On the origin of the Itararé Group basal nonconformity and its implications for the Late Paleozoic glaciation in the Paraná Basin, Brazil

Accepted Manuscript On the origin of the Itararé Group basal nonconformity and its implications for the Late Paleozoic glaciation in the Paraná Basin, Brazil

Claus Fallgatter, Paulo S.G. Paim PII: DOI: Reference:

S0031-0182(16)30424-2 doi: 10.1016/j.palaeo.2017.02.039 PALAEO 8225

To appear in:

Palaeogeography, Palaeoclimatology, Palaeoecology

Received date: Accepted date:

2 September 2016 8 February 2017

Please cite this article as: Claus Fallgatter, Paulo S.G. Paim , On the origin of the Itararé Group basal nonconformity and its implications for the Late Paleozoic glaciation in the Paraná Basin, Brazil. The address for the corresponding author was captured as affiliation for all authors. Please check if appropriate. Palaeo(2016), doi: 10.1016/ j.palaeo.2017.02.039

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ACCEPTED MANUSCRIPT C. Fallgatter and P.S.G. Paim

On the origin of the Itararé Group basal nonconformity and its implications for the Late Paleozoic glaciation in the Paraná Basin, Brazil CLAUS FALLGATTERa*, PAULO S. G. PAIMa a

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PPGGeo, Universidade do Vale do Rio dos Sinos - UNISINOS -, Av. Unisinos, 950, Setor C02, São Leopoldo-RS, Brazil (E-mail: [email protected]) * Current address: School of Geosciences, Department of Geology and Petroleum Geology; University of Aberdeen, Aberdeen AB24 3UE, UK

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ABSTRACT

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Outcrop studies of the Itararé Group in the Alfredo Wagner area, along the eastern margin of the Paraná Basin in Brazil, reveal a partially exhumed, NNW-SSE oriented glacial paleovalley. The Carboniferous Rio do Sul Formation represents the basal interval of the glacially influenced Itararé Group, and rests directly on Precambrian basement rocks. This nonconformity is thought to represent an areal scouring surface, as indicated by the series of glacially-carved linear depressions that were incised into the glaciated bedrock during glacial advances. These large-scale erosional landforms differ in their extent by orders of magnitude, and include wide, deep troughs as well as narrower, shallower gouges. Glacial troughs and gouges usually display smaller-scale, subglacial erosion landforms, such as polished and grooved striations on their walls and floors. These indicate the NNW motion of a warm-based, grounded glacier. In addition, large-scale erosional features represent local topographic lows within the major glacial paleovalley. These local lows are large enough to contain part of the succeeding deglacial deposits, which thin toward and onlap the wall of these depressions. We grouped these deglacial deposits into six genetic facies associations (FA-1 to -6) within three fining-upward glacial cycles. FA-1 is composed of black shale, and represents an initial marine flooding event. It can be correlated with the regionally widespread Lontras Shale. FA-2 is composed of sandstones and mudstones associated with floodderived hyperpycnal flows, and occasionally with surge-type turbidites. FA-3 is represented by plastered or injected, matrix-supported conglomerates emplaced as lodgment tills. FA-4 comprises matrix- and clast-supported conglomerates, sandstones, and mudstones, attributed to ice-rafting and concentrated flows in grounding-line fan facies. This unit also contains ice-keel dragging features, carved into soft substrate. FA5 is composed of dropstone-bearing black shale, related to a second major marine flooding event within the paleovalley. FA-6 comprises sandstones associated with channelized, surge-type turbidites, forming the second turbidite system in the succession. As these deposits are confined to a paleovalley, we infer that sedimentation was influenced by topographically controlled ice streams, instead of non-confined, widespread ice lobes as previously suggested. Reconstructing South America and southern Africa in their positions during the LPIA suggests that the glacial valleys then present in the Namibia area may have extended further, into the eastern margin of the Paraná Basin. In this scenario, confined outlet glaciers would have flowed out from the Namibian Windhoek ice cap, and terminated along the eastern margin of the Paraná Basin. Keywords: Subglacial erosion; ice stream; paleovalley; topographic control; glacial paleogeography. 1

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1. Introduction

During the Carboniferous and Permian, small continental ice caps and alpine glaciers influenced deposition in many sedimentary basins throughout the Gondwanan continents (Isbell et al. 2003, 2012; Fielding et al., 2008a; Limarino et al., 2006, 2014a). The supercontinent Gondwanaland was affected by multiple episodes of glaciation, known collectively as the Late

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Paleozoic Ice Age (LPIA; Fielding et al., 2008; Isbell et al., 2012). The LPIA represents the longest glacial interval of the Phanerozoic (Isbell et al., 2003; Fielding et al., 2008a; Rocha-

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Campos et al., 2008), but not all of the glacial centers were active at the same time (Visser, 1997). The subglacial and proglacial sediments deposited over much of Gondwanaland as a

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result of the LPIA have been a major focus of research.

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The glacially influenced deposits of the Paraná Basin, which are included in the Itararé Group (Schneider et al. 1974; França and Potter, 1988), have been studied since the first half of

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the twentieth century. However, the role played by topography in determining patterns of deposition has not been fully explored in the literature, possibly due to the lack of suitable exposures. Topographic control of deglacial deposition in the Paraná Basin has been proposed

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(Gravenor and Rocha-Campos, 1983; Santos et al., 1996) but can only rarely be demonstrated

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(França et al., 1996). In addition, the timing and style of Late Paleozoic glaciation, in the Paraná Basin and elsewhere, remain poorly understood (Isbell et al. 2003; Fielding et al., 2008a).

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Depositional models and paleogeographic reconstructions of the Itararé Group have always interpreted late Paleozoic glacial ice to have been unconstrained by topography (França

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and Potter, 1988; Santos et al., 1996; Rocha-Campos et al., 2008; Vesely et al., 2015). Reconstructions of the glacial paleogeography are mostly based on subglacial erosion features and the succeeding deposits, which show ice movement mainly to the N and NW (Rosa et al., 2016 and references therein). Based on these reconstructions, late Paleozoic glacial deposits in the eastern Paraná Basin have been attributed to the advance and retreat of an ice sheet (Bigarella et al. 1967; Crowell and Frakes, 1970), or more recently to the lobate termini of ice streams (Santos et al. 1996), or multiple, small ice lobes (Vesely et al. 2015).

