Journal of Sedimentary Research, 2021, v. 91, 611–635
A LITHOFACIES ANALYSIS OF A SOUTH POLAR GLACIATION IN THE EARLY PERMIAN:
PAGODA FORMATION, SHACKLETON GLACIER REGION, ANTARCTICA
LIBBY R.W. IVES AND JOHN L. ISBELL
University of Wisconsin–Milwaukee, Department of Geosciences, Lapham Hall, 3209 North Maryland Avenue, Milwaukee, Wisconsin 53211, U.S.A.
ABSTRACT: The currently favored hypothesis for Late Paleozoic Ice Age glaciations is that multiple ice centers
were distributed across Gondwana and that these ice centers grew and shank asynchronously. Recent work has
suggested that the Transantarctic Basin has glaciogenic deposits and erosional features from two different ice
centers, one centered on the Antarctic Craton and another located over Marie Byrd Land. To work towards an
understanding of LPIA glaciation that can be tied to global trends, these successions must be understood on a local
level before they can be correlated to basinal, regional, or global patterns. This study evaluates the sedimentology,
stratigraphy, and ﬂow directions of the glaciogenic, Asselian–Sakmarian (Early Permian) Pagoda Formation from
four localities in the Shackleton Glacier region of the Transantarctic Basin to characterize Late Paleozoic Ice Age
glaciation in a South Polar, basin-marginal setting. These analyses show that the massive, sandy, clast-poor
diamictites of the Pagoda Fm were deposited in a basin-marginal subaqueous setting through a variety of
glaciogenic and glacially inﬂuenced mechanisms in a depositional environment with depths below normal wave
base. Current-transported sands and stratiﬁed diamictites that occur at the top of the Pagoda Fm were deposited as
part of grounding-line fan systems. Up to at least 100 m of topographic relief on the erosional surface underlying
the Pagoda Fm strongly inﬂuenced the thickness and transport directions in the Pagoda Fm. Uniform subglacial
striae orientations across 100 m of paleotopographic relief suggest that the glacier was signiﬁcantly thick to
‘‘overtop’’ the paleotopography in the Shackleton Glacier region. This pattern suggests that the glacier was likely
not alpine, but rather an ice cap or ice sheet. The greater part of the Pagoda Fm in the Shackleton Glacier region
was deposited during a single retreat phase. This retreat phase is represented by a single glacial depositional
sequence that is characteristic of a glacier with a temperate or mild subpolar thermal regime and signiﬁcant
meltwater discharge. The position of the glacier margin likely experienced minor ﬂuctuations (readvances) during
this retreat. Though the sediment in the Shackleton Glacier region was deposited during a single glacier retreat
phase, evidence from this study does not preclude earlier or later glacier advance–retreat cycles preserved
elsewhere in the basin. Ice ﬂow directions indicate that the glacier responsible for this sedimentation was likely
ﬂowing off of an upland on the side of the Transantarctic Basin closer to the Panthalassan–Gondwanide margin
(Marie Byrd Land), which supports the hypothesis that two different ice centers contributed glaciogenic sediments
to the Transantarctic Basin. Together, these observations and interpretations provide a detailed local description of
Asselian–Sakmarian glaciation in a South Polar setting that can be used to understand larger-scale patterns of
regional and global climate change during the Late Paleozoic Ice Age.
Strata of the Transantarctic Basin (TAB) contain a complete South Polar
sedimentary record of the global ‘‘ icehouse’’ to ‘‘ greenhouse’’ transition
during the Early Permian (Collinson et al. 1994, 2006; Isbell et al. 2008b).
Sedimentation in the TAB was dominated by glaciogenic processes during
the Asselian–Sakmarian (Isbell et al. 2008c). This interval was part of the
Late Paleozoic Ice Age (LPIA, ~374–256 Ma) (Fielding et al. 2008c;
nez and Poulsen 2013). Widespread glaciation across Gondwana
characterized the LPIA, as did low pCO
, high pO
, generally low eustatic
levels with large magnitude ﬂuctuations, low solar luminosity, and
O and d
C values relative to the rest of the Phanerozoic
(Gastaldo et al. 1996; Raymond and Metz 2004; Monta˜
nez and Soreghan
2006; Fielding et al. 2008d; Rygel et al. 2008; Monta˜
nez and Poulsen
The currently favored hypothesis for LPIA glaciations is that multiple
ice centers (ice sheets or ice caps) were distributed across Gondwana, and
that these ice centers grew and shank asynchronously over the LPIA’s ~80
Myr duration (Fielding et al. 2008c; Isbell et al. 2012; Monta˜
Poulsen 2013; L´opez-Gamund´ı et al. 2021; Rosa and Isbell 2021). The
character, distribution, and resulting sedimentary records of these glaciers
would have been driven by global, regional, and local climatic and
geologic inﬂuences (Isbell et al. 2012; Monta˜
nez and Poulsen 2013;
L´opez-Gamund´ı et al. 2021). The potential for local and regional
heterogeneity of LPIA glaciogenic strata is therefore extremely high. To
work towards an understanding of LPIA glaciation that can be tied to
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global trends, these successions must be understood on a local level before
they can be correlated to basinal, regional, or global patterns.
In this paper, we evaluate the sedimentology, stratigraphy, and ﬂow
directions of four glaciogenic (Pagoda Fm) successions in the Shackleton
Glacier region of the TAB (Fig. 1B). The Pagoda Fm in the Shackleton
Glacier region has not previously been described and analyzed at the level
of detail reported in this study. The Shackleton Glacier region can offer a
different perspective to better-studied areas of the TAB (e.g., the
Beardmore Glacier region) because it was located in a basin marginal
position on the non-cratonic (or, ‘‘Panthalassan proximal’’) side of the
basin during the deposition of the Pagoda Fm (Fig. 1).
Asselian–Sakmarian glaciogenic strata of the Transantarctic Basin
(TAB) occur in discontinuous outcrops along the margin of the East
Antarctic Craton from Victoria Land, near Australia, to Dronning Maud
Land, near southern Africa (Frakes et al. 1971; Collinson et al. 1994; Isbell
et al. 2008c) (Fig. 2A). During the lower Cisuralian, the TAB was a narrow
(~100–200km-wide), trough-shaped basin that formed parallel and
proximal to the Gondwanide margin of the East Antarctica Craton, but
inboard of the Panthalassic margin (Fig. 3B) (Collinson et al. 1994; Elliot
2013; Isbell 2015; Elliot et al. 2017). In the central Transantarctic
Mountains and Victoria Land, glaciogenic strata occur in four sub-basins:
the Ohio Range to the Scott Glacier (Horlick Sub-basin), the Amundsen
Glacier to the Darwin Glacier area (Beardmore Sub-basin), south Victoria
Land (SVL), and north Victoria Land (NVL) (Figs. 1B, 3) (Frakes et al.
1966; Isbell et al. 2008c; Isbell 2010; Cornamusini et al. 2017). The
Shackleton Glacier region is located near the southern edge of the
The origin and nature of the TAB during the Lower Permian is not well
understood. Hypotheses include intracratonic and extensional settings
(Collinson et al. 1994; Isbell 2015; Elliot et al. 2017). Regardless of what
processes drove basin formation at that time, the TAB was a narrow,
trough-shaped basin, with Proterozoic and early Paleozoic basement
shoulders, that paralleled the Panthalassan margin of the East Antarctic
Craton during deposition of the Pagoda Fm (Fig. 3) (Isbell et al. 1997b).
Though there is no evidence for upper Carboniferous to Sakmarian
orogenic activity in the central Transantarctic Mountains or adjacent Marie
Byrd Land, volcanic arcs and tectonic compression were occurring
elsewhere along Gondwana’s Panthalassan margin during that time,
including in eastern Australia, the Andean margin of South America and
Patagonia, the Ellsworth Mountains, Thurston Island, and the Antarctic
Peninsula (Fielding et al. 2001; Elliot 2013; Viza
´n et al. 2017). This same
margin was extremely active and subject to repeated, complex accretion
events throughout the Paleozoic (Veevers et al. 1994; Domeier and Torsvik
2014; Goodge 2020). As a result of this activity, the TAB evolved into a
foreland basin later in the Permian (Collinson et al. 1994; Elliot et al.
2017). During the Lower Jurassic, strata in the central Transantarctic
Mountains were pervasively intruded by sills associated with Ferrar Group
volcanism and the breakup of Gondwana (Elliot 1992).
SEDIMENTOLOGY AND STRATIGRAPHY OF THE PAGODA FORMATION
The Pagoda Fm is the basal unit in the Permian–Early Jurassic Victoria
Group (upper Beacon Supergroup) in the Beardmore Sub-basin of the
TAB. Rare palynomorphs and conchostracans suggest that the Pagoda and
Mackellar fms are Asselian–Sakmarian (Masood et al. 1994; Askin 1998;
Babcock et al. 2002). The Pagoda Fm overlies both the Kukri and Maya
regional erosional surfaces (Figs. 1B, 3) (Collinson et al. 1994; Isbell 1999;
Elliot 2013). The Maya Erosional Surface is a disconformity that separates
Devonian(?) clastics of the lower Beacon Supergroup from the Victoria
Group (Isbell 1999). The Kukri Erosional Surface separates the Beacon
Supergroup from underlying Ross Orogeny intrusions and associated
metasediments. Signiﬁcant relief of at least 150 m occurs on these
unconformities (Fig. 4) (Isbell et al. 1997a, 2008c; Isbell 1999). Both the
Pagoda Fm, and the overlying, post-glacial Mackellar Fm, lap onto the
erosional surfaces, indicating that the Pagoda Fm and its equivalents in
other sub-basins often did not overtop the relief (Isbell et al. 1997a). The
lower Beacon Supergroup units are not present in the Shackleton Glacier
region, and the erosional surface underlying the Pagoda Fm is merged
Maya and Kukri surface (Isbell et al. 2008c). At all sites in this study, the
Pagoda Fm overlies Ross Orogeny granites.
