 2. The lower and upper zones are dominated by ductile deformation structures 1 2 145 (e.g. rotational features, dispersion tails, clast boudinage). The intermediate zone, meanwhile, 3 4 5 146 is dominated by brittle deformation styles (e.g. fractured and crushed grains). 6 7 147 Micromorphological terminology follows the style of Brewer (1976), as adapted by Menzies 8 9 10 148 (2000), Zaniewski & van der Meer (2005) and Phillips et al. (2011) (see 8. Glossary). 11 12 13 149 4.1. Lower ductile zone (0-21 m) 14 15 16 150 4.1.1. Macroscale description 17 18 19 20 151 The basal 3 m of the Chuos Formation in Log 3 (Fig. 2) is characterised by massive to 21 22 152 crudely laminated, poorly sorted diamictite. Evidence of deformation in this horizon is 23 24 153 restricted to fractures traversing the larger outsized clasts (Fig. 3d). These sediments pass 25 26 27 154 gradually upwards through moderately to well laminated diamictites. Approximately 14 m 28 29 155 above the base of the section, these well laminated units exhibit asymmetric folds (Fig. 3e), 30 31 32 156 clast attenuation and development of asymmetric pressure shadows. These features 33 34 157 collectively record top-to-the-NW sense of simple shear. 35 36 37 38 158 At the south-eastern margin of the outcrop, these shear structures are overlain by a finely 39 40 159 laminated siltstone (Fig. 3f), which coarsens upwards with the input of granule to small 41 42 160 pebble sized clasts. The siltstone exhibits a concave-upward basal surface and planar upper 43 44 45 161 surface, and pinches out laterally (to the NW), where no disconformity in the diamictite is 46 47 162 visible. Above, the diamictite remains well laminated with deformation features including 48 49 50 163 rotational turbate structures (Fig. 3g), clast dispersion tails and pervasive lineations on the 51 52 164 clast surfaces, trending NNW (344°). 53 54 55 165 4.1.2. Microscale description 56 57 58 59 60 7 61 62 63 64 65 166 Micro-scale observations support a crudely developed lamination in the basal 3 m of the 1 2 167 logged section, passing gradually upwards into more delicately laminated intervals (Fig. 4A). 3 4 5 168 Rounded to sub-angular grains consist predominantly of quartzite, with minor feldspar, 6 7 169 siltstone and clay intraclasts. Clast long axis orientations are variable, although a high 8 9 10 170 proportion of grains exhibit a sub-horizontal microfabric (oriented N/NW to S/SE). 11 12 13 171 Planar features (e.g. linear grain alignments, symmetrical pressure shadows) are restricted to 14 15 172 the base of this zone, whereas rotational deformation structures (e.g. turbate structures, 16 17 18 173 dispersion tails, asymmetric pressure shadows and clast boudinage) are dominant throughout 19 20 174 (Fig. 4A), and become progressively more abundant upsection. This is reflected in the 21 22 23 175 development of necking structures between adjacent turbates (Fig. 5a). Bands of birefringent 24 25 176 clay material (plasmic fabric) also become more distinct and pervasive upsection, varying 26 27 28 177 from skel-masepic to skelsepic and unistrial. In places, this clay birefringence also outlines 29 30 178 distinct S-C fabrics (Fig. 5e). These features, alongside the rotational structures, support a 31 32 179 top-to-the N/NW shear sense. Towards the top of this zone, these structures and plasmic 33 34 35 180 fabrics are cross-cut by sub-horizontal, non-birefringent clays, which in places demonstrate 36 37 181 vertical, flame-like structures along their upper boundary. 38 39 40 41 182 4.2. Middle brittle zone 42 43 44 183 4.2.1. Macroscale description 45 46 47 184 At 21 m, poor exposure precludes detailed observation. At this interval, a metre thick unit of 48 49 50 185 massive diamictite crops out, without evidence of macro-scale deformation. 51 52 53 186 4.2.2. Microscale description 54 55 56 187 The thin section is composed of poorly sorted, angular to sub-rounded quartz and feldspar 57 58 59 188 grains which demonstrate no preferred orientation. Brittle deformation features, including 60 8 61 62 63 64 65 189 crushed quartz and fractured clasts, are common throughout, whereas ductile features were 1 2 190 not observed. The abundant fracture surfaces can be correlated between adjacent grains 3 4 5 191 across most of the section, resembling a ‗jig-saw fit‘ pattern (Figs. 5 & 7e). A distinct 6 7 192 birefringent fabric between the grains cannot be discerned. 