A source of ice in South Africa may represent the spreading center from which glaciers reached the eastern Paraná Basin. During the LPIA, the Neoproterozoic Dom Feliciano Orogenic Belt, which composes most of the Santa Catarina Shield in the study area, was connected to the 2

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Gariep Orogenic Belt of Namibia and South Africa (Basei et al., 2005), supporting previous models in which ice flows from South Africa, over Precambrian terranes, to reach the Paraná Basin (Santos et al., 1996; Rocha-Campos et al., 2008; Vesely et al., 2015). As glaciers moved northwestward from Africa, they would have overrode the Santa Catarina Shield, and migrated into the Paraná Basin (Gravenor and Rocha-Campos, 1983). Therefore, it is reasonable to propose that the Itararé Group is confined to glacial troughs or valleys in parts of the eastern Paraná Basin, as has been postulated for its southern margin (Tedesco et al. 2016; Rio Grande do

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Sul Shield), and elsewhere in the Gondwana during the LPIA (Kneller et al. 2004; Dykstra et al. 2006; Henry et al. 2010; Aquino et al. 2014; Limarino et al. 2014b; Argentina; Dickins, 1996;

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Jones and Fielding, 2004; Fielding et al. 2008b; Australia; Visser, 1981; Visser and Kingsley, 1982; Visser, 1987, 1990; Smith et al. 1993; von Brunn, 1996; Isbell et al. 2008; South Africa

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and Bussert, 2010; Ethiopia).

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The Itararé Group has long been recognized as evidence for late Paleozoic glaciation in the Paraná Basin. Studies have focused on the deposition of deglacial (and rarely subglacial)

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strata. However, questions concerning the absolute age of these deposits, and glacial paleogeography at the time of deposition, merit further, more detailed work. No evidence for topographic control on sedimentation has been documented from the Precambrian basement

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rocks cropping out along the eastern margin of the Paraná Basin. If such glacially-carved

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features could be identified, they might provide a clue as to the behavior, extent, and timing of glaciation across Gondwana. Therefore, the main objectives of this study include: i) description of the Itararé Group basal nonconformity, and interpretation of its origin and morphology; ii)

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description of the facies architecture, and construction of a high-resolution stratigraphic framework for the sedimentary fill of the larger subglacial erosion features; iii) reconstruction of

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the local subglacial and proglacial history; and iv) assessment of the paleogeographic implications of our results.

2. Geologic Setting

The Paraná Basin is a Paleozoic-to-Mesozoic, northeast-southwest oriented, elongate intracratonic basin developed on the South American platform. It contains nearly 7,000 meters of sedimentary and volcanic strata (Zalán et al. 1990; Milani et al., 2007). The studied succession is part of the Itararé Group, which represents the base of a complete transgression-regression cycle named the Gondwana I Supersequence (Fig. 1; Milani et al., 1998). The Itararé Group is late Bashkirian to late Moscovian in age (Souza, 2006; Cagliari et al., 2016) and records a prolonged 3

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interval of continental ice sheet influence in the Paraná Basin during the LPIA. It includes a number of subglacial features, dropstone-bearing intervals, and mass and gravity flow deposits (Castro, 1980; Gama Jr. et al., 1992; Eyles et al., 1993; Vesely and Assine, 2002; Vesely, 2007; d’Ávila, 2009; Suss et al., 2014; Aquino et al., 2016; Carvalho and Vesely, 2016). The orientation of subglacial features indicates that ice movement was mainly to the north and northwest (Rosa et al., 2016). The Itararé Group is more than 1,300 meters thick, and crops out

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along the basin margins.

Based on outcrop sections along the eastern margin of the basin, Schneider et al. (1974)

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divided the Itararé Group into the Campo do Tenente, Mafra and Rio do Sul formations. The studied interval represents the lower portion of the Rio do Sul Formation (Fig. 1). França and

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Potter (1988), working mainly with subsurface data, divided the Itararé Group into three major fining-upward glacial-deglacial cycles, with different names (Fig. 1). Canuto et al., (2001) also

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reported higher frequency glacial-deglacial cycles, represented by seven fining- or coarseningupward successions exposed in the north of Santa Catarina State. Later, Vesely and Assine

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(2006) documented five fining-upward sequences in the Paraná State, associated with the retrogradation of proglacial, glaciomarine depositional systems. In both cases, the high

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frequency cycles are mostly composed of diamictites and dropstone-bearing, fine-grained strata.

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The Santa Catarina Shield is represented in the study area by the Precambrian igneous rocks of the Mantiqueira Province (Heilbron et al., 2004). This province includes a Neoproterozoic mobile belt, subdivided into the diachronous Araçuaí, Ribeira, Dom Feliciano

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and São Gabriel orogens (Almeida et al., 1981; Heilbron et al., 2004). In the Santa Catarina Shield, the Dom Feliciano Orogenic Belt can be further subdivided into the Florianópolis

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Terrane, the Brusque Metamorphic Complex, and the Itajaí foreland basin. In the Alfredo Wagner locality, the Itararé Group (the basal of Rio do Sul Formation) sits directly on the Florianópolis Terrane (Fig. 2), which is locally composed of granites and granitoids of the Águas Mornas Complex and Paulo Lopes Intrusive Suite. The Alfredo Wagner area contains wellknown exposures of the Itararé Group along the eastern margin of the Paraná Basin, and has been the subject of investigation since the nineteen-eighties (Castro, 1980; Rocha-Campos et al., 1988; Canuto, 1993). The outcrops described here are located along the Picadas, Lessa and Adaga rivers, and along two main roads (BR 282 and SC 22) that run parallel to these rivers (Fig. 2). The study area also includes younger Paraná Basin strata, which crop out along the escarpments of the Serra do Faxinal (Fig. 2).

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3. Methods We implemented an outcrop-based research strategy in the eastern Paraná Basin, focusing on the Alfredo Wagner locality in southeastern Santa Catarina State, Brazil. We identified and described several sedimentary facies, then grouped them into six facies

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associations representing six successive deposits. Using this framework, we assembled an 8 km long, strike-oriented panel correlation showing the vertical stacking of facies as well as the

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morphology of the basal nonconformity (Fig. 3). The morphology of the basal nonconformity

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was determined by a combination of extensive field mapping, sedimentary unit thickness changes, onlap relationships, and paleocurrent trends, and revealed local topographic highs and

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lows within a major trunk valley. Sedimentary units were logged at several sites and their relative timing were established based on their vertical succession and distribution above the

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Precambrian basement.

Eight 1:100 scale sedimentary logs, containing ~154 m of total strata, were generated in

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the lower part of the Rio do Sul Formation. Each sedimentological log is paired with a gamma ray log, with readings (counts per second) taken at intervals of 30 cm. The highest gamma-ray

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readings coincide with the highest thorium and uranium content, seen in a thick black shale interval (FA-5, see below) interpreted as a maximum flooding surface, and used as a datum for correlation. Sedimentological and gamma ray logs, combined with outcrop descriptions and

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ground photomosaics, were used to interpret depositional settings. Sediment dispersal patterns

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were inferred from more than 50 paleocurrent readings.