Since their discovery, the Pagoda Fm and its equivalents throughout the
Transantarctic Mountains have been unanimously interpreted as glacio-
genic or glacially inﬂuenced because their predominant lithologies are
massive and laminated, sandy and silty diamictites (Long 1964a; Lindsay
1970a; Coates 1985; Barrett et al. 1986; Collinson et al. 1994; Isbell et al.
2008c). Minor lithologies of the Pagoda Fm include conglomeratic
sandstones, sandstones, mudrocks, and lonestone-bearing mudrocks (Isbell
et al. 2008c). Besides diamictites and lonestones, evidence for a glacial
origin for the Pagoda Fm includes striated and polished basement surfaces,
the prevalence of striated and faceted clasts, and a clear relationship
between local basement composition and lithologies of large clasts in the
diamictites (Lindsay 1969; Coates 1985). Detailed interpretations of
depositional environments have been made for a few Pagoda Fm localities
(Lindsay 1970a; Waugh 1988; Miller 1989; Isbell et al. 2001; Lenaker
2002; Long et al. 2008–2009; Koch 2010; Koch and Isbell 2013) and its
equivalents in Victoria Land (Askin et al. 1971; Barrett 1972; Barrett and
McKelvey 1981; Isbell 2010; Cornamusini et al. 2017), Horlick Mountains
(Frakes et al. 1966; Aitchison et al. 1988), and Ellsworth Mountains
(Ojakangas and Matsch 1981; Matsch and Ojakangas 1991). With few
exceptions, these analyses have invoked subaqueous, glacial-proximal
depositional settings. This is in contrast to early surveys that interpreted
diamictites as subglacially deposited ‘‘tillites’’ (Lindsay 1970a; Coates
1985; Miller 1989; Isbell et al. 1997b).
Isbell et al. (2008c) separated the Permian glaciogenic units in the
Transantarctic Mountains into basin-margin and basinal facies associa-
tions. Basin-margin successions are predicted to occur near basement highs
and along basin margins, are relatively thin (,100 m), contain evidence
for subglacial deformation and erosion, have deformation resulting from
proglacial glaciotectonism, and small (m-scale) gravity-driven deposition.
Basinal successions are thicker (100–500 m), have little-to-no evidence for
subglacial processes, and are more likely to contain stratiﬁed diamictites,
lonestone-bearing mudrocks, mudrocks, and larger (up to tens of meters)
mass-transport deposits. Evidence of grounded ice and grounding-line
processes have been identiﬁed in both basinal (e.g., Koch and Isbell
(2013)) and basin-margin (e.g., Isbell (2010)) facies associations. Based on
its paleogeographic position and Pagoda Fm thickness, the Shackleton
Glacier area is here predicted to contain the basin-margin facies.
In the Shackleton Glacier region, the glaciogenic facies of the Pagoda
Fm are underlain by a non-glacial, lacustrine facies association at a single
site on Mt. Butters (site MB-17). This facies and its depositional
environment are described in detail by Isbell et al. (2001). Below this
contact, the ﬁne-grained lacustrine facies are pervasively sheared, likely
subglacially (Isbell et al. 2001). Lonestones, interpreted to be iceberg-
rafted debris, occur in the lower post-glacial Mackellar Fm in the
Shackleton Glacier region (Seegers 1996; Seegers-Szablewski and Isbell
1998). This suggests that glaciers were still present in the Transantarctic
Basin even after glaciogenic sedimentation was no longer dominant.
STUDY AREA AND METHODS
The sedimentary sections described in this paper were examined as part
of the U.S. Antarctic Program’s helicopter-supported Shackleton Glacier
Deep-Field Camp during the 2017–2018 austral summer (Table 1; Fig.
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FIG. 1.—Generalized geologic maps of study area. Maps are South Polar projections. A) Geologic map of the Central Transantarctic Mountains and Victoria Land, with
relevant outlet glaciers and mountain ranges labeled. Modiﬁed after Elliot (2013), Goodge and Fanning (2016), and Estrada et al. (2016). Box on inset map indicates extent of
this geologic map. Red box on geologic map indicates the extent of ‘‘map B.’’ B) Regional geologic map of the Shackleton Glacier area, noting the locations of sections
described in this study. MM-17 is Mt. Munson, MB-17 is Mt. Butters 1, MBSE-17 is Mt. Butters 2, and RS-18 is Reid Spur. Geology adapted from McGregor and Wade
(1969), and Mirsky (1969), aerial photos from LIMA Landsat imagery (Bindschadler et al. 2008).
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FIG. 2.—Paleogeographic reconstructions of Gondwana near the Carboniferous–Permian Boundary. All maps are south-polar projections. Star indicates the approximate
location of the Shackleton Glacier area. Continent distributions and paleolatitudes are based on Lawver et al. (2011) and copied from Isbell et al. 2012). Note that there are
differences in the positioning of some crustal blocks (e.g., Patagonia and New Zealand) between this reconstruction and Figure 3, which is modiﬁed after Elliot (2013). A)
Yellow regions indicate the modern extent of sedimentary basins containing Late Paleozoic Ice Age strata. Abbreviations include: Falkland Islands–Malvinas (FI), Ellsworth
Mountain block (EM), Antarctic Peninsula (AP), Thurston Island (TI), Marie Byrd Land (MBL), and the Challenger Plateau–western New Zealand (ChP). Basins are adapted
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1B). These sections are located on the Mt. Butters Massif (MB-17, MBSE-
17) on the west side of the Shackleton Glacier, on the east face of Reid
Spur (RS-18) of the Ramsey Glacier, and Mount Munson (MM-17) at the
head of Barrett Glacier. The two sections at Mt. Butters are separated by
approximately 2 km.
Sites were selected for section descriptions based on accessibility,
continuity of exposure, and preliminary observations from previous
expeditions to the area. At MB-17 several shorter sections were measured
to capture lateral variability. Sedimentological data (texture, grain shape,
sedimentary structures, etc.) as well as paleotransport indicators (including
cross stratiﬁcation, primary current lineations, striae, slickenslides, fold
hinge lines, and thrust-plane orientation) were logged in each section.
Measured sections were placed in context using outcrop-scale photographs
taken from helicopters (Fig. 4).
Structural and sedimentary orientations were measured using a Brunton
compass, with the azimuth set to 0008. Measurements were corrected for
magnetic declination using the NOAA Magnetic Declination Estimated
Value Calculator (NOAA 2019). Orientations were corrected for structural
dip and aggregate orientations calculated using Stereonet (v. 11) software
(Allmendinger et al. 2012; Cardozo and Allmendinger 2013). Some
measurements from Mt. Munson were collected by JLI during the 1997–
1998 ﬁeld campaign to the Shackleton Glacier region. New ﬂow-direction
measurements are in Table 2 and site descriptions are in Appendix A.
The following descriptions are of glaciogenic facies associations (FA)
that constitute the Pagoda Fm in the Shackleton Glacier region (Table 3).
We use the term glaciogenic to refer to sedimentary systems whose
components are dominantly derived from glacial erosion and/or transport,
and whose depositional processes are glacier-driven. For example, a
succession that is the result of a plume from a subglacial jet would be
glaciogenic, but largely non-glacially-derived deep-water sediments with
the occasional outsized clast would instead be considered ‘‘glacially
inﬂuenced.’’ The characteristics of distinct sediment grain sizes are
consistent throughout these successions. Very-ﬁ ne-grained sandstones and
shales are black in color. Fine- to medium-grained sandstones are generally
quartz-rich with some lithic and potassium feldspar grains. Cobble- and
boulder-size clasts are sourced from local basement lithologies (Fig. 1B),
and include predominantly phaneritic granitoids, with some gneiss,
quartzite, and gray, ﬁne-grained metasedimentary rocks. All sand-size
and coarser-grained material in the Pagoda Fm occurs in all categories of
particle roundness (angular to well rounded), although ﬁner-grained sands
are typically better-rounded than medium- and coarse-grained sands.
Striations occur on large clasts throughout the Pagoda Fm but are not
common. The lack of striations is possibly due to the hardness of
individual clasts, which are primarily composed of granite, quartz, and
feldspar (Dowdeswell et al. 1985; Bennett et al. 1997).
Massive Sandy Diamictite Facies Association (MSD)
MSD Description.—This diamictite is the dominant FA of the Pagoda
Fm. Similar lithologies occur throughout the Transantarctic Basin. In the
Shackleton Glacier region, the thickness of this FA ranges from 3 m at Mt.