8 9 10 11 193 4.3. Upper ductile zone 12 13 14 194 4.3.1. Macroscale description 15 16 17 195 Well laminated diamictite returns at 28 m, wherein a high proportion of clast long axes 18 19 20 196 parallel the sub-horizontal lamination, and are in places attenuated along this fabric (Fig. 3h). 21 22 197 Deformation is again dominated by rotational features including asymmetric boudins, S-C 23 24 198 fabrics and turbate structures. This interval is cross-cut by a carbonate sedimentary dyke, 25 26 27 199 consisting of brecciated fragments of the surrounding well laminated diamictite, set in a 28 29 200 carbonate mudstone matrix (Fig. 3i). 30 31 32 33 201 4.3.2. Microscale description 34 35 36 202 This zone is characterised by a much smaller clast population than the underlying lower 37 38 203 ductile and middle brittle zones, and becomes progressively more matrix-rich upsection 39 40 41 204 accompanied by more distinct horizontal lamination (Fig. 4C). As in the lower ductile zone, 42 43 205 rotational deformation features dominate, although conversely no planar features are present 44 45 46 206 (e.g. linear grain alignments, symmetrical pressure shadows). Turbate structures and 47 48 207 dispersion tails are abundant at the base of this zone, but diminish upsection where 49 50 51 208 asymmetric pressure shadows and clast boudins are more prevalent (Fig. 5f-h). These features 52 53 209 support a top-to-the NW sense of shear. Unistrial and skelsepic plasmic fabrics are highly 54 55 210 developed throughout, as well as sub-horizontal clay-rich layers which parallel and cut 56 57 58 59 60 9 61 62 63 64 65 211 oblique to the plasmic fabric (Fig. 5f-h). As in the lower ductile zone, these clay layers 1 2 212 exhibit flame structures along their upper boundary. 3 4 5 6 213 5. Discussion 7 8 9 214 5.1. Syn-sedimentary evolution of the Chuos Formation 10 11 12 215 Within the stratified diamictite facies, the widespread occurrence of outsized clasts with 13 14 15 216 impact-related deformation structures (Fig. 3a) is interpreted as evidence of ice-rafted debris 16 17 217 (IRD) (Thomas and Connell, 1985; Bennett et al., 1996; Condon et al., 2002), whereby 18 19 20 218 derivation via sediment gravity flow rafting is rejected owing to the absence of characteristic 21 22 219 clast imbrication, or correlation between bed thickness and maximum clast size (e.g. Martin 23 24 220 et al, 1985; Postma et al., 1988). Water depths were sufficient to accumulate IRD at repeated 25 26 27 221 stratigraphic intervals throughout deposition, but in the absence of diagnostic indicators, 28 29 222 accumulation within a glaciomarine or glaciolacutrine basin cannot be ascertained. Large- 30 31 32 223 scale cross-bed foresets within this facies (Log 2; Fig. 2) record tractive deposition and 33 34 224 development of a simple barform, prograding towards the east. In view of the evidence 35 36 37 225 favouring ice-rafting, these features are interpreted as ice-proximal subaqueous fan deposits 38 39 226 (e.g. Powell & Domack, 1995; Powell, 2003; Hornung et al., 2007). The overlying soft- 40 41 42 227 sediment striated surfaces (Fig. 3c), identical in morphology to those of the Hirnantian glacial 43 44 228 record of North (Sutcliffe et al., 2000; Le Heron et al., 2005), support the intrastratal 45 46 229 transmission of shear stresses and subglacial deformation following subaqueous fan 47 48 49 230 progradation. The absence of slickencrysts and polish on these surfaces discounts a later 50 51 231 tectonic origin (Petit & Laville, 1987; Eyles & Boyce, 1998). 52 53 54 55 232 In the massive diamictite facies, glacial indicators are typically absent. This is considered to 56 57 233 reflect glaciogenic debris flow remobilisation, consistent with the proposed ice-proximal 58 59 60 10 61 62 63 64 65 234 environment. In this setting, dynamic grounding-line oscillations would contribute to high 1 2 235 rates of sediment supply, supported by the presence of subaqueous fan deposits and common 3 4 5 236 coarsening upward profile of the diamictites (e.g. Benn, 1996; Evans et al., 2012), leading to 6 7 237 rapid accumulation and oversteepening of the sediment pile. Resultant re-working of the 8 9 10 238 glaciogenic sediments may also account, at least in part, for the unusual absence of striated 11 12 239 and faceted clasts within the Chuos Formation. 