4. Stratigraphy

The studied interval was divided into six stacked facies associations (FA-1 to -6), which record an upward fluctuation of the glacier margin (see Table 1). Some of these facies associations were further subdivided into ‘a’, ‘b’, and ‘c’ categories based on their genetic and stratigraphic relationships. A panel correlation showing the distribution of these facies associations in the study area is shown in Fig. 3. FA-1 through 4 are confined to local topographic lows, while FA-5 and 6 are more widespread, however, all of these deposits probably represent the fill of a much larger glacial erosional feature (Fig. 3), based on their onlap 5

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relationships and the confined nature of coeval strata in nearby areas (Puigdomenech et al., 2014; Fallgatter, 2015).

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4.1. Basal nonconformity

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4.1.1. Areal scouring and subglacial erosion

The basal nonconformity of the Itararé Group in the Alfredo Wagner area is represented

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by a surface of areal scouring, up to 8 km wide (Fig. 3), which include NNW-oriented, large linear depressions and smaller streamlined landforms. Linear depressions of different orders of

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magnitude occur throughout this surface, forming topographic highs and lows in the granite bedrock (Fig. 3). They can be informally divided into a deeper and wider class (with long

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wavelength and amplitude) and a shallower and narrower class (with short wavelength and amplitude). Deep, wide depressions are the most prominent linear erosion landforms, and are estimated to be up to 560 m wide and 40 m deep in the study area; similar erosional features to

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the southwest of the study area suggest that these depressions can be even wider and deeper. In

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contrast, the shallower and narrower scours are up to 15 m wide and 8 m deep (Fig. 3); RochaCampos et al. (1988) described them as forming an undulating, irregular surface. They display smooth, polished striations and grooves, on walls that range from gentle to steep. Both classes of

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(Fig. 4A and B).

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erosion features are large enough to have trapped deglacial sediments, which onlap their walls

The bedrock surface is characterized by grooved striations, plastic molded forms (‘pforms’; Bennett and Glasser, 2009), and polished surfaces (Fig. 3), most of them previously undescribed. These structures are observed only in the walls and floors of smaller carved depressions. Grooved, striated pavements display lineations up to 2 mm deep and 1 cm wide (Figs. 4C and D). Striations in the Rio Picadas III outcrop (Figs. 2 and 4C) were described by Rocha-Campos et al. (1988), who also identified grooves and friction cracks in lunate and crescentic forms. In at least one locality (the Rio Lessa outcrop, see Figs. 2 and 3), striations are superimposed on larger, grooved lineations (p-forms) associated with a polished wall (Fig. 4E). Polished surfaces are widespread on the basal nonconformity of the Itararé Group, and are

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usually associated with grooves and striations along linear depressions, directly overlain by deglacial deposits. The areal scouring surface and subglacial erosion landforms observed in the Alfredo Wagner area indicate that ice slid over the Precambrian bedrock. In general, this nonconformity is thought to represent the sides and floor of a broad, NNW-oriented glacial trunk valley. A basement high to the ENE is believed to represent the eastern side of the paleovalley.

covered by younger strata, and does not crop out at the surface.

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Unfortunately, due to the regional dip of the Parana Basin to the west, its western margin is

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These subglacial erosion landforms are related to glacial advances that molded the

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basement, resulting in a series of linear, glacially-carved lows preserved as deep, wide troughs or shallow, narrow gouges. Their differences in magnitudes may be the result of differential

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abrasion, as rates of abrasion are controlled mainly by changes in pressure, the velocity of sliding ice, and the concentration of debris at the glacier sole (Boulton, 1979; Hallet, 1981; Iverson, 2002) or subglacial hydrology (Herman et al., 2011). Comparably large subglacial erosion

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landforms have been reported from the late Paleozoic of Western Australia (Eyles and Broekert, 2001), and from the late Pleistocene of Europe and North America (Eyles and McCabe, 1989;

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Wingfield, 1990; Stewart and Lonergan, 2011). They have been attributed to glacial tunnel valleys (Benn and Evans, 2010), which require subglacial meltwater and glacial scouring, both

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of which are more common in wet-based glaciers.

The streamlined character of the glaciated granite surface also suggests that it resulted

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from abrasion, quarrying and subglacial meltwater erosion combined with the bedrock properties (Krabbendam and Glasser, 2011). The polished and striated walls of glacial troughs and gouges

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indicate that these were made by individual clasts travelling at the base of the glacier and that linear depressions were at least partially carved by ice. The origin of friction cracks has been related to the removal of rock slabs by subglacial quarrying (Boulton, 1979; Iverson, 2002; Glasser and Bennett, 2004). Grooved striations superimposed on p-forms suggest a combination of intense meltwater flow and secondary glacial abrasion. Mechanical abrasion implies that the glacier was carrying a basal debris load, and eroded the basement primarily by linear scratching (Glasser and Bennett, 2004), however, grooved-striations such as p-forms can form by either glacial abrasion or meltwater erosion. The polished walls and floors observed in some glacial troughs and gouges are the result of mechanical shearing of fine-grained sediments or rock flour over bedrock (Boulton, 1979; Bennett and Glasser, 2009; Benn and Evans, 2010). 7

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These features suggest a grounded, warm-based glacier that moved via basal sliding, as has been suggested for other sites in the Paraná Basin during the same interval (Gravenor and Rocha-Campos, 1983; Santos et al. 1996 and Vesely et al., 2015). The areal scouring landscape and major subglacial erosion landforms indicate a generally fast and highly erosive ice flow (Glasser and Bennett, 2004; Kleman et al., 2006). After this interval of subglacial erosion, and following the retreat of the glacier, the basal nonconformity was defined by topographic highs

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and lows that controlled the deposition of the succeeding deposits.

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4.2. Subglacial and deglacial deposits

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4.2.1. Facies Association 1 (FA-1)

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FA-1 is composed of facies Fsh and FCm (Table 1), and confined to the bases of glacial troughs. It overlies and onlaps Precambrian granites (Fig. 5A). This facies association contains up to 10 m of black shale, which is eventually replaced by mm-scale, normally-graded siltstone

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to mudstone couplets. A few disconnected outcrops to the southwest of the study area suggest that this unit may be thicker, filling even deeper incised glacial troughs. Gamma-ray readings are

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usually quite high (up to 560 cps).

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Black shales are typical of anoxic settings. This unit likely corresponds to the Lontras Shale, a well-known stratigraphic marker in the Itararé Group that is often used for regional correlations (Schneider et al., 1974; França and Potter, 1988). The Lontras Shale includes

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assemblages of marine fossils, such as brachiopods, bivalves, and conodonts (Rocha-Campos and Rösler 1978; Simões et al., 2012; Neves et al., 2014; Wilner et al., 2016). Therefore, FA-1 is

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interpreted to record a major marine flooding event. However, as the study area is believed to have been a marginal setting during deposition of the Lontras Shale, the unit occurs as disconnected bodies confined to topographic lows, making it difficult datum to use for stratigraphic correlation. The thin, normally-graded silt-to-mud beds within the black shale are interpreted as distal turbidity plumes. High gamma ray readings are attributed to its high organic matter content.