Munson (MM-17) to 73 m at Mt. Butters (MB-17C). Since this diamictite
is almost wholly massive, there is no clear way to further subdivide these
successions. Where exposed, the lower contact overlies either a striated and
polished unconformity with the Queen Maud Batholith (MM-17 and
MBSE-17) or subglacially deformed lacustrine sediments (Isbell et al.
2001; MB-17C). The upper contact of this FA is sharp or erosional with
current-transported facies, including the Cross Bedded Sandstone (CBS)
facies association, the Heterogenous Sandy (HS) facies association, and the
Mackellar Fm. The contact between this FA and the HS facies association
is also gradational where facies HS1 (stratiﬁ ed diamictite) is present above
the contact. The upper part of this FA are intercalated with turbidites (see
LS facies interpretation), and may interﬁnger with stratiﬁed diamictites and
mass-transport deposits (see HS facies association interpretation).
Approximately 90% of this FA is clast-poor to clast-rich diamictite
(Hambrey and Glasser 2003; Hambrey and Glasser 2012) with minor
amounts of sorted sands and gravels (Fig. 5). Clast abundances ﬂuctuate
throughout the succession. Some intervals are sufﬁciently clast-poor that
they could be classiﬁed as muddy sandstones with dispersed clasts (,1%
clasts) (Fig. 5F) (Moncreiff 1989; Hambrey 1994; Hambrey and Glasser
2012). Most clast-rich parts of this diamictite contain 10–15% clasts (Fig.
5A, E), but some very limited areas contain up to ~30% clasts (Fig. 5C).
Clasts range from pebble-size to 4 m in diameter. Clast compositions
includes granite, feldspar, vein quartz, gneiss, and ﬁne-grained metasand-
stone. Clast shape ranges from rounded to angular. Faceted clasts are
common, but bullet-shaped and striated clasts are rare. The matrix is very
poorly sorted, with sizes ranging from muds through granule-size grains.
Matrix grain-size distributions remain constant within and between
outcrops of this facies, though mean matrix grain size increases slightly
in clast-rich sections relative to clast-poor sections.
The diamictite in this FA is massive, with very rare exceptions.
However, broadly deﬁned zones of clast-poor or clast-rich diamictites do
from Isbell et al. (2012). B) Proposed positions of glacial centers during the Early Permian based on ﬂow directions and position of basins and ‘‘ highlands.’’ Illustrated ice
centers are not meant to represent the whole possible extent of each proposed glacier, but where proposed glaciers were likely to be nucleated. The arrows reﬂect ﬁeld
measurements of ﬂow directions reported in the studies cited for each ice center. However, ﬂow directions of glaciers are highly variable, both spatially and temporally, and the
true ﬂow paths of these ancient ice centers were likely much more variable than the arrows on this map. Conﬁdence is based on abundance of available lithologic data, and
both relative and absolute ages. Ice centers are as follows: MBL, the proposed Marie Byrd Land ice center, discussed in this study as the most likely source for the glaciogenic
sediments of the Pagoda Fm in the Shackleton Glacier region (Isbell et al. 1997b; Isbell 2010), a) Uruguay (Crowell and Frakes 1975; Assine et al. 2018; Fedorchuk et al.
2019), b) Asunsci´on (Frakes and Crowell 1969; Fran¸ca and Potter 1988; Limarino et al. 2014), c) Windhoek–Koakoveld Highlands (Martin 1981; Visser 1987; Fran¸ca et al.
1996; Rosa et al. 2016; Tedesco et al. 2016; Assine et al. 2018; Dietrich et al. 2019; Fallgatter and Paim 2019), d) Cargonian Highlands (Crowell and Frakes 1972; Visser
1997; Isbell et al. 2008a; Dietrich et al. 2019), e) Cape–Ventana Fold Belt (Visser 1997; Isbell et al. 2008a; Wopfner 2012), f) East African Thermal Rise (Rust 1975; Wopfner
2012), g) Patagonian Western Magmatic Arc (Pauls 2014; Survis 2015; Marcos et al. 2018), h) Zimbabwe (Wopfner 2012; Dietrich et al. 2019), i) Madagascar–SW India
(Veevers and Tewari 1995; Isbell et al. 2012), j) Chotanagpur and Chhattisgarh (Veevers and Tewari 1995; Dasgupta 2006; Isbell et al. 2012), k) Pilabra–Yilgarn, l) Kimberly
(see Mory et al. (2008); Martin et al. (2019), and references therein, m) Arunta–Musgrave (Mory et al. 2008 and Martinet al. 2019) and references therein, n) Bowen–
Gunnedah–Sydney (Fielding et al. 2008a, 2008b, 2010), o) Galilee (Fielding et al. 2008a, 2008b, 2010; Isbell et al. 2012), p) Wilson (Hand 1993; Rocchi et al. 2011; Jordan et
al. 2013), r) East Antarctic (Isbell et al. 1997a; Isbell 2010), s) Ellsworth (Frakes et al. 1971; Ojakangas and Matsch 1981; Matsch and Ojakangas 1992; Visser 1997).
TABLE 1.—Names and locations of sedimentary sections described in this
Name Geographic Coordinates
Mt. Butters 1 MB-17 S84851.0290W177825.216090 m
Mt. Butters 2 MBSE-17 S84853.0030W177822.354077 m
Reid Spur RS-18 S84847.035 0E178846.6800.62 m
Mt. Munson MM-17 S84845.3590E173841.11805m
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FIG. 3.—Regional geologic context and tectonic setting of the Transantarctic Basin during the Asselian–Sakmarian. A) Stratigraphy of the Beacon Supergroup across
different regions of the Transantarctic Basin. Adapted from Elliot (2013), Cornamusini et al. (2017), and Elliot et al. (2017). B) Tectonic setting of southern Gondwana during
the Permian, adapted from Elliot (2013). Regions of sedimentary deposition are shaded yellow. Note that there are differences in the positioning of some crustal blocks (e.g.,
Patagonia and New Zealand) between this reconstruction and Figure 2. C) Modern extent and isopach map of the Asselian–Sakmarian glaciogenic facies (Pagoda Fm and
equivalents) in the Horlick Sub-basin and Beardmore Sub-basin in the Transantarctic Mountains. Gray areas are outcrops (nunatuks). Lines show isopachs of the Pagoda Fm
(Beardmore Sub-basin) and Buckeye Fm (Horlick Sub-basin) from Isbell et al. (2008c).
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occur. These zones are anywhere from 1 to 30 m thick. While we tracked
these changes in clast abundance vertically through the measured sections,
it was not our impression that these ‘‘zones’’ have any sort of horizontal
organization as might be implied by the terms ‘‘layers’’ or ‘‘ horizons,’’ and
that there are no distinct bounding surfaces between these zones. The
transitions between these zones are gradational. These gradational
transitions occur on the decimeter to meter scale. There was no clear
relationship between transition thickness and any other property of these
rocks. Zones of diamictites with similar clast concentrations cannot be
correlated between outcrops; the distribution of clast-poor and clast-rich
diamictites appears to be unique to each locality.
Crude and chaotic bedding occur very occasionally in the otherwise
massive diamictite. Where bedding can be discerned, the beds are 2–10 cm
thick, laterally discontinuous, and internally massive. Rare sedimentary
structures include ruck structures beneath large clasts (Fig. 5D) and
thinning of beds over large clasts (Fig. 5H). Rare beds of boulders and
cobbles occur in the lower part of this facies at both Mt. Butters sites (Fig.
5G). In all ‘‘boulder beds’’ the clasts were not striated, polished, or
Sandstone and/or conglomerate bodies are also rare within this FA.
These bodies occur in the diamictite and are poorly to moderately sorted.
The grain compositions in these bodies are similar to sand- to gravel-size
grains in the diamictite. Most sand bodies are massive but can contain
layers of discrete grains sizes. However, these layers are often highly
deformed. Bodies of sorted sand and/or gravel occur in two distinct sizes.
Small bodies consist of sorted sediments that most often occur as
irregularly shaped, horizontally elongate lenses, ‘‘whisps,’’ and boudinage-
like bodies that are up to 20 cm thick and 100 cm long (Figs. 5A, B, E, F).
Such structures may occur alone or, more frequently, in bands and zones up
to 1 m thick. Occasionally, diapir-like structures made of sand and/or
gravel, 1 to 2 m thick, are also present. These structures project upward
into the diamict but are not associated with other sand or gravel bodies.
Larger bodies of well-sorted sand or gravel range in thickness from 1 to 3
m and are laterally discontinuous. These bodies have the same grain-size
distributions as the smaller bodies. Large sand bodies are typically
massive, but where stratiﬁcation does exist in the sands and/or gravels the
beds are largely deformed and display water-escape structures. The large
sand bodies thin laterally and have overturned folds on their thicker ends.
In one instance, the main sand body was accompanied by smaller sand
bodies that trailed off away from its thick end (comet-like structure) (Fig.
5F; Sandstone with dispersed clasts). The lower contact between the
comet-like bodies and the surrounding diamictite has a slope of 308to 358
above horizontal, and the underlying diamictite sometimes has a ﬁssile
structure, indicating shearing. Occasionally such contacts overlie smaller
sand lenses, sheath folds, ‘‘whisps,’’ and boudins.