13 14 15 240 In light of the location of the study area at the northern margin of an intracratonic fold belt 16 17 18 241 (Miller, 2008), distinguishing the effects of soft-sediment or tectonogenetic deformation 19 20 242 within the sheared and laminated diamictite facies remains paramount. Overall, the lack of 21 22 23 243 bedding-discordant fabric development or metamorphic mineral overprint, and largely 24 25 244 undeformed nature of the underlying and overlying formations indicate a soft-sediment 26 27 28 245 genesis. In support, towards the top of Log 3 (Fig. 2) a carbonate dyke intrudes and brecciates 29 30 246 the sheared diamictite (Fig. 3i), with evidence of liquefaction of clay and silt grade material 31 32 247 along the intrusive contact. This is used to support a porewater-induced origin for the dyke, 33 34 35 248 representing hydrofracturing of the sediment pile (e.g. van der Meer et al., 2009), and thus 36 37 249 acts to support continued syn-sedimentary deformation after pervasive shearing and 38 39 40 250 attenuation of the diamictite. Moreover, kinematic indicators throughout the section, all 41 42 251 demonstrate top-to-the N/NW sense of shear (Figs. 3-5), consistent with the N-S strike of the 43 44 45 252 grooves on the soft-sediment striated surfaces. Conversely, the dominant structural grain 46 47 253 produced during Damaran orogenesis generated ENE-trending structures (Miller, 1983; Gray 48 49 et al., 2008), cutting oblique to the trend of the sedimentary structures within the Chuos 50 254 51 52 255 Formation. 53 54 55 256 56 57 58 59 257 60 11 61 62 63 64 65 258 5.2. Glacial vs. non-glacial deformation history 1 2 3 259 A striking feature of the described section, both on macro- and micro-scale, is the upsection 4 5 6 260 increase in deformation intensity, reflected in the increased abundance and lateral attenuation 7 8 261 of individual deformation structures. This incremental strain profile is a common feature 9 10 11 262 within subglacial regimes (Boulton and Hindmarsh, 1987; Hart and Boulton, 1991; Benn and 12 13 263 Evans, 1996; Evans et al., 2006; Hart, 2007), where the highest stress conditions are 14 15 264 encountered at the ice-bed interface, towards the top of the bed, and diminish downwards. In 16 17 18 265 contrast, deformation structures in a mass flow deposit exhibit the highest stress 19 20 266 characteristics at base (see Fig. 6), where friction between the flow and the underlying 21 22 23 267 substrate is greatest, thereby resulting in development of a basal shear zone accompanied by 24 25 268 an upward decrease in strain intensity (Nardin et al., 1979; Nemec, 1990; Hart & Roberts, 26 27 28 269 1994; Mulder and Alexander, 2001). 29 30 31 270 Features considered diagnostic of sediment remobilisation, including flow noses and tile 32 33 271 structures (Hart & Roberts, 1994; Bertran and Texier, 1999; Lachniet et al., 2001; Menzies & 34 35 36 272 Zaniewski, 2003), are conspicuously absent from the Chuos Formation. The former are 37 38 273 indicative of low shear, downslope slumping, and are hence rarely preserved under high 39 40 41 274 stress subglacial deformation (Lowe, 1982; Hart & Roberts, 1994), whilst the latter appear to 42 43 275 be a unique feature associated with deceleration and dewatering of sediment gravity flows 44 45 46 276 (Menzies & Zaniewski, 2003; van der Meer & Menzies, 2011). Similarly, clasts with coatings 47 48 277 of diamictite on the macro-scale, or concentrically laminated grain coatings on the micro- 49 50 278 scale, though not diagnostic, would support sediment re-working if present (Phillips, 2006; 51 52 53 279 Kilfeather et al., 2009). Furthermore, the absence of load structures and rare evidence of 54 55 280 vertical to sub-vertical water escape structures may be considered atypical of mass flow 56 57 58 59 60 12 61 62 63 64 65 281 deposition (Lowe, 1982; Visser et al., 1984; Hart & Roberts, 1994; Menzies & Zaniewski, 1 2 282 2003). 