4.2.2. Facies Association 2 (FA-2)

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FA-2 is an interval composed of sandstones and associated mudstones, up to 30 m thick. Like the previous facies association, it is completely confined to glacial troughs scoured into the Precambrian basement (Fig. 3). This interval is composed of facies Sg, Si, Hsm and FCm (Table 1). Sandstone beds thicken upwards, with a resulting increase in the sand/mud ratio, but pinch out as they onlap onto the trough walls. The sandstone bodies have low gamma ray signature, typically ranging from 200 to 250 cps. The sediment transport direction was to the NW, although

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a minor SSW component is also present. The contact with the underlying FA-1 is sharp.

Individual sandstone beds are up to 2 m thick, fine- to medium-grained, and usually

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amalgamated (facies Si), forming bedsets up to 5 meters thick (Fig. 5B and C). Sandstones can also occur as normally-graded beds (facies Sg) up to 3 m thick. These latter beds are mostly

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massive, with only their upper 20 to 10 cm showing normal grading (Fig. 5D). Sandstone bedsets are intercalated with heterolithic intervals (Fig. 5E), consisting of thin (mm- to cm-scale),

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normally-graded sand-to-mud couplets exhibiting abundant sinusoidal ripple-cross lamination

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(facies Hsm). These beds often contain plant debris.

Reverse grading, intra-bed scours, and alternation of planar-bedding and ripple cross lamination within the same deposit all suggest deposition in long-lived flows with waning and

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waxing pulses, which we propose can be ascribed to flood-driven hyperpycnal flows (Mulder

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and Syvitski 1995; Mulder et al., 2003; Plink-Björklund and Steel, 2004; Zavala et al., 2006) or waxing, depletive flows (Kneller, 1995; Kneller and Buckee, 2000). Fluctuations in flow velocity may be related to peaks in meltwater-derived river discharge, resulting in sandstone

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beds with the primary structures described above. Hyperpycnal flows occur when fluvial discharge is denser than the water in the receiving basin (Bates, 1953), which is more common in

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lakes than marine settings (Mulder et al., 2003; Plink-Björklund and Steel, 2004). However, in areas such as Alfredo Wagner, the topography of the receiving basin may have lead to basinal restriction, and freshwater input from melting glaciers could potentially create favorable conditions for the development of hyperpycnal flows. In addition, plant debris within sandstone beds has often been interpreted as evidence for gravity-flow deposits related to river floods (Plink-Björklund and Steel, 2004).

Massive, normally-graded beds are likely related to short-lived, surge-type turbidity currents (Lowe, 1979; Kneller, 1995; Mulder and Alexander, 2001; Meiburg and Kneller, 2010) with delayed grading (Lowe, 1982). Surge-type beds are often associated with downslope resedimentation processes. In this case, high sedimentation rates associated with high meltwater 9

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discharge may have caused rapid progradation of the delta, resulting in over-steepening, instability, and the eventual collapse of the delta-front facies. However, this type of sandstone bed is only a minor component of FA-2.

Evidence for the northwestward migration of current ripples is in agreement with glacial striations and trough orientations. Those related to a SSW secondary mode are interpreted to be the result of currents deflected by the trough walls. The thickening-upward pattern and increase

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in the sand/mud ratio can be interpreted either as a result of delta progradation, or a consequence of flow ponding (Sinclair and Tomasso, 2002; Hodgson and Pickering, 2007). Low gamma ray

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readings are consistent with clean sandstones with low mud content.

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4.2.3. Facies Association 3 (FA-3)

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Deposits of FA-3 are poorly exposed and scarce throughout the study area, being restricted to topographic lows in the basement bedrock (Fig. 3). The unit is composed of the

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Gmso lithofacies (Table 1), and it is associated with striated pavements found on the floors of glacial gouges cut into Precambrian basement. These gouges are a shallower class of depressions

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relative to the glacial troughs described above, where FA-1 and 2 are typically found (Fig. 3).

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FA-3 is composed of angular to subangular granite clasts, floating within a medium- to coarse-grained sandstone matrix (Fig. 6A); individual deposits are up to 0.6 m thick. Granite clasts are faceted, and range from granule to pebble size. They are randomly distributed

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throughout the deposit with no clear orientation, and resemble the underlying bedrock. Deposits show complex relationships with basement rock on a local scale, and seem to have been injected

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into fissures in the underlying granites (Fig. 6B). Deformation of the conglomerates is more pronounced in these injected deposits. Gamma ray readings are typically high (Fig. 3).

The restricted distribution of FA-3 in negative relief areas of the bedrock surface, and the oligomitic nature of the conglomerate (which is mostly composed of chaotically distributed angular clasts), suggest that these represent lodgment tillites (Dreimanis and Schlüchter 1985; Evans et al. 2006). Granite clasts within the deposits are similar to the basement underneath, implying a local source. Furthermore, their association with glacial striations and sheared surfaces indicates that these deposits were generated at the base of a sliding glacier, as a result of frictional drag (Smith, 1984; Dreimanis and Schlüchter 1985; Hart and Roberts 1994; Evans et al. 2006; Benn and Evans, 2010). Abrasion and plucking of the substrate likely lead to the 10

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plastering of basal debris into depressions in the Precambrian bedrock (Dreimanis and Schlüchter 1985).

The apparent injection of tillites into weakness zones in the bedrock may be related to ice loading during glacial advances, producing high pressures and drag forces. Differential pressure may cause joints to form in the bedrock, or glacial plucking may exploit pre-existing joints, facilitating the injection of tills (Boulton, 1979; Smith, 1984). This unit is uncommon in the

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Itararé Group, mainly due to its low preservation potential. Most of the matrix-supported conglomerates described in the Itararé Group literature are not true lodgment tillites, as they are

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the result of resedimentation processes (e.g., d´Avila, 2009), or concentrated density flows (e.g., Aquino et al., 2016; see FA-4). High gamma ray signatures (up to 400 cps) are probably due to

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the high granite clasts content.

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4.2.4. Facies Association 4 (FA-4)

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FA-4 is generally confined to shallow topographic lows, filling up glacially-carved gouges (Figs. 7A and B; see also Figs. 4A, B and E). We have divided FA-4 into three subfacies (a, b and c) based on facies associations, spatial distribution, and genetic relationships. Subfacies

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FA-4a represents the base of FA-4, sitting either directly on Precambrian granites (Fig. 7C) or

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FA-3 deposits, but its characteristic lithofacies also occur intercalated with subfacies FA-4b and FA-4c. It is represented by deposits of facies FSm (Table 1), reaching up to 3 m in thickness. It includes granule- to boulder-sized clasts with smooth-faceted shapes, which disrupt the

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surrounding laminae (Fig. 7D). Subfacies FA-4a is deformed and slumped in places.