MSD Interpretation.—This FA is most likely glaciogenic or glacially
inﬂuenced. Evidence for glacial transport of sediment in this FA includes
striations and polish on basement granites where diamictites rest directly
on granite (MBSE-17 and MM-17), the very poor sorting of sediment in
the system, and the presence of angular to rounded grains of all sizes,
faceted and striated clasts, and large boulders composed of local basement
lithologies. In glacial settings, massive diamictites, like the facies described
here, may be the result of subglacial till deposition (Evans et al. 2006),
settling from suspension of a meltwater plume (Visser 1994), settling from
suspension and rain-out from icebergs and iceberg scouring (Dowdeswell
et al. 1994; Lisitzin 2002), mass-transport deposits (Rodrigues et al. 2019),
or debris ﬂows (Powell and Molnia 1989). Glacial depositional
environments that include these processes are subglacial, proglacial
proximal (but outside the inﬂuence of the grounding zone), and
grounding-zone-environments including ground-zone wedges (Batchelor
and Dowdeswell 2015; Demet et al. 2019; Dietrich and Hoffmann 2019),
morainal banks (Eidam et al. 2020), and ice-contact fans (Powell 1990;
FIG. 4.—Photographs of Mt. Butters outcrops showing the stratigraphy of the
Pagoda Fm and postglacial Mackellar Fm, and well as the relief of the Maya
Erosional Surface, in this area. Formations are labelled, and the Maya Erosional
Surface is marked by a green, dashed line. A) A photograph of section MB-17 taken
from a helicopter; view to the south. The saddle in the background of this photo is
approximately halfway between section MB-17 and MBSE-17. B) An outcrop that is
part of the Mt. Butters Massif, and occurs on a spur located southeast of section MB-
17. View is toward the southeast from MB-17. This site has never been visited on
foot, so the scale is not certain. Note the dip of the granitic basement towards the
southwest, which is consistent with basement dip measurements at site MB-17 and
MBSE-17 in this study. This exposure is the same as Figure 4 in (Isbell et al. 2001).
C) Photograph of section MBSE-17, view toward the SW. Purple oval shows
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Benn 1996). The most likely depositional processes and environments for
the MSD facies can be inferred by considering the unique features of this
The subglacial deposition of this FA is not likely to have been the
dominant process, but cannot be wholly ruled out. A tillite interpretation
for the MSD FA is supported by the massive, poorly sorted nature of the
diamictites. However, there is no strong evidence for glacier grounding (for
example, striated pavements or continuous erosional surfaces) above the
base of this facies. The boulder and cobble beds in this FA do not contain
uniform oriented, bulleted, or striated clasts. This suggests that the boulder
beds are lags; that they formed due to winnowing (Eyles 1988), not
This FA was more likely deposited subaqueously than subaerially. In
most of the sections measured in this study, the upper contact of the
diamictite is conformable with, and in some cases, gradationally transitions
into, various subaqueous facies (HS facies association). At section RSP-18,
TABLE 2.—Summary of paleo-transport measurements from this study (mean is Fisher mean 6a99). Note that due to the proximity of the study
locations to the south pole that the orientation of compass directions are not the same across the basin. Mack, Mackellar Fm; L. Pag, lower Pagoda
lacustrine facies (Isbell et al. 2001); T, trend; P, plunge; D, dip angle; DD, dip direction; S, strike; *, measured during 1997 ﬁeld season.
Section (Height) Facies/Fm Feature Measurement Orientation N
10 m above basement Mack Asymmetrical ripples Transport T: 3228, 3128, 00283
0 m - Striae on basement Lineation T: 18081
*0 m - Striae on basement Lineation T: 095863.788
*4 m above base of Mackellar Mack Asymmetrical ripples Transport T: 1578628.587
13–17.5 m Mack Highly deformed slump features Vergence T: 1098, 1048, 11483
*4–48 m of Mackellar Fm Mack Asymmetrical ripples Transport T: 1098620.0810
0 m - Dip and dip direction of basement Mean pole to plane T: 26285
Plane from mean S: 3528
pole D: 118
MB-17A/B L. Pag Symmetrical ripple crest axes Mean lineation T: 346866.0 817
MB-17B L. Pag Slickenside lineation Mean T: 006 81
MB-17C (12 m) MSD Fold axes and small thrust faults Mean Vergence T: 22086 3187
MB-17C (20 m) MSD Plane of thrust faults Vergence T: 2568, 25182
MB-17C (58 m) MSD Slide surface Planes DD: 2438, 19182
Plane from mean S: 1178
pole D: 238
DD: 206 8
MB-17C/D HA Asymmetrical ripples Mean transport T: 325830
(73–81m) Spread 2368- 0558
MB-17C/D HA Cross beds Transport T: 3468, 3568, 35685
MB-17D HA Climbing Ripples Transport T: 2218, 25682
MB-17C (84 m) HA Grooves Direction of T: 2538624 85
(iceberg keel marks?) shallowing
(18 m above Pagoda) Mack Asymmetrical ripples Mean Transport T: 2198639 87
0 m - Striae on Basement Lineation T: 3118, 31682
MBSE-17 (29 m) MSD Sheath fold hinge Orientation T: 0168,P:2081
Vergence 286 8??
MBSE-17 (31–32 m) MSD Thrust faults Mean pole to plane T: 26387
Plane from mean S: 3538
pole D: 218
Vergence: T: 2638
MBSE-17 (54 m) HA Crenulations on slide sufrace Lineation T: 2418, 25182
MBSE-17 (61–68 m) CBS Cross beds and asymm. ripples Mean transport T: 20886 29 87
MBSE-17 (21–26 m above Pagoda) Mack Asymmetrical ripples Mean transport T: 29886 6823
24–48 m LS Cross beds, asymm. ripples, and PCL Mean transport T: 17686 29 818
68–70 m Mack Asymmetrical ripples Transport T: 2618, 2418, 21184
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TABLE 3.—Description of facies and facies associations.
Name Thickness Lithology Structures Formative Process
– 5–75 m
Clast-poor to clast-rich,
sandy diamictite; muddy
sand matrix; minor
discontinuous sand and
Diamictites are massive to
crudely bedded and
ungraded, sands are
generally massive, and
sometime laminated, but
always highly deformed
processes; likely a
combination of mass
subglacial till deposition
–~17 m Coarse to fine-grained,
Fine- to medium-grained sands
that occur in thin, planar
beds with primary current
sandstones are trough cross-
bedded; fine- to very-fine-
grain sandstones are
laminated, or thin-bedded
with unidirectional ripples
and/or a transitional
flow (Mulder and
Distal or medial
portion of an ice-
contact fan or delta
1–10 m Clast-rich, sandy
diamictite; matrix is
3–7 cm beds that are massive
with sharp, planar, and
contacts; soft sediment
deformation associated with
processes; likely a
combination of iceberg-
rain-out and plume
0–15 m Conglomerates to very
Bedding is usually massive,
soft-sediment deformation is
pervasive; few primary
structures include fold
noses, boudinage, faulting,
shear structures above and
below contacts, and ruck
driven processes in the
form of slides, slumps,
0–15 m Coarse- to very f ine-
Medium- to coarse-grained
sandstones are thickly
laminated to bedded with
planar cross beds, trough
cross beds, climbing ripples,
3D ripples that are
asymmetrical or climbing,
hummocky and swaly cross
symmetrical ripples with
bundled upbuilding; very-
fine and fine sandstones are
laminated or thin-bedded
ripples, some flaser ripples
and climbing ripples;
occasional, small-scale soft-
sediment deformation occurs
in all lithologies
transport and deposition,
with some slumping;
unconfined flow; poorly
sorted sediment source;
large variations in
10–30 m Well-sorted, medium- to
quartz arenite sandstone;
rare pebbles, cobbles,
Low-angle and trough
crossbeds; crossbed sets
range in thicknesses from
~15 cm to ~1.5 m; rare
thin beds with asymmetrical
ripples; occur within
amalgamated channels in
sand sheet ,2 km wide
Strong tractive flow
confined to a series of
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FIG. 5.—Photographs of the Massive Sandy Diamictite (MSD) facies association. Rulers are 50 cm long when folded in half and 1 m long when unfolded. Marks on rulers
are in cm. The weathered color of the diamictite is gray to beige and tends to be redder when the matrix has a higher percentage of sand. Small sand bodies are highlighted by
blue where present, and orange line indicates important bedding planes. A) A characteristic clast-poor section of the MSD facies association with a sand ‘‘whisp’’ (MB-17C).
B) A diamictite section of this facies that is very clast poor, yet contains several large boulders (MB-17C). C) A clast-rich section of the MSD diamictite (MB-17C). D) A ruck
structure in the diamictite made by a boulder (MM-17). E) Examples of small sand bodies in the MSD diamictite that experienced simple shear strain, resulting in sheath fold,
‘‘stringers,’’ boudinage, and en echelon structures. (MBSE-17). F) Example of small sand-body bands in an otherwise massive, clast-poor diamictite. (MB-17C). G) A boulder
and cobble bed in an otherwise massive diamictite (MBSE-17). H) Diamictite beds onlapping on to a large, 4 m boulder in the diamictite. Inset picture shows the whole, ~4-
m-diameter boulder in outcrop, with person for scale.