3 4 5 6 283 Rotational deformation structures are ubiquitous in the Chuos Formation, particularly on the 7 8 284 micro-scale (Figs. 4-5), and are encountered in both subglacial settings and sediment gravity 9 10 11 285 flows (Lachniet et al., 2001; Menzies and Zaniewski, 2003; Phillips, 2006). In the former, 12 13 286 these features are interpreted as the product of shearing within the deforming bed (van der 14 15 287 Meer, 1993, 1997), whereby stress is accommodated around a rotating nucleus (consisting of 16 17 18 288 a core stone or stiff matrix), leading to preferential alignment of smaller clasts at the nucleus 19 20 289 periphery. Similar mechanisms are envisaged in sediment gravity flows, although rotation 21 22 23 290 acts as a product of transient turbulent cells within the depositing flow (Phillips, 2006). 24 25 291 Although the latter mechanism cannot be excluded, the absence of other features indicative of 26 27 28 292 turbulence (e.g. normally graded beds), as compared to abundant evidence of pervasive 29 30 293 shearing, is compatible with a glacitectonic origin for the turbates. This is also supported by 31 32 294 their association with planar shear fabrics and structures throughout (cf. Hiemstra & Rijsdijk, 33 34 35 295 2003). 36 37 38 296 Additional deformation structures frequently cited as evidence of subglacial processes are 39 40 41 297 prevalent throughout (see Fig. 6), including pervasive tectonic lamination, unidirectional 42 43 298 folding, pressure shadows and clast dispersion tails (e.g. van der Meer, 1993; Hart & Roberts, 44 45 46 299 1994; Benn & Evans, 1996; Menzies et al., 1997; Lachniet et al. 2001; Carr et al., 2006; 47 48 300 Menzies et al., 2006). These features support ductile deformation under high cumulative 49 50 301 stress, facilitated by elevated porewater pressures, which succeed in lowering the effective 51 52 53 302 stress of the sediment (Menzies, 2000; Phillips et al., 2007; Lee & Phillips, 2008). These 54 55 303 conditions are commonly encountered in subglacial settings under the high overburden 56 57 58 304 pressure of ice and abundant basal meltwater supply (see section 5.3). This is also supported 59 60 13 61 62 63 64 65 305 by the presence of fractured and crushed grains, which frequently develop in zones of high 1 2 306 hydrostatic pressure near the ice-bed interface (Hiemstra and van der Meer, 1997; Menzies, 3 4 5 307 2000; Carr et al., 2006). 6 7 8 308 5.3. Hydrology of the subglacial bed 9 10 11 309 In view of the sedimentological evidence of primary deposition as a subaqueous diamictite, it 12 13 14 310 would be plausible to consider the sediments as water-saturated, and thus with elevated 15 16 311 porewater contents, prior to subglacial deformation. Nonetheless, the observed upsection 17 18 19 312 increase in strain intensity, alongside attenuation and lateral isolation of individual 20 21 313 microstructures, reflects sustained and increasing high porewater pressures throughout 22 23 24 314 deformation. 25 26 27 315 In a subglacial environment, the effect of overriding ice on porewater state will be threefold: 28 29 30 316 1) overburden pressure will increase confining pressure on the deforming bed, 2) the ice will 31 32 317 act as an impermeable seal inhibiting vertical water escape, and 3) friction at the ice-bed 33 34 318 interface will generate abundant basal meltwater, thereby increasing porewater content 35 36 37 319 (Evans and Hiemstra, 2005; Phillips et al., 2007; Lee & Phillips, 2008, 2011). A common 38 39 320 process in this scenario will be the development of lateral water escape features (Roberts and 40 41 42 321 Hart, 2005; Lee & Phillips, 2008), in this succession generating abundant sub-horizontal clay 43 44 322 –filled conduits. These features, in conjunction with well developed plasmic fabrics 45 46 47 323 throughout, support high concentrations of impermeable clay minerals within the sediment, 48 49 324 which act to further retard water escape from the deforming bed (Denis et al., 2009; 50 51 325 Lesemann et al. 2010). These factors will thus enable increased dilation of the sediment, 52 53 54 326 whereby the zone of subglacial shearing can extend deeper within the deforming sediment 55 56 327 pile (Lee & Phillips, 2008). 