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Subfacies FA-4b is up to 6 m thick, and is composed of facies Gmsp (Table 1; Fig. 7E). It sits directly on the Precambrian basement wherever subfacies FA-4a and FA-3 are absent. In general, subfacies FA-4b is represented by an amalgamation of several diamictite beds, forming lenticular bodies separated by mudstones with outsized intraclasts, and very thin sandstone intervals of subfacies FA-4a. The majority of subfacies FA-4b is contained within the walls of glacial gouges. Diamictite beds are thickest (up to 0.5 m) along the central axes of these gouges, pinching out toward the gouge walls (Fig. 7A). However, some of the upper beds of subfacies FA-4b seem to have spilled out of the topographic lows formed by the gouges, draping over the irregular surfaces of the surrounding local highs (Fig. 3). These deposits are apparently interbedded with and/or replaced by subfacies FA-4c along glacial gouges.

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Subfacies FA-4c is up to 4 m thick, and either sits directly on Precambrian basement or on subfacies FA-4a deposits (Figs. 8A and B). This interval consists of sigmoidal beds (i.e., gently dipping clinoforms) ranging from 0.20 to 0.60 meters thick (facies Gcs, Sm and Sp), which fill up at least two minor glacial gouges (Fig. 3). The succession displays a fining-upward trend. These deposits transition, over short distances in the downflow direction, into fine- to medium-grained, moderately- to poorly-sorted sandstone beds with undulating geometries (Fig. 8C; facies Sd, Sr, and Sh). Thicker sandstone beds are separated by interbeds of the FSm

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lithofacies. The tops of sandstone beds show soft-sediment striations, in the form of grooves with marginal berms, indicating NW transport (Fig. 8D). The progradation of clinoforms indicates

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transport in an E–W orientation, transverse to the axis of these depressions (Fig. 3).

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Large, smooth-faceted intraclasts that disrupt the surrounding laminae, matrix- and clastsupported conglomerates, and coarse-grained sandstones defining small clinoforms, are all likely

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indicators of subaqueous grounding-line fan or outwash fan deposits (Lønne, 1995; Koch and Isbell, 2013; Aquino et al., 2016). FA-4 is interpreted to have developed along a proximal,

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calving glacier margin, and was deposited via a combination of several sedimentary processes. Three main processes seem to have operated during deposition of the grounding-line fan, all related to meltwater discharge from subglacial tunnels: 1) deposition of fine-grained sandstone

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and mudstone from meltwater hypopycnal plumes (Powell, 1990), or hyperpycnal flows

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associated with ice-rafted debris (IRD) during the progressive melting of calving icebergs (subfacies FA-4a; Gilbert, 1990; Brodzikowski and Van Loon, 1991); 2) efflux jets of concentrated (or hyperconcentrated) gravity-driven flows (Mulder and Alexander, 2001), or

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cohesionless debrites (Postma, 1986), during peaks in discharge (subfacies FA-4b); and 3) bedload dominated, unidirectional currents (Lønne, 1995, Hornung et al., 2007) and deposition in

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minor subaqueous deltas close to the ice margin (subfacies FA-4c; Powell, 1990; Russell and Arnott, 2003; Aquino et al., 2016).

Subfacies FA-4a records minor intervals of decreased meltwater discharge from subglacial tunnels, and floating ice derived from tidewater glaciers. In contrast, subfacies FA-4b records transient increases in meltwater efflux, though it may also be related to the relatively steep slopes of grounding-line fans resulting from resedimentation of subaqueous jet deposits (Powell, 1990). Facies similar to those in the lower part of subfacies FA-4b were earlier interpreted as lodgment tillites by Rocha-Campos et al., (1988).

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Soft-sediment striations on the sandstone beds of subfacies FA-4c are related to ice-keel scouring (Woodworth-Lynas and Dowdeswell 1994; Eyles et al., 2005), and are comparable to features described elsewhere in the Itararé Group (Santos et al. 1992; Vesely and Assine, 2002; Vesely et al., 2015). The presence of dropstones within subfacies FA-4c corroborates a calving glacier-margin setting, with drifting icebergs dragging through the substrate as they progressively melt. The fining-upward sequences are associated with the backstepping of FA-4

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during glacial retreat, and the ensuing deposition of FA-5 (see below).

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4.2.5. Facies Association 5 (FA-5)

FA-5 is widespread throughout the study area, and was used as a datum for stratigraphic

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correlation (see Fig. 3). This unit is not confined to local topographic lows; it onlaps the basement in the northeastern part of the study area (see Fig. 3), and rests directly on FA 4 strata

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elsewhere. FA-5 is up to 6 m thick, and is composed of facies Fsh and FCm (Table 1) with abundant granule- to pebble-sized clasts with a basement affinity (Fig. 9). It also contains intervals of very thin, normally-graded silt-to-mud couplets. The gamma ray signature is

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typically quite high (up to 560 cps).

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Like FA-1, the deposition of black shales in FA-5 indicates an anoxic environment. FA-5 is also interpreted as the record of a significant marine flooding event (Rocha-Campos and Rösler 1978; Simões et al., 2012; Neves et al., 2014; Wilner et al., 2016). Basement-affinity

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clasts within the black shales are interpreted as dropstones, and indicate the presence of floating ice derived from calving tidewater glaciers. Although this unit was attributed to the Lontras

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Shale (Schneider et al., 1974; França and Potter, 1988) by Rocha-Campos et al., (1988), we believe that it may be correlative with the younger Passinho Shale (Schneider et al., 1974, Puigdomenech et al., 2014), with the lower black shale interval (FA-1) being equivalent to the Lontras Shale. Its high gamma ray readings are related to its high organic matter content.

4.2.6. Facies Association 6 (FA-6) FA-6 can be subdivided into ‘a’ and ‘b’ subfacies based on their associations and spatial distribution. Subfacies FA-6a is up to 22 m thick, and is widespread throughout the study area (Fig. 3). It is composed of the Hssm lithofacies (Table 1). Its lower contact with FA-5 is sharp, 13

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and is either marked by an abrupt change in color, or a sudden loss of outsized clasts. These deposits are interbedded with thin, very fine-grained sandstones that show ripple-cross lamination indicating transport to the NW. Subfacies FA-6a deposits onlap the Precambrian basement towards the northeastern margin of the study area (Fig. 3). The unit includes at least two major sandy lenses (subfacies FA-6b). Gamma ray readings are usually high (up to 410 cps).