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the massive diamictite is interstratiﬁed with turbidites, a subaqueous
process (see LS facies interpretation). Additionally, there are no
characteristics in this FA that suggest subaerial exposure (e.g., paleosols,
desiccation features, wind-transported sediments).
The lack of clear or laterally continuous stratiﬁcation, as well as a wide
range of grain sizes in the matrix throughout this diamictite, indicate that
the deposition of these sediments was not principally controlled by
‘‘sorting’’ processes such as currents and/or low-density ﬂows. The
massive nature of beds may be a primary depositional feature, as is the
case in subglacial deposits, supraglacial debris, or subaqueous plume
sedimentation. Alternatively, this facies may also be the result of
secondary processes, such as redeposition by mass-transport and high-
density gravity-driven processes–ﬂows (Wright and Anderson 1982;
Vesely et al. 2018), or homogenization due to iceberg scour (Dowdeswell
et al. 1994), dewatering (Collinson and Mountney 2019), or bioturbation
(Svendsen and Mangerud 1997; Murray et al. 2013). Where crude
stratiﬁcation does occur, there are indicators that settling from suspension
may have been a key depositional process. Ruck structures beneath large
clasts suggest those clasts were dropped into the surrounding diamictite
from ice-rafted debris (Fig. 5D) (Thomas and Connell 1985). The
thinning of diamictite beds over large clasts suggests that there was some
component of settling from suspension in their depositional process (Fig.
5H). In glaciomarine settings, this most often occurs as meltwater plume
Small, deformed, sorted sediment bodies occur in many depositional
settings alongside diamictites, including in till (Kessler et al. 2012) and
proximal glaciomarine sediments (Domack 1983; Sheppard et al. 2000).
Depending on their depositional context, small sand bodies are proposed to
be sourced from iceberg dumps, winnowing due to dewatering, or from
incorporating subglacial sediment into till through freeze-on. The small
sand and gravel bodies all indicate pervasive simple shear (whisps,
stringers, boudinage, and sheath folds) or loading (diapir structures). In at
least one case, shearing was associated with an overlying, thicker
sandstone body (Fig. 5F). If the diamictite experienced similar strain
conditions, representative structures would likely be impossible to observe
in outcrop, due to the homogeneous natures of the diamictite. The large
sand bodies are most likely mass-transport deposits that underwent
slumping or non-turbulent ﬂow (Posamentier and Martinsen 2011;
Rodrigues et al. 2019).
The combination of processes (subglacial till deposition, iceberg
rain-out, plume sedimentation, iceberg scouring, and mass transport)
that potentially contributed to the deposition of the massive diamictite
and related facies are most likely to occur in a glacier-proximal to
glacier-intermediate (not distal or proximal) setting, with water depths
below storm wave base (i.e., offshore–transition to offshore) (Licht et
al. 1999; Powell and Cooper 2002). In the Cenozoic, comparable
depositional models have been proposed for similar successions in the
Yakataga Fm, Alaska (Eyles and Lagoe 1990), Weddell Sea, Ross Sea,
and George V regions of Antarctica (Anderson et al. 1980; McKay et al.
2009), as well as St. George’s Bay, Newfoundland (Sheppard et al.
2000). Of these examples, St. George’s Bay is likely the most analogous
to the Shackleton Glacier region during the Permian, since it is not an
open-shelf setting, but an embayment whose topography is controlled
by much older basement rocks (Batterson and Sheppard 2000; Shaw
Plumes emitted from subglacial and englacial meltwater jets were
likely the primary sources of sediments in this system. Plume
sedimentation is not likely to occur where the glacier meltwater is
denser than the ambient water in the depositional environment, which is a
condition associated with lacustrine conditions and resulting hyperpycnal
ﬂows. Therefore the deposition of this facies most likely occurred in
marine or estuarine conditions (Powell 1990). Variations in iceberg
calving, ﬂuctuating glacial hydraulic systems, and minor movement of
the ice margin may explain the variation of matrix grain size and clast
abundance throughout the facies. Glacial hydraulic systems and icebergs
capable of producing sufﬁcient sediment to create this FA are
characteristic of temperate to ‘‘mild’’ subpolar glaciers (Matsch and
Ojakangas 1991; Hambrey and Glasser 2012; Dowdeswell et al. 2016;
Kurjanski et al. 2020).
The time frame in which the deposition of this facies occurred is
difﬁcult to infer, especially without any evidence in the facies for glacier
grounding above its lower contact. Sedimentation rates in glaciomarine
settings are highly variable, even for the same glacier, and strongly depend
on glacier conditions and proximity to the ice front (Hallet et al. 1996).
Rates of accumulation will also depend on physiography of the
depositional area (e.g., fjord vs. open shelf). Accumulating ~100 m of
glaciomarine diamictite could take anywhere from a few years (Cowan and
Powell 1991) to a few millennia (Partin and Sadler 2016; Domack and
Powell 2018). The lack of non-glaciogenic deposits (see LS and HS facies
descriptions for glacial interpretations) interstratiﬁed with the MSD facies
suggests that the deposition of this facies occurred on the shorter end of
this time scale.
Laminated Sands Facies Association (LS)
LS Description.—This is a sandstone FA that occurs only at site RS-18
and is interstratiﬁed with facies MSD. This succession is ~17 m thick and
laterally continuous across the outcrop. Its lower contact is erosional above
MSD, and its upper contact was covered. Internally, this FA consists of
ﬁning-upward packages 3–5 m thick (Fig. 6). This FA consists of coarse-
to ﬁne-grained, well-sorted sandstones. The dominant lithology in this FA
is ﬁne- to medium-grained sands that occur in thin, planar-laminated beds
with primary current lineations. Coarse-grained sandstones at the bases of
some packages are trough cross-bedded. Fine- to very-ﬁne grained
sandstones are laminated, or thin-bedded with unidirectional cross-laminae
and/or ripples. The sandstone is quartz-rich, and grains are subangular to
rounded. Pebbles up to 8 cm in diameter occur at the bases of some cross-
beds. The uppermost parts of ﬁning-upward packages sometimes include
ﬁne- to very-ﬁne-grained black-colored sandstone.
LS Interpretation.—This FA is most likely the result of a series of
noncohesive density-ﬂow events, in the form of high-density turbidites
and/or a transitional concentrated density ﬂows (Mulder and Alexander
2001). Fining sequences such as these could also be formed in a ﬂuvial or
shallow marine setting, where bedload-dominated currents are common.
However, a ﬂuvial setting is inconsistent with the lack of channelization or
indication of surface exposure in the RSP-18 succession. Similarly, a
shoreface setting is unlikely because there is no evidence of emergence or
wave action in the succession. The current directions measured in this unit
are unidirectional, which is also not indicative of a shoreface. Since these
turbidites are interstratiﬁed with facies MSD, they are most likely the distal
or medial part of an ice-contact fan or delta (Lønne 1995; Dowdeswell et
Heterogenous Sandy Facies Association (HS) HS Stratigraphy.—
This FA (Fig. 7) occurs at both Mt. Butters sites above a gradational
contact with the massive diamictite (MSD), and below a sharp, erosional
contact with the cross-bedded sands facies (CBS) at MBSE-17 and a sharp,
planar contact with the Mackellar Fm at MB-17 (Fig. 8). The lower part of
the facies association begins as interbedded, discontinuous bodies of
deformed, sorted sands and gravels (facies HS2) in stratiﬁed diamictites
(facies HS1). Undeformed, moderately sorted, stratiﬁed sandstone bodies
with a range of grain sizes (facies HS3) occur in the middle of the
succession and eventually become the dominant facies near the top of the
succession. This FA ranges in thickness from 1 to 15 m. Lithologies in this
facies are interstratiﬁed.
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Facies HS1: Stratiﬁed diamictite
Description.—This facies is a clast-rich, sandy diamictite (Figs. 7D, G,
I, 8). This diamictite facies is similar to the massive diamictite facies
(MSD), but is consistently stratiﬁed, is more clast-rich, and the matrix is
better sorted. The matrix in this facies is moderately to well sorted. Matrix
grain sizes range from medium to very ﬁne sand. The mean matrix grain
size varies between beds. Most beds have a mean matrix grain size of
medium sand, but some beds have a dominantly very ﬁne grained sand
matrix. Clasts in this facies are angular to subrounded, and have a size
range similar to MSD, granule to 1 m boulders. The beds in this facies are
3–7 cm thick, are planar, laterally discontinuous, and have sharp contacts.
The distribution of clasts is random and unrelated to bedding planes.
Larger clasts often punctuate bedding planes. This facies has a gradational
lower contact overlying the massive diamictite (MSD) facies. Vertical and
lateral contacts with other facies in this FA (HS3 and HS2) are most often
sharp, sometimes loaded, deformed, or erosional. This bedded diamictite is
frequently interbedded with facies HS2. Strata in this facies adjacent to
contacts with facies HS2 often display soft-sediment deformation,
including load structures and sheared contacts.