57 58 59 60 14 61 62 63 64 65 328 If porewater pressures continue to rise, stress at the ice-bed interface will reach critical levels 1 2 329 enabling the ice to decouple from its substrate (Evans et al., 2006), especially in contact with 3 4 5 330 underlying bed irregularities. This will typically result in infill of fine grained sediments 6 7 331 within the subglacial cavity, preserved as discontinuous lens-shaped beds (Evans & Benn, 8 9 10 332 2004; Lesemann et al., 2010). This interpretation is favoured for the finely laminated siltstone 11 12 333 recorded at the south-eastern margin of Log 3 (Fig. 3f). In contrast, when porewater pressures 13 14 15 334 are reduced, e.g. in response to enhanced water escape or freezing of the subglacial bed, 16 17 335 porewater influenced deformation will be inhibited, potentially leading to ‗locking-up‘ of the 18 19 336 deforming material (Evans et al., 2006; Lee & Phillips, 2008). As a result, the previously 20 21 22 337 water-saturated sediment may undergo brittle brecciation, as recorded in the middle brittle 23 24 338 zone (Fig. 4B). The jig-saw fit pattern of adjacent fractured surfaces (Fig. 5c) supports in situ 25 26 27 339 brecciation of this unit, consistent with rapid de-watering and deceleration of the mobile 28 29 340 deforming bed. This de-watering horizon in the middle of the sequence may therefore be used 30 31 32 341 to distinguish at least two phases of increasing porewater pressure during deformation of the 33 34 342 Chuos Formation, in-keeping with the characteristic polyphase nature of subglacial 35 36 343 deformation regimes (e.g. van der Meer, 1993; Menzies, 2000; Phillips et al., 2007, 2008). 37 38 39 40 344 5.4. Ice marginal model for the deposition and deformation of the Chuos Formation 41 42 43 345 Recent studies of stratified glacigenic diamictites within ice marginal environments have 44 45 46 346 advocated accumulation of thick, variably deformed sediment piles through the combined 47 48 347 effects of high sediment supply and glaciotectonic thrusting (e.g. Evans & Hiemstra, 2005; Ó 49 50 51 348 Cofaigh et al., 2011; Evans et al. 2012). In these settings, pre-existing stratification within the 52 53 349 sediment, commonly produced in response to the heterogeneous sediment inputs encountered 54 55 350 at the ice margin, encourages deformation partitioning along bed surfaces. Repeated 56 57 58 351 oscillations of the ice margin will therefore exploit these pre-existing structural weaknesses, 59 60 15 61 62 63 64 65 352 leading to glaciotectonic thrusting and stacking, and hence incremental thickening of the 1 2 353 succession. Depending on the extent of ice advance, these oscillations can also lead to 3 4 5 354 overriding of the sediment pile (Ó‘Cofaigh et al., 2011), resulting in subglacial as well as ice 6 7 355 marginal deformation. 8 9 10 11 356 A similar setting is envisaged to account for the sedimentological and structural features 12 13 357 preserved within the Chuos Formation. The development of sub-horizontal lamination at the 14 15 358 relatively less deformed base of Log 3 (Fig. 2), alongside the widespread occurrence of 16 17 18 359 stratified diamictites throughout the Chuos Formation, support the generation of a syn- 19 20 360 depositional, or ‗pre-existing‘ stratification. Subsequent deformation of the sediment pile 21 22 23 361 clearly resulted in deformation partitioning along bed/lamina contacts since tectonic 24 25 362 lamination, shear structures and plasmic fabrics are bed-parallel throughout. However, in 26 27 28 363 contrast to previous studies (e.g. Evans & Hiemstra, 2005; O‘Cofaigh et al. 2011; Evans et al. 29 30 364 2012) a significant thrust component was not recorded in the Chuos Formation. It is possible 31 32 365 that sub-horizontal shear surfaces within the sediment also operated as thrust planes during 33 34 35 366 proglacial to submarginal deformation, leading to progressive stacking and thickening of the 36 37 367 sediment pile. Alternatively, in light of the evidence for high porewater contents and 38 39 40 368 significant ice overburden pressure, vertical stacking of the deforming bed may not have been 41 42 369 possible, resulting in lateral attenuation and ‗smearing‘ of the sediment pile as opposed to 43 44 45 370 thrust-related aggradation. This may also reflect the subglacial position of the sediment 46 47 371 throughout deformation, where it would have been sheltered from proglacial and ice marginal 48 49 tectonics. In contrast, discrete bed-scale horizons of subglacial deformation associated with 50 372 51 52 373 the grounding-line fan deposits (Log 2; Fig. 2) likely reflect periodic overriding in ice 53 54 374 marginal positions, with potentially thrust-related shearing accommodated in the soft- 55 56 57 375 sediment striated surfaces (Fig. 3c). 