Lenses of subfacies FA-6b are thickest along their central axis (up to 8 m), and pinch out

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towards the margins of larger erosional features (paleovalley). These lenses consist of several amalgamated beds of the Sg and FCm lithofacies (Table 1). In general, these lenticular beds

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thicken upward, lack coarse-grained sandstones, and are fairly well sorted. Some sandstone beds are partially scoured (Fig. 10A), forming small depressions filled with thin-bedded, very fine-

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grained sandstone and mudstone (Fig. 10B). Paleocurrent indicators suggest flow directions

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ranging from W to NE, with a mean vector to the NW.

Subfacies FA-6b is slumped and deformed along the axes of these lenticular bedsets.

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Sandstone beds are fragmented, mainly via stretching, into blocks up to several meters across (Fig. 10C). The blocks are mostly massive, but some preserve primary sedimentary structures, mainly in the form of ripple cross lamination. In general, the sandstone blocks become smaller

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and less common toward the tops of lenticular bedsets. Minor fragments of diamictites are also

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dispersed throughout the slumped zones. Gamma ray readings of subfacies FA-6b show low cps values.

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Normally-graded, massive to stratified sandstone beds suggest short-lived, surge-type turbidity currents (Lowe, 1979; Kneller, 1995; Mulder and Alexander, 2001; Meiburg and

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Kneller, 2010), representing a second expressive turbidite system in the Rio do Sul Formation that has only rarely been recognized (Castro, 1980; Puigdomenech et al., 2014). The vertical transitions within individual beds suggests a initially high-density turbidity current, associated with frictional freezing, followed by a fully turbulent low density flow and mud fallout (i.e., a bipartite turbidity current). Scour features may be related to the migration of minor channels within the larger channel represented by the lenticular sandstone bedset. This interpretation is supported by the wide range of apparent palaeocurrent directions. The channelized bodies (subfacies FA-6b) are enclosed within a mudstone-rich succession (subfacies FA-6a) ascribed to a prodelta setting. Subfacies FA-6b likely represents channelized, surge-type turbidity currents associated with proglacial delta front failures. Delta front collapse may have been promoted by

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the high sedimentation rates associated with a proglacial setting, and the resulting deposition of a sand bar above muddy prodelta deposits (Suss et al., 2014). 5. Depositional Model

The preserved succession in the Alfredo Wagner region consists of three glacial cycles (1 to 3); the facies associations described above were deposited during these intervals of ice

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advance and retreat (Fig. 11). FA-1 to -6 are stacked in a vertical order and record the local glacial scenario and its depositional setting. Subglacial and deglacial sedimentation were mainly

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controlled by the pre-existing glacially-carved topography, and fluctuations of the glacier margin

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related to climate and relative sea level.

The first glacial cycle (Figs. 11A and B) includes an initial phase of advancing ice, which

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eroded into the Precambrian substrate and produced a series of subglacial erosion landforms manifested as linear troughs (Fig. 11A). These landforms are diagnostic of a warm-based glacier,

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which flowed toward the northwest via basal sliding, sculpting the granite bedrock and creating an areal scouring landscape. This landscape was probably produced by a combination of fast ice sliding, abrasion, and high rates of subglacial meltwater discharge, producing glacial troughs

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large enough to contain part of the overlying sedimentary succession. As the ice margin retreated

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to the southeast, a significant marine flooding event took place, as recorded by a regional-scale, dropstone-free black shale unit (FA-1; Fig. 11B). During this phase, the excavated glacial valley was progressively flooded, and became a fjord. The sediments deposited during this marine

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flooding event constitute an important stratigraphic marker in the Itararé Group, known as the Lontras Shale (Schneider et al., 1974; França and Potter, 1988), and may record the period of highest sea level within the studied interval. This black shale, and the succeeding unit, are both

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confined to the deepest negative depressions on the nonconformity surface. The high organic matter content of this shale indicates anoxic conditions, while the accumulation of land plant debris (wood fragments, spores, and pollen) and algae records an amelioration of the climate during this interval of inlet flooding.

The second glacial cycle (Fig. 11C to F) records a more complete succession, one that includes both the waxing and waning phases of the valley glacier. The initial phase of advancing ice is marked by a relative sea level fall. It is recorded geological by the abrupt superposition (i.e. correlative conformity) of hyperpycnal turbidites (FA-2) above the black shales of FA-1 (Fig. 11C). Plant debris within the hyperpycnal turbidites indicate fluvial erosion of a vegetated 15

ACCEPTED MANUSCRIPT C. Fallgatter and P.S.G. Paim

landscape, suggesting that an outwash plain and delta developed between the margin of the advancing ice and the marine waters of the fjord. The hyperpycnal flow deposits are fully constrained by the topography, and constitute the infill of the largest glacial troughs in the study area (conduits for underflow currents). Deposition of flood-derived underflows within a prodelta environment indicates shallower water, probably on the order of tens of meters. Features related to direct glacial influence, such as dropstones and soft-sediment striations, are absent in this

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initial phase of deposition.

The following phase records the progressive advance of ice through the study area (Fig.

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11D), represented by glacial gouges and lodgment tillites (FA-3). Subglacial tills were lodged into topographic lows in the basement, or injected into the bedrock (subglacial injectites),

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implying intense friction at the base of the glacier. The restriction of lodgment tills to the shallow, striated gouges attests to an additional phase of excavation and scouring. Differential

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abrasion of the substrate may have occurred, due to differences in pressure, variable ice velocity, or non-uniform concentrations of debris at the base of the glacier (Boulton, 1979). A warming of

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the climate and consequent retreat of the ice margin followed this pulse of glacial advance (Fig. 11E). During this retreat phase, meltwater-derived grounding-line fans and associated ice-rafted debris filled in and over-spilled the glacial gouges (FA-4). The presence of dropstones and

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iceberg-keel scouring of soft marine sediments suggest calving of the tidewater glacier front. As

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deglaciation proceeded, melting caused an increase in relative sea level, resulting in an extensive marine flooding event and the re-establishment of a fjord environment. This transgression is recorded by the dropstone-bearing black shales of FA-5 (Fig. 11F), which imply that the glacier

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was in contact with a stagnant, anoxic, and relatively deep water body.

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The third glacial cycle (Fig. 11G) is incomplete, and is represented by an initial phase of ice advance and relative sea-level fall. Like the base of glacial cycle 2, it contains an abrupt superposition (i.e., correlative conformity) of turbidites (FA-6) above the previous transgressive black shales (FA-5). No evidence for direct glacial influence was observed within FA 6, however, high rates of fluvial sediment supply from meltwater allowed for the formation of a delta at the fjord head. Overloading of delta front sands may have triggered channelized, surgetype turbidity currents, resulting in the deposition of a second expressive turbidite succession (Fig. 11G).