Interpretation.—Similar to facies MSD, this facies was likely
deposited in a glacier-proximal glaciomarine setting, but was dominated
by plume sedimentation and iceberg rainout (Eyles 1987; Licht et al. 1999;
Powell and Domack 2002; McKay et al. 2009). The consistency of the
stratiﬁcation in this facies compared to the MSD facies suggests that the
sediment in this facies was not as frequently subject to remobilization,
either through gravity-driven transport or iceberg scour. The better sorting
of the matrix in this facies relative to the massive diamictite facies indicate
that sediment-sorting processes were more active than during the
deposition of MSD. In a glacial-proximal setting, this likely means that
turbulence kept ﬁne-grained sediment suspended in the water column and
did not allow it to settle out. The clast-rich composition and the thin nature
of the beds also suggest that depositional processes were relatively
constant, compared to the high variability of clast contents in facies MSD.
Loaded, deformed, and erosional contacts in this facies, and with other
facies in this FA, indicate that this facies was rheologically ‘‘ weak,’’ or
experienced ductile deformation at low strain magnitudes. Therefore,
during and shortly following deposition, facies HS1 was likely water-
saturated. This also suggests that that sedimentation was rapid, that this
facies was subject to gravity-driven processes, and that these sediments
were deformed by gravity-driven deposits (Figs. 7D, F, 8) (Visser 1994).
Facies HS2: Chaotic Sandstones
Description.—This facies consists of very ﬁne- to coarse-grained
sandstones (Fig. 7A–C). Beds in this facies are laterally discontinuous
(Figs. 7G, 8). The thicknesses of sandstone bodies are laterally inconsistent
and range in thickness from 0.5 m to 2 m. Widths of sandstone bodies
range from ~1 m to outcrop scale. Sandstone beds may be interbedded
with one another, but are dominantly interbedded with the surrounding
either massive (MSD FA) of stratiﬁ ed (HS1) diamictite facies. Sandstone
bodies are irregularly shaped, but generally have planar to lenticular
shapes. Lower and lateral contacts are deformed, sharp, or erosional, and
often show evidence of soft-sediment deformation (Fig. 7D, F). Upper
contacts are sharp and conformable. Contacts between sandstone bodies
are erosional or deformed.
Beds in this facies are often internally massive, and soft-sediment
deformation is pervasive (Fig. 7E, F). Primary sedimentary structures are
sometimes preserved, but this is rare and only occurs in a small area of any
given bed. Secondary structures in this facies include fold noses (Fig. 7E),
boudins, faults, and other simple shear structures above and below contacts
(Fig. 7D–G), and ruck structures associated with rare outsized clasts (Fig.
7I). Grain size in this facies ranges from conglomerate to very-ﬁ ne-grained
sandstone. Fine- to medium-grained sandstones tend to be well sorted,
while coarse-grained sandstones and conglomerates are poorly sorted. The
medium- and coarse-grained sandstones occur more frequently than ﬁner-
Interpretation.—The sandstone bodies in this facies are most likely the
result of mass-transport, gravity-driven processes (Posamentier and
Martinsen 2011; Sobiesiak et al. 2018; Rodrigues et al. 2019). Preserved
primary sediment structures in some of these bodies indicate that sediment
sorting due to current transport likely occurred before the remobilization
and ﬁnal deposition of these sediments. Irregular lateral contacts between
this facies and the two diamictites facies (HS1 and MSD) indicate that
mass-transported bodies were sandstone rafts deposited into pre-existing
diamictites by gravity driven processes. The deposition of mass-transport
FIG. 6.—Photograph of the Laminated Sands
(LS) facies association at site RS-18 (Reid Spur).
Triangles show ﬁning-upward packages separated
by erosional surfaces. Reddish lithologies are
medium-grained sandstone to coarse-grained
sandstone, and have planar lamination and low-
angle cross-stratiﬁcation. Black lithologies are
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deposits (MTDs) into the diamictite indicates that diamictite deposition
was contemporaneous with MTD emplacement. This also suggests that the
diamictite was weak (highly susceptible to deformation), likely due to both
high pore-water pressures and a lack of consolidation.
Facies HS3: Stratiﬁed Sandstones
Description.—This facies consists of these facies very ﬁne- to coarse-
grained sandstones (Fig. 7A–C). Medium- to coarse-grained sandstones
are thickly laminated to bedded. Common sedimentary structures in
medium- to coarse-grained sandstones include planar cross beds, trough
cross beds, climbing ripples, and 3D ripples that are asymmetric or
climbing (Fig. 7A). Rare sedimentary structures include hummocky and
swaly cross stratiﬁcation and symmetrical ripples with bundled upbuilding
(Fig. 7B, C). Lateral and vertical variations in sedimentary structures
within a unit of the same lithology or grain-size is common. Coarser sands
are occasionally massive or contain laterally discontinuous sand or gravel
lenses. Trough cross-beds occasionally have pebbles at the bases of
FIG. 7.—Photographs of the Heterogenous Sandy (HS) facies association at Mt. Butters in section MB-17 and MBSE-17. Marks on all rulers are in cm. Rulers are50cm
when folded in half and 1 m long when unfolded. White dashed lines have been used to highlight contacts and important bedding surfaces. A) Medium-grained sandstone
layer in facies HS3. The sandstone body is composed mostly of climbing dunes (cross beds) and capped by asymmetrical ripples. Note sharp contact with black-colored,
bedded diamictite (facies HS1) in lower part of image (MB-17). B) Fine- to medium-grained sandstone in facies HS3 with up-building symmetrical ripples (MB-17). C)
Swaly and hummocky cross-stratiﬁcation in sandstone layer of facies HS3 (MB-17). D) Loaded, possibly boudinage, contact between a massive sandstone (HS2) and bedded
diamictite (HS3) (MB-17). E) Sandstone with a soft-sediment recumbent fold nose in facies HS2, likely at the front of a slump (MB-17). F) Deformed contact between facies
HS2 and HS3. Jacob’s staff for scale; marking is every 10 cm. (MB-17). G) Outcrop showing interﬁ ngering between facies in this facies association. Black-colored sediments
are bedded diamictites (facies HS1), other lithologies show sand and gravel of facies HS2. (MB-17). H) Groove structures on top of facies HS1. View approximately toward
north (down the Shackleton Glacier). I) Granitic outsized clast punctuating interlaminated sandstone beds in the bedded diamictite facies HS1 (MB-17).
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FIG. 8.—Photo mosaics (A, B, and C) and interpretive sketches (B0and C0) of the upper part of section MBSE-17 at Mt. Butters. Mosaics A and B were taken from a
helicopter, and mosaic C from the ground. This is the opposite side of the outcrop shown in Figure 4C. View is to the north, and the ridge runs roughly east–west. Part A
shows lateral variations in the architecture of the cross-bedded sandstone (CBS) facies. Figures B and B0highlight the stratigraphic relationships between the Massive
Diamictite (MSD) facies association, the Heterogenous Sandy (HS) facies association, and the CBS facies association. Black lines in CBS denote channel erosional surfaces.
Black lines in HS3 indicate soft-sediment deformation. The part of this outcrop highlighted by C and C 0is located within the red box on Part B. Part C shows a part of the
outcrop that is characteristic of the HS facies association, and was selected to illustrate the pervasive nature of soft-sediment deformation in this facies association.
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troughs. Very ﬁ ne- and ﬁne-grained sandstones are laminated or thin-
bedded. Sedimentary structures in ﬁne-grained and very ﬁne-grained
sandstones include unidirectional cross-laminae, ripples, planar and wavy
laminae, some ﬂaser-bedded cross-laminated units, climbing-ripple
laminae, and rare outsized clasts with ruck structures. Lithologies of all
grain sizes also have minor amounts of soft-sediment deformation,
including dewatering structures, minor folds, and loading.
At the upper contact of the Heterogenous Sandy FA with the Mackellar
Fm in section MB-17, these sandstones had at least ﬁve large, shallow,
east–west-oriented grooves (Table 2) that occur in a massive, well-sorted
sandstone (Fig. 7H). All grooves were ~1–2 m wide and 10 cm deep.
Berms ~10–20 cm high bound the grooves on their long sides. All of the
grooves gradually shallow towards the west (2158), and two of the grooves
had prow-like berms at their eastern terminus.
Strata in this facies are laterally continuous across the outcrop.
Sandstone bodies are generally wedge-shaped and thicken in the direction
of ﬂow. Erosional surfaces are common in the facies. This facies has an
erosional lower contact with both diamictite facies (MSD and HS1), and a
sharp, conformable contact with the overlying UFG facies association.
Interpretation.—This facies was likely deposited rapidly in an
unconﬁned, subaqueous setting with limited reworking by waves and
icebergs. The sedimentary structures in this facies indicate that current-
dominated transport and deposition occurred, followed by syndepositional
or postdepositional slumping of some deposits. The wide range of grain
sizes and sedimentary structures indicate a sediment source with a wide
range of grain sizes and huge variations in current velocities during
deposition. Common sedimentary structures, such as planar cross beds,
trough cross beds, and asymmetrical ripples, suggest unidirectional,
relatively high-velocity currents. Climbing ripples suggest decrease in ﬂow
velocity downcurrent, which is characteristic of unconﬁned ﬂows, which is
supported by the lack of channelized deposits. Flaser bedding, as well as
abrupt changes in grain size and sedimentary structures within and
between beds, indicates that current velocities were highly variable and
ﬂuctuated. Rare soft-sediment deformation suggests high pore-water
pressure during rapid deposition. Rare hummocky and swaly cross-
stratiﬁcation and symmetrical ripples form under oscillatory ﬂow
conditions created by surface waves, possibly during reworking by storms
(Reineck and Singh 1980; Dumas and Arnott 2006; Collinson and
Mountney 2019). Their rare occurrence and interstratiﬁcation with ﬁne-
grained sediments suggest that these wave features formed below normal
wave base. The shape of the grooves on the upper contact of this FA at site
MB-17, their position in the upper contact of a massive (homogenized)
sandstone, and immediately below the contact with the dropstone-bearing
lower Mackellar Fm suggest that these features are iceberg keel marks
(Dowdeswell et al. 1994; Vesely and Assine 2014). These massive
sandstones were likely deposited in a similar way to other sandstones in
this facies, but were homogenized by iceberg actions.