58 59 60 16 61 62 63 64 65 376 6. Conclusions 1 2 3 377 The Chuos Formation in the Otavi Mountainland, northern Nambia, accumulated in an ice- 4 5 6 378 proximal subaqueous environment prior to secondary subglacial deformation. Detailed 7 8 379 analysis of soft sediment deformation structures was critical in determining the presence and 9 10 11 380 influence of grounded ice at this time. These features, in conjunction with the deposition of 12 13 381 ice-proximal subaqueous fan deposits, and abundant ice-rafted debris at recurrent 14 15 382 stratigraphic intervals throughout the Chuos diamictites act to support dynamic oscillations of 16 17 18 383 the ice grounding-line. Therefore, the unusual paucity of ‗classic‘ glacial indicators (i.e. 19 20 384 striated and facteted clasts, striated pavements) does not preclude Neoproterozoic glaciation, 21 22 23 385 as frequently argued under sediment gravity follow hypotheses. Soft sediment deformation 24 25 386 structures are thus considered as key, and largely under-considered, palaeoclimate proxies, 26 27 28 387 with significant implications for determining the glacigenic origin of pan-global diamictite 29 30 388 successions, as well as the nature of subglacial bed conditions during the Neoproterozoic 31 32 389 icehouse. 33 34 35 36 390 7. Acknowledgments 37 38 39 391 The authors wish to thank Paulus N. Mungandjera and Ralph J.C.M. Muyamba, UNAM, for 40 41 42 392 their invaluable field assistance, and the owners of Ghaub and Varianto Farms for granting 43 44 393 permission to examine these sections on their land. Thanks also to Kevin D‘Souza, RHUL, 45 46 394 for assistance with thin section photography. 47 48 49 50 395 8. Glossary 51 52 53 396 Based on Brewer (1976), as adapted by Menzies (2000), and Zaniewski and van der Meer 54 55 397 (2005,) unless otherwise stated. 56 57 58 59 398 Dispersion tail: concentration of smaller grains or plasma in the lee of a larger grain. 60 17 61 62 63 64 65 399 Necking structure: variety of turbate structure. Alignment of smaller grains occurs between 1 2 400 adjacent larger grains (Lachniet et al. 1999). 3 4 5 6 401 Plasma: particles of colloidal size (<20 μm), including mineral and organic material, within 7 8 402 which individual grains cannot be discerned. 9 10 11 403 Plasmic fabric: orientation of plasma domains based on the optical properties (birefringence) 12 13 14 404 of aligned plasma particles. Common varieties include: 15 16 17 405 Asepic: anisotropic plasma domains with little to no preferred orientation. Sub- 18 19 20 406 varieties include Argillasepic (dominantly clay-sized particles) and Silasepic 21 22 407 (dominantly silt-sized particles). 23 24 25 408 Masepic: short plasmic fabric domains with a single preferred orientation. 26 27 28 29 409 Bimasepic: plasma particles exhibit two dominant preferred orientations. Termed 30 31 410 Lattisepic where these directions are perpendicular. 32 33 34 411 Insepic: small clusters of oriented plasma particles where clusters show no preferred 35 36 37 412 orientation. 38 39 40 413 Multisepic: multiple (>2) preferred plasmic fabric orientations. 41 42 43 44 414 Skelsepic: plasmic particles preferentially oriented around skeleton grains. 45 46 47 415 Omnisepic: random orientation of various plasmic fabric domains. 48 49 50 416 Unistrial: elongate, discrete bands of birefringent clay plasma. 51 52 53 417 Pressure shadow: Typically massive domain of lower strain adjacent to a clast. Synonymous 54 55 56 418 with strain shadow (Phillips et al. 2011). 57 58 59 60 18 61 62 63 64 65 419 Turbate structure: circular arrangement of grains around a core stone or stiff matrix. Long 1 2 420 axes of oriented grains exhibit a parallel or radial orientation relative to the margins of the 3 4 5 421 core stone (Hiemstra & Rijsdijk, 2003). 