6. Glacial paleogeography 16

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During the LPIA, The Gondwanan supercontinent experienced multiple episodes of glacial advance and retreat. Paleogeographic reconstructions of the eastern margin of the Paraná Basin suggest the presence of an ice sheet (Crowell and Frakes 1970), the lobate termini of ice streams (Santos et al. 1996), or several smaller ice lobes (Vesely et al., 2015). Subglacial erosion features in the eastern Paraná Basin show a consistent NW orientation, with minor deviations to the north (in the present-day continental configuration), suggesting a rather constrained ice flow. The presence of ice lobes cannot satisfactorily explain the ubiquitous NW trend of ice motion

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along the eastern margin of the Paraná Basin. The NW trend of ice movement was likely a response to the position of the Paraná Basin depocenter, possibly facilitated by pre-existing,

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NW-oriented tectonic lineaments in the Precambrian basement (Zalán et al., 1990).

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Reconstructing the positions of South America and southern Africa during the Carboniferous (Fig. 12A), an ice center located in the Huab and Kunene areas of western

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Namibia (the Windhoek Ice Cap; Santos et al. 1996, Stollhofen et al., 2000) can be identified as the most probable source of ice to the eastern Paraná Basin (Fig. 12B; Santos et al. 1996; Vesely

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et al., 2015). Deglacial deposits in the Namibia area are confined within paleovalleys (Martin, 1981; Visser, 1987), suggesting that ice was topographically constrained during its flow out from the Windhoek ice sheet. The presence of paleovalleys in southern Africa led Santos et al. (1996)

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to interpret these flows as "outlet glaciers", which had ample room of spread out into wide ‘ice

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lobes’ as they reached their terminus at the Paraná Basin, but in fact these constrained flows may have extended all the way into Brazil.

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In contrast to the lobate model, the data presented here clearly show that part of the Itararé Group was confined to a paleovalley, with local basement highs and lows controlling the

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earliest deglacial deposition. This implies that a confined glacier slid along its base, and carved into the Precambrian basement, instead of flowing in the less erosive form of ice lobes. This paleovalley would have extended into the nearby Vidal Ramos area (Fig. 12C), about 35 km north of Alfredo Wagner, where deposits confined to a glacial trough have also been postulated (Puigdomenech et al., 2014; Fallgatter, 2015). This connection is suggested by the relative positions of the two areas, the alignment of glacial striations (Fig. 12C), and the similar distribution of facies. Further north, a connection between the Alfredo Wagner paleovalley and the Rio do Sul embayment (Canuto, 1993; Santos et al. 1996) is highly possible, and to the southeast, the paleovalley may have connected with one of the similar topographic features that have been described in Namibia. Alternatively, sediments sourced from local hinterland glaciers at the head of the fjord between Brazil and Namibia is also plausible. 17

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Taking into account that deposition was confined to valley glaciers in the study area, ice that reached the eastern margin of the Paraná Basin can either be explained by ice streams flowing out from continental ice sheets, or by alpine outlet glaciers flowing out from mountain ice fields. A connection with an ice cap is more probable, due to the presence of elevated areas near the pole. Ice streams are corridors of fast-flowing ice, typically hundreds of kilometers long and tens of kilometers wide (Bennett, 2003). They form the arteries through which most of the

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ice and sediments are discharged from ice sheets and ice caps (Bennett, 2003; Ben and Evans, 2010). Corridors of ice flow bounded by rock may be classified as ‘topographically controlled

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outlet glaciers’ (Bentley, 1987), or as ‘topographic ice streams’ (Stokes and Clark, 1999). Topographically-controlled ice streams derived from the Windhoek Ice Cap may have been the

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predominant type of ice flows in both Namibia and southern Brazil. As they advanced towards the Paraná basin depocenter, these outlet glaciers passed into confined or unconfined tidewater

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termini, with calving icebergs drifting away from the ice front.

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In light of the probable link between the southeastern margin of Brazil and southwestern margin of Africa during the LPIA, an accurate paleogeographic reconstruction should contain elements from several of the models discussed above. Any reconstruction should include an ice

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cap (Visser, 1987), from which topographically controlled ice streams (glacial valleys) flowed

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downslope across Namibia and the eastern margin of the Paraná basin, before finally spreading out over the flatter, unconfined inner areas of the basin to produce lobate termini, as postulated by Santos et al., (1996). Alternatively, glaciers may have overtopped the valley walls in some

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parts of the eastern Paraná Basin, and flowed in a more unconstrained fashion as proposed by

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Vesely et al. (2015).

7. Conclusions

This study allows us to draw significant conclusions about the erosional and depositional processes associated with the LPIA, and recorded in the Itararé Group. A detailed study of the basal nonconformity reveals partially exhumed glacial troughs and gouges, carved into Precambrian basement. These large (10-100 m wide) subglacial erosion features were formed during two glacial cycles, in response to glacial abrasion and meltwater discharge. The floors and walls of these troughs and gouges are usually polished, and associated with grooved striations. The depressions were filled by deglacial sediments, which thin toward the margins of

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the troughs and onlap the walls. To create such extensive areal scouring, the glacier would have to have been warm-based, grounded, fast moving, and highly erosive.

The presence of deglacial deposits filling up glacially-carved structures strongly suggests that deposition of the Itararé Group was partially controlled by the pronounced basin floor topography, and contradicts other paleogeographic models proposed for the eastern Paraná Basin. It is likely that the eastern Paraná Basin was characterized by ice streams flowing out

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from an ice cap centered in Southern Africa via topographically controlled corridors, rather than as unconstrained ice lobes, as previously suggested, though ice lobes may have existed at the

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termini of ice streams, where the ice was free to spread. Based on these constraints, we can infer a paleogeography that comprises an ice cap, feeding topographically controlled ice streams with

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ice lobes forming at their termini, during the deposition of the Itararé Group. Continued mapping of the basal nonconformity following the ice stream track, combined with information about its

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morphology obtained from geophysical surveys, may provide valuable insight into the style and dynamics of glaciation during deposition of the Itararé Group, further improving

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paleogeographic reconstructions of the LPIA in Gondwana.

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Acknowledgments

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The authors would like to express their gratitude to BG Brasil E&P Ltda. for funding this research. We would also like to thank the Agencia Nacional do Petróleo (ANP) and the Universidade do Vale do Rio dos Sinos (UNISINOS) for their consistent support. We extend our

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special thanks to Carolina D. Aquino and Ilana Lehn, for their help with fieldwork, and to the editor of this manuscript, Thomas Algeo, as well as Christopher Fielding and an anonymous

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reviewer for their valuable feedback. We are grateful to the Alfredo Wagner community, which allowed us access to the private properties where most of the outcrops described in this study are located. The second author wishes to acknowledge the Conselho Nacional de Desenvolvimento Cientifico e Tecnológico (CNPq) for their long-term support.

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Fig. 1: Stratigraphic chart of the Itararé Group (modified from França and Potter, 1988, after França et al., 1996).