HS Depositional Environment.—The three facies in this FA represent
a complex depositional environment that is characteristic of subaqueous,
glacier-intermediate to -proximal settings in front of the terminus of
temperate to ‘‘mild’’ subpolar glaciers. Evidence for the glaciogenic origin
of this FA includes pebbles with ruck structures (representing ice-rafted
debris), iceberg keel marks (facies HS3), the very poor sorting of sediment
in the system, and the wide range of sedimentary grain shapes in the
Pagoda Fm sandstones, which are described at the beginning of the facies
The stratiﬁed diamictite (facies HS1) was deposited primarily through
plume sedimentation in a glacier-proximal setting, and is the dominant, or
‘‘background,’’ sedimentation type in this FA. The gradational contact
separating MSD (massive diamictite) and facies HS1 suggests a gradual
shift in depositional environments between the two. Sediment composition
is consistent between the two diamictite facies, suggesting that the
sediment source did not change, but that the depositional environment
shifted from glacier-intermediate to glacier-proximal. The deposition of
both diamictites was likely controlled by the same processes (i.e., plume
sedimentation, iceberg rain-out, iceberg scouring, and mass transport) but
to different degrees. This shift from glacier-intermediate to glacier-
proximal was most likely driven by a minor readvance of the glacier
margin, but it may have also been an apparent effect caused by the
progradation of the overlying grounding-line fan system (HS2 and HS3).
The sandstone facies in this FA (facies HS2 and HS3) most likely
represent the medial part of a subaqueous grounding-line fan(s) system
(Powell 1990; Lønne 1995; Dowdeswell et al. 2015). In facies HS3, the
high-velocity, unidirectional current transport combined with abrupt
changes in grain size (i.e. current velocity), unconﬁned ﬂow, and
interstratiﬁcation with the bedded diamictite (facies HS1) are characteristic
of grounding-line fans (Powell 1991). The gravity-driven transport of
facies HS2 sandstone bodies were likely derived from deposits similar to
(or the same as) facies HS3. ‘‘Shedding’’ of sediments is characteristic of
the rapid sedimentation in grounding-line fan systems (Benn 1996; Powell
and Alley 1997; Lønne et al. 2001). Intense, ductile deformation and
loading along contacts throughout this FA indicate that all facies were
water-saturated, unconsolidated, and generally had the consistency of soup,
suggesting rapid deposition and that they were therefore prone to
resedimentation (Fig. 8B).
The wave reworking of some sandstone beds, as indicated by
hummocky and swaly cross-stratiﬁcation and wave-ripple stratiﬁcation
(symmetrical ripples) in facies HS3 suggest that this depositional
environment was occasionally subjected to surface-wave activity (below
normal wave base). These features indicate that there was not perennial ice
cover during the deposition of this FA. This evidence for wave reworking
suggests a similar, wave-winnowing origin for the boulder beds in facies
Where this FA is well developed in outcrops at Mt. Butters (sites MB-17
and MBSE-17), the succession has a general coarsening and increase in
sorting trends upward. This trend indicates the increase in the proximity of
the energy and sediment source, either through progradation of the
grounding-line system and/or advance of the glacial front.
Cross Bedded Sandstone Facies (CBS)
CBS Description.—This facies occurs at Mt. Butters section MBSE-17,
and consists of an erosionally-based, laterally extensive, channel-form
sandstone body 10–30 m thick and several hundred meters wide that cuts
into and through a laterally continuous thick sandstone sheet at the top of a
coarsening-upward succession of the HS (1–3) facies association (Figs. 8,
9). The sandstone body is laterally continuous across outcrop MBSE-17
but is not present at section MB-17, which is ~2 km north (Fig. 1B). The
basal CBS erosional surface has a relief of up to 10 m, and lower contacts
with all HS facies and the MSD facies (Fig. 10). The upper contact of this
facies with the overlying Mackellar Fm is sharp and horizontal.
This facies occurs in multistoried, multilateral sand-ﬁlled channel-form
bodies displaying nonsequential, lateral compensational stacking patterns
(Fig. 9). Individual channels are meter-scale thick and tens of meters wide,
trough-shaped in cross section, and are ﬁlled by either vertical or
downstream accretion dipping to the east. Channels are truncated by the
bases of overlying channel bodies. Channel stacking is nonsequential and
disorganized, with some aggradation. This facies is composed of well-
sorted, medium- to very coarse-grained quartz sandstone, with minor
occurrences of conglomerate lenses and beds (Fig. 9). Mudrocks were not
observed in the sandstone bodies. Very rare pebble- and small-cobble-size
clasts occur throughout the sandstones. Those clasts have lithologies
similar to clasts observed in both diamictite facies (HS1 and MSD).
Sedimentary structures are almost exclusively 0.15–1.5-m-thick sets of
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low-angle stratiﬁcation and trough cross-beds (Fig. 9A). Thin beds with
asymmetrical ripples also occur but are rare.
Adjacent to the described section (MBSE-17), the edge of the channel-
form body appears to extend across the top of the underlying strata as a
wing-like extension (Fig. 8A). Channel-form sandstone bodies occur in the
wing. These bodies appear to transition laterally into other thick sandstones
with channelized bases. Most notably, the contact between the channelized
sandstones in the wing and the underlying HS facies appears to be sharp
CBS Interpretation.—The CBS facies in the Pagoda Fm was deposited
by strong, tractive, conﬁned ﬂow as indicated by the occurrence of the
basal erosion surface and the internal channel bodies ﬁlled by downstream-
accreting bar forms and cross-stratiﬁcation. The trough shape of the
internal sandstone-ﬁlled channels and their multistoried and multilateral
characteristics suggest that the channels were stationary during ﬂow in the
channels and that they did not migrate until channel switching occurred
and new channels formed as older channels ﬁlled and were abandoned
(Friend 1983). The occurrence of low-angle stratiﬁcation and meter-scale
trough cross beds organized into downstream-accreting bodies with
massive bedded to lenticular gravels suggest high ﬂow velocities. Such
features are characteristic of highly dynamic systems where aggradation in
channels likely forced channel switching to adjacent areas on the
depositional surface. The occurrence of this facies association on top of
unconﬁned coarsening-upward HS (1–3) facies association and the
occurrence of wings that appear as a continuation of the HS (1–3)
coarsening-upward succession suggest that the CBS facies formed as part
of a HS–CBS larger-scale dispersal system. The presence of wings also
suggests that parts of the CBS system were unconﬁned and represent
‘‘overbank’’ deposition on surfaces in areas between channels. Together,
these patterns are most characteristic of an unconﬁned, distributive setting
(Funk et al. 2012). This unit is similar to some grounding-line fan systems
that authors have called subaqueous outwash fans (Visser et al. 1987;
Thomas and Chiverrell 2006; Rose 2018).
Whether this facies was deposited subaerially or subaqueously is
unclear. However, the CBS sandstone body does not contain evidence for
shallow-water wave reworking, pedogenesis, or subaerial exposure,
whereas facies both below and above this facies represent subaqueous
deposition below normal wave base, and likely below storm wave base.
Therefore, a subaqueous setting seems likely.
Most of the Pagoda Fm in the Shackleton Glacier Area is composed of
massive diamictites (facies MSD) that likely formed in glacier-proximal to
glacier-intermediate environments, at depths largely below normal wave
base, through a variety of glaciogenic and glacially inﬂuenced processes.
These massive diamictites are also conformably overlain by, and
interstratiﬁed with, grounding-line fan deposits (facies associations LS,
HS, and CBS). These glacially derived lithologies are conformably succeed
by prodeltaic, ﬁne-grained facies of the Mackellar Fm.
Evidence for the grounded advance of a glacier(s) in the Shackleton
Glacier region is present at base of the MSD facies at both Mt. Butters sites
(MB-17 and MBSE-17) and at Mt. Munson (MM-17). The lower contact
of the MSD facies is not exposed at Reid Spur, so similar inferences cannot
be made for that locality. No conclusive evidence for subglacial
deformation or erosion was observed higher in the Pagoda Fm at any
site examined, though a subglacial origin for the MSD facies cannot be
wholly ruled out. This advance was likely made by a glacier whose
thickness exceed 100 m and ﬂowed from north to south across the
Shackleton Glacier region. The advance would have come from the
direction of the present Ross Sea and crossing the TAB’s margins
perpendicular to the elongate trend of TAB (See discussion in prior
section; Fig. 11D). This observation, along with other data from the TAB,
strongly suggest that there were at least two ice centers in Antarctica during
the Permian, one located on the East Antarctica craton and one in present
day West Antarctica (Isbell 2010).