6 7 8 422 9. References 9 10 11 12 423 Allen, P.A., Etienne, J.L., 2008. Sedimentary challenge to Snowball Earth. Nature 13 424 Geoscience, 1, 817-825. 14 15 425 Arnaud, E., 2008. Deformation in the Neoproterozoic Smalfjord Formation, northern 16 17 426 Norway: an indicator of glacial depositional conditions? Sedimentology, 55, 335-356. 18 19 427 Arnaud, E., 2012. The paleoclimatic significance of deformation structures in Neoproterozoic 20 428 successions. Sedimentary Geology, 243–244, 33–56. 21 22 23 429 Benn, D.I., 1996. 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Geological Survey of Finland, Special Paper, 3, p. 77-89. 16 17 641 van der Meer, J.J.M. 1993. Microscopic evidence of subglacial deformation. Quaternary 18 642 Science Reviews, 12, 553-587. 19 20 21 643 van der Meer, J.J.M. 1997. Subglacial processes revealed by the microscope: particle and 22 644 aggregate mobility in till. Quaternary Science Reviews, 16, 827-831. 23 24 645 van der Meer, J.J.M., Kjær, K.H., Krüger, J., Rabassa, J., Kilfeather, A.A., 2009. Under 25 26 646 pressure: clastic dykes in glacial settings. Quaternary Science Reviews, 28, 708-720. 27 28 647 van der Meer, J.J.M., Menzies, J., 2011. The micromorphology of unconsolidated sediments. 29 648 Sedimentary Geology, 238, 213-232. 30 31 32 649 Visser, J.N.J., Colliston, W.P., Terblanche, J.C. 1984. The origin of soft sediment 33 650 deformation structures in Permo-Carboniferous glacial and proglacial beds, South 34 651 Africa. Journal of Sedimentary Petrology, 54, 1183-1196. 35 36 37 652 Zaniewski, K., van der Meer, J.J.M. 2005. Quantification of plasmic fabric through image 38 653 analysis. Catena, 63, 109-127. 39 654 40 41 655 42 43 44 656 45 46 657 47 48 658 49 50 51 659 52 53 660 54 55 661 56 57 58 662 59 60 25 61 62 63 64 65 663 Figure captions 1 2 3 664 Figure 1: A. Map outlines stratigraphic framework of Namibia, and the location of the study 4 665 sites in the Otavi Mountainland, along the south-eastern flank of the Owambo Basin, after 5 6 666 Miller (2008). B. Stratigraphy of the Cryogenian succession exposed in the study area, after 7 8 667 Hoffman & Prave (1996). 9 10 11 668 Figure 2: Logged sections of the Chuos Formation, located on inset geological map of Ghaub 12 13 669 and Varianto Farm study areas. Note the overall lateral transition from massive to stratified 14 670 diamictites towards the west/north-west. The stratigraphic location of images shown in 15 16 671 figures 3-5 are indicated by their corresponding numbers adjacent to the logs. Map modified 17 18 672 after Geological Survey of Namibia (2008). 19 20 21 673 Figure 3: Macro-scale sedimentary and deformation structures. Coin and lens cap for scale 22 23 674 measure 2 cm and 5 cm, respectively. A) Outsized clast with impact-related deformation 24 675 structure, indicated by white arrows. B) Fractured clast infilled by diamictite. C) North-south 25 26 676 trending linear grooves and ridges interpreted as soft-sediment striated surfaces. D) Fractured 27 28 677 outsized clast within the comparatively undeformed base of Log 3 (Fig. 2). E) Asymmetric 29 30 678 fold demonstrating top-to-the-NW vergence. Note clast attenuation parallel to deformed 31 32 679 laminae. F) Laminated siltstone lens, interpreted as ice-bed separation feature, restricted to 33 34 680 the south-eastern margin of the exposed section (Log 3; Fig. 2). G) Rotational turbate 35 681 structure showing preferential alignment of smaller clasts around the margins of the central 36 37 682 obstacle clast (picked out by white dashed lines). Note top-to-the-SE shearing and abrasion of 38 39 683 core stone. H) Pervasive tectonic lamination. Asymmetric clast boudinage and S-C structures 40 41 684 define a top-to-the-NW shear sense. I) Carbonate dyke cross-cuts sheared diamictite (G). 42 43 685 Note brecciated fragments of deformed diamictite within the dyke. 44 45 686 Figure 4: Paired thin section photograph and interpretation of Chuos Formation 46 47 687 microstructures. Lower Ductile Zone: A) Well developed sub-horizontal fabric and abundant 48 49 688 rotational deformation structures (turbate structures, asymmetric boudins, clast dispersion 50 51 689 tails and asymmetric pressure shadows).Microstructures support top-to-the-NW shear sense. 