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Fig. 2: A) Location of Santa Catarina State in southern Brazil. B) Location of the Alfredo Wagner area, in the Santa Catarina Shield. C) Geological map of the study area displaying positions of logged stratigraphic sections and glacial striations, which are oriented to 340° (Rio Adaga), 357° (Rio Lessa), 350° (Rio Picadas I) and 345° (Rio Picadas III). D) Composite stratigraphic section, showing the vertical distribution of the facies associations described in this study. It = Itararé Group, SC = Santa Catarina State, FA = Facies Associations.

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Table 1: Sedimentary facies within the lower part of the Rio do Sul Formation in the Alfredo Wagner area.

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Fig. 3: A) Sedimentary and gamma-ray logs acquired across an 8 km long, strike-oriented stratigraphic section in the Alfredo Wagner area. The irregular topography of the Precambrian bedrock represents a surface of areal scouring that constitutes the basal nonconformity of the Itararé Group. Notice that FA-1 to -6 are confined to depressions of distinct scales along the basal nonconformity. Rose diagrams presenting the palaeocurrents readings are associated with distinct stratigraphic intervals. B) Section between Rio Lessa and Rio Adaga outcrops. C) Line drawing of (B) showing two different magnitudes of subglacial erosional features and their infill by deglacial sediments.

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Fig. 4: Subglacial erosion features in Precambrian basement rocks. Sense of ice movement is indicated by solid arrows. A) View of a partially exhumed glacial gouge, exhibiting a polished and striated floor, and filled by deglacial strata (Rio Adaga II outcrop). B) Strike and dip view of a glacial gouge carved into Precambrian granites, and filled by FA-4 deposits. Note the half-cylindrical exposure of deglacial beds onlapping the right side of a gouge wall (Rio Lessa outcrop). C) Plan view of glacial striations on a gouge floor (Rio Picadas III outcrop; see RochaCampos et al. 1988). D) Poorly preserved, striated pavement on an extremely polished granite surface, exposed along Adaga River. E) Grooved (p-form) surface in the polished Precambrian granites of a gouge wall (Rio Lessa outcrop). Note grooved striations superimposed on the p-form along the wall, and the onlap of the deglacial deposits (dotted half arrows). A few friction cracks are also visible.

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Fig. 5: Examples of FA-1 and FA-2. A) Black shales (facies Fsh of FA-1) onlapping the granites of a partially exposed glacial trough wall (yellow dotted line; Rio Adaga outcrop). B) Thick, sharp-based sandstone bedset of FA2. C) Thick sandstone bed (facies Si), showing an alternation of planar-bedded and ripple cross-laminated intervals with intra-bedding erosion surfaces (white arrows). Person for scale. D) Sandstone bed showing a thick, massivelybedded interval, overlain by a 10 cm-thick interval displaying normal grading and planar-bedding, followed by ripple cross-lamination (facies Sg). E) Thin-bedded sand-to-mud couplets, with abundant sinusoidal ripple cross lamination (facies Hsm). Examples B through E are from the Rio Adaga IV outcrop.

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Fig. 6: A) Matrix-supported conglomerates (facies Gmso of FA-3), showing chaotically distributed, poorly sorted, angular to subangular clasts of granite. B) Deposits of FA-3 lodged into a granite joint, believed to represent an injected conglomerate (facies Gmso). Examples are from the Rio Lessa outcrop.

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Fig. 7: Examples of FA-4. A) Oblique section through a nearly 8 m deep glacial gouge, showing the onlap of facies Gmsp deposits onto one of its walls (Rio Lessa outcrop). Solid yellow lines represent main bedding surfaces. B) Along-strike view of a glacial gouge, filled by deposits of subfacies FA-4a and FA-4b (facies FSm and Gmsp). This outcrop was described by Rocha-Campos et al. (1988). C) and D) Examples of facies FSm, including granule- to boulder-sized dropstones with a local basement affinity, embedded within fine grained sandstones and mudstones. Examples are from the Rio Picadas I and Rio Adaga II outcrops, respectively. E) Detailed view of facies Gmsp (subfacies FA-4b), the dominant facies in FA-4 (Rio Picadas II outcrop).

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Fig. 8: Examples of FA-4c. A) Several conglomeratic, sigmoidal sandstone beds. Clinoforms indicate progradation to the west (facies Sp). Note that these beds overlie slumped deposits of FA-4a. B) Downcurrent transition from conglomeratic sandstones (photo A) to fine- to medium-grained sandstones with undulating geometry (facies Sp and Sr), onlapping one of the glacial gouge walls. C) Sandstone beds, with preserved dune bedforms and associated cross-stratification (facies Sd). D) Soft-sediment striations on top of a sandstone bed (facies Sm). Note that grooves are associated with marginal berms (compass for scale). Examples A and B are from the Rio Adaga II outcrop, examples C and D are from the Rio Picadas I outcrop.

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Fig. 9: FA-5, consisting of dropstone-bearing black shales (facies Fsh). Dropstones vary in size from A) granules to B) pebbles, and are composed mainly of local basement-derived clasts. Examples are from the Rio Lessa outcrop.

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Fig. 10: Examples of FA-6. A) Cut-and-fill deposits of sandstone with lenticular geometries (facies Sg). Note that the erosive surface truncates primary stratification of the underlying bed. B) Normally-graded sandstone beds, showing a massively-bedded basal interval followed by planar-bedded and ripple cross-laminated beds (facies Sg). C) Slumped zone, represented by thick, disrupted sandstone beds and blocks. These deposits are restricted to the axes of major sandstone lenses in subfacies FA-6b. Examples A and B are from the Rio Adaga III outcrop, C is from the Rio Picadas II outcrop.

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Fig. 11: Reconstruction of the glacial landscape and proglacial depositional stages recorded in the Itararé Group in the Alfredo Wagner area. Schematics A through G represent three proposed glacial cycles, defined by the waxing and waning of glacial ice streams (see text for details).

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Fig. 12: A) Reconstruction of the South American and African continents during the LPIA. The Paraná Basin is indicated by light grey color in Brazilian territory. B) Juxtaposition of the Brazil and Namibia areas (modified from Stollhofen et al., 2000; Vesely et al., 2015 and Tedesco et al., 2016). Note that Brazil was reoriented in accordance with reconstruction in (A). Oriented paleovalleys (letters) and subglacial erosion landforms (numbers) described to date are labeled (see Rosa et al., 2016 for references). C) Regional geological map, showing the area covered by the paleovalley proposed in this study (modified from Wildner et al. 2014), showing its extent and approximate dimensions. Note the location of the striated surfaces described in this work: 9a = BR 282 road section, RochaCampos et al., (1988); 9b = Rio Lessa; 9c = Rio Adaga, and 9d = Rio Picadas.

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