52 53 690 Middle Brittle Zone: B) This sample is characterised by fractured and crushed quartz grains 54 55 691 (white) and magnetite crystals (black). Fractured surfaces can be traced between adjacent 56 692 clasts, akin to a ‗jig-saw‘ pattern, indicating in situ clast breakage. Sense of deformation 57 58 693 cannot be ascertained. Upper Ductile Zone: C) Microstructures are highly attenuated, defined 59 60 26 61 62 63 64 65 694 by the predominance of sheared clast boudins. These features, alongside asymmetrical 1 2 695 pressure shadows and rare rotational turbates support top-to-the-NW deformation. Note sub- 3 4 696 horizontal clay layers cutting oblique to tectonic lamination, with flame-like structures on 5 697 their upper boundary. 6 7 8 698 Figure 5: Micro-scale deformation structures. Photos A-D captured under plane-polarized 9 10 699 light, E-H under cross-polarized light. White bar for scale measures 1 mm. Lower Ductile 11 12 700 Zone: A) Necking structure developed between adjacent turbate structures. B) Central 13 14 701 rotational deformation structure with associated ‗tails‘ of silt-grade sediment and small sand 15 702 clasts. Top-to-the-NW shear sense also supported by asymmetric pressure shadow in the 16 17 703 uppermost portion of the image. Middle Brittle Zone: C) Dashed lines highlight ‗jig-saw fit‘ 18 19 704 fractured grain surfaces. Upper Ductile Zone: D) Rotational structure with fabric-parallel 20 21 705 ‗tails‘ of clay-grade sediment and small sand grains. Abrasion and fracturing of core stone 22 23 706 interpreted as product of pervasive shearing. E) Birefringent clay particles outline S-C 24 25 707 fabrics, supporting top-to-the-N shear sense. F) Asymmetric pressure shadow outlined by 26 708 skelsepic-unistrial plasmic fabric. Note discontinuous clay layers trending both parallel and 27 28 709 oblique to the plasmic fabric (black arrow). G-H) Well developed unistrial to skelsepic 29 30 710 plasmic fabric surrounding boudinaged skeleton grains. Note discontinuous clay layers 31 32 711 cutting oblique to the plasmic fabric (black arrow). 33 34 35 712 Figure 6: Schematic diagram highlights dominant style and association of deformation 36 713 structures typically encountered within a sediment gravity flow as compared to the 37 38 714 assemblage identified in the Chuos Formation. Length of black arrow corresponds to strain 39 40 715 intensity, wherein the former would be expected to demonstrate a basal shear zone with 41 42 716 upward decreasing strain intensity, vertical water-escape structures and abundant evidence of 43 44 717 re-working (e.g. diamictite intraclasts, laminated clast coatings, slump folds). In contrast, the 45 46 718 observed upsection increase in strain intensity, abundance of ductile deformation features, 47 719 pervasive tectonic lamination, and sub-horizontal water-escape features act to support 48 49 720 subglacial deformation under high cumulative stress and elevated porewater pressures. 50 51 721 Adapted after Evans et al. (2012) and Phillips (2006). 52 53 54 722 Figure 7: Ice marginal model for the deposition and subsequent deformation of the Chuos 55 56 723 Formation. Ice-proximal subaqueous deposition occurs through the combined input of ice- 57 724 rafted, englacial, supraglacial and subglacial debris, building out as ice-contact subaqueous 58 59 60 27 61 62 63 64 65 725 fans. During grounding-line oscillations, these sediments undergo subglacial deformation 1 2 726 processes behind the ice front, periodic mass wasting in advance of the ice front, as well as 3 4 727 proglacial to ice marginal intrastratal shear. Modified after Ó‘Cofaigh et al. (2011) and Evans 5 728 et al. (2012). 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22 23 24 25 26 27 28 29 30 31 32 33 34 35 36 37 38 39 40 41 42 43 44 45 46 47 48 49 50 51 52 53 54 55 56 57 58 59 60 28 61 62 63 64 65 Figure Click here to download high resolution image Figure Click here to download high resolution image Figure Click here to download high resolution image Figure Click here to download high resolution image Figure Click here to download high resolution image Figure Click here to download high resolution image Figure Click here to download high resolution image