When the glacier margin retreated from the Mt. Butters, Mt. Munson,
and Reid Spur sites, the deposition of glacier-proximal deposits was
FIG. 9.—Photographs of the Cross Bedded Sandstone (CBS) facies at section
MBSE-17 on Mt. Butters. A) An example of low-angle and trough cross-bed sets in
this facies. Measuring stick is 1 m long. B) Photograph of facies in outcrop, noting
occurrences of minor lithologies.
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FIG. 10.—Sedimentary logs and paleotransport directions from sites described in this study. Paleotransport directions are plotted to align with the maps of modern
Transantarctic Mountains presented in this study, so that north is toward the bottom of the page and varies by geographic position (see inset map). Details of paleotransport
measurements are available in Table 2. MM-17 is Mt. Munson, MB-17(A–C) is Mt. Butters 1, MBSE-17 is Mt. Butters 2, and RS-18 is Reid Spur. Lithologies are grouped by
their interpreted facies or facies association. Colored bars next to each log are used to indicate the distribution of facies associations in these sections, and colored areas in
between sections show interpretation of each facies extent outside the section. Dark green corresponds to the localized lacustrine facies of the Pagoda Fm at Mt. Butters
described by Isbell et al. (2001). Bright green represents the Massive Sandy Diamictite (MSD) facies association. Different shades of purple represent grouping of the
Heterogenous Sandy facies association (HS1, stratiﬁed diamictite; HS2, chaotic sandstones; HS3, stratiﬁed sandstones). The inset map shows the location of each section and
isopachs of the Pagoda Fm in the Beardmore Sub-basin from Isbell et al. (2008c). Light blue represents the Cross Bedded Sandstone facies (CBS). Dark blue represents the
Laminated Sandstone facies association (LS). Red represents the Mackellar Fm. The datum for these columns were chosen using the last evidence for glaciogenic sediments,
either the uppermost outsized clast or diamictite, which is a marker that also serves as an upper sequence boundary.
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FIG. 11.—Box diagrams showing progressive phases of the depositional model for the Pagoda Fm and lowermost, glacially inﬂuenced Mackellar Fm in the Shackleton Glacier region, alongside maps showing the modern
locations of the sites described in this study with transport orientations related to each part of the depositional model. See Table 3 for ﬂow directions. A) Map of the modern central Transantarctic Mountains. Gray areas are
approximate locations of nunatuk regions. Solid black lines are isopachs of the Pagoda Fm, copied from Isbell et al. (2008c). Red square indicates map area in Parts B–F. Note that the smaller map areas are rotated relative to
the larger ‘‘map A.’’ Yellow stars show site locations described in this study. B) Paleotopography of the Shackleton Glacier Area before deposition of the Pagoda Fm. Map shows strike and dip of granite surface underlying
the Pagoda Fm at Mt. Butters site MB-17, corrected for modern structural conditions. This map also includes the Isbell et al. (2008c) isopach lines and derived basin-axis orientation for reference. C) Proposed depositional
conditions for the lacustrine facies at the base of the Pagoda Fm at site MB-17 (Isbell et al. 2001). The green, double-sided arrow shown on the map shows the orientation of symmetrical (wave) ripple crests, which parallels
the strike of the underlying basement. D) Proposed depositional conditions for Massive Sandy Diamictite (MSD) facies and Laminated Sands (LS) facies during retreat of the glacier out of the Shackleton Glacier area. On
the map, purple wedge indicates range of ﬂow direction in the LS facies, and green wedges indicate down-slope transport direction in the MSD facies. Blue double-headed arrows show glacier ﬂow directions measured in
this study. E) Proposed conditions during the deposition of the grounding-line fan represented by facies associations MSD, Heterogenous Sandy (HS), Cross-bedded Sands (CBS), and Mackellar Fm. The red wedge shows
range of ﬂow directions in the Mackellar Fm, blue and purple show the range of ﬂow direction for the CBS and HS ﬂow directions, respectively, and the blue arrow indicates the mean ﬂow direction of the CBS facies at site
MBSE-17. F) Proposed conditions during the deposition of the lower Mackellar Fm.
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initiated. The massive diamictite (MSD) facies that dominated the Pagoda
Fm at Mt. Butters and Reid Spur (RSP-18) was likely deposited through a
combination of glaciogenic depositional processes characteristic of (sensu
lato) glaciomarine settings, a combination of settling from suspension of
neritic sediments, plume sedimentation from subglacial and englacial jets,
as well as iceberg sedimentation and mixing. The prevalence of plume
sedimentation and lack of hyperpycnal (under-) ﬂow indicate that the water
in the TAB was likely either marine or brackish. It should be noted that the
TAB did not have a shelf, in the formal sense, as it was a trough-shaped
basin that was not directly connected to the open ocean.
The Shackleton Glacier region during the Cisuralian was not an open
shelf, but a near-coast setting with ample topographic relief, and water
depths below normal wave base, such as St. George’s Bay, Newfoundland
(Sheppard et al. 2000). Ultimately, many of these sediments were likely
remobilized by gravity-driven slides, slumps, and ﬂows. This is especially
true at the Mt. Butters site, where the vergence of soft-sediment
deformation features in the MSD facies follow the local paleotopographic
slope (Fig. 11B, E). Soft-sediment deformation caused by mass-transport
deposits (facies HS) in bedded diamictites at Mt. Butters shows that the
diamictites were unconsolidated, water-saturated, and subject to resedi-
mentation, suggesting relatively rapid deposition.
At Mt. Butters and Reid Spur, massive diamictites are interstratiﬁed
with and overlain by grounding-line fan deposits in the form of the LS, HS,
and CBS facies associations. The stacked density ﬂows (facies LS) at Reid
Spur represent the medial to distal part of a fan, likely in relatively deep
water. Flow direction in these sediments are towards the basin axis (Fig.
11D), suggesting that the basin geometry controlled topography in this
area. On the other hand, the fan deposits at Mt. Butters are more proximal
to the glacier margin. The laminated diamictites, reactivation surfaces, and
storm-wave deposits in the fan at the top of the Mt. Butters section (MB-17
and MBSE-17) suggest that this fan was built gradually along a relatively
stable margin and not during a single, catastrophic drainage event
(Dowdeswell et al. 2016). The successions at Mt. Butters from the HS
facies association through the CBS facies association represents either the
progradation of one of these fans, a minor readvance associated with the
deposition of the fan, or a combination of the two (Fig. 11E). The
dispersed ﬂow directions in the HS facies are characteristic of a fan, though
their orientation broadly toward the north is away from the basin axis and
opposed to glacier ﬂow directions. Flow directions in the CBS facies are
well clustered toward the southwest (Table 2, Fig. 11E). Inﬂuence of the
paleotopographic slope at Mt. Butters and the radial nature of fan
geometries are the most likely cause from these seemingly antagonistic
ﬂow directions (Fig. 11E). The orientation of the ice front would have also
likely been inﬂuenced by this topography.
Basin-Margin vs. Basinal Facies Associations
Isbell et al. (2008c) described two generalized facies association of the
Pagoda Fm: one that occurs in basinal settings and another that occurs
along the basin margins. The sites described in this study from Mt.
Munson and Mt. Butters in the Pagoda Fm are characteristic of the Basin
Margin FA, because they are relatively thin successions (,100 m) and
have evidence for subglacial erosion in the form of polished and striated
bedrock (MBSE-17 and MM-17), and subglacially sheared lacustrine
sediments (MB-17; Isbell et al. 2001). The site at Reid Spur (RSP-18) is
also most characteristic of the Basin Margin FA, because diamictite facies
there are thick and unstratiﬁed and has a poorly sorted matrix, suggesting
plume sedimentation and gravity-driven redeposition. However, this
section also likely represents a transition between the two facies
associations. This is indicated by the LS facies at this site, which is
attributable to the mid- to distal part of a grounding-line fan, and by the
inferred thickness of the Pagoda Fm at Reid Spur (~100 m), which is the
same as the transition thickness between Isbell et al. (2008c)’s two FAs.
Effects of Topography
The transport directions in these successions strongly suggest that
paleotopography (relief on the Maya Erosional Surface) played a
signiﬁcant role in the deposition of the Pagoda in the Shackleton Glacier
region. The inﬂuence of topography is particularly evident in the section
MB-17 (Fig. 11). The surface of the basement at MB-17 dips toward the
west at 118(Figs. 4, 11B; Appendix A). Wave-ripple crests in facies BFG
are parallel to the strike of the basement surface at MB-17 (Fig. 11C),
suggesting that the paleotopography created by this surface was sufﬁcient
to affect and orient wave action. The transport directions of slumping and
other gravity-driven processes in facies MSD at both sites MB-17 and
MBSE-17 are also generally towards the west (Figs. 10, 11D), following
the same slope. Flow directions in the grounding-line fan facies
associations (HS and CBS) have a wide spread ranging from the southwest
toward the east. However, in the HS facies association at MBSE-17,
gravity-driven transport is still towards the west (Figs. 10, 11E). The south-
southeast ﬂow directions in the turbidite facies (LS) at site RSP-18 do not
align with the ﬂow direction at Mt. Butters, suggesting that those facies
have a separate origin than the Mt. Butter’s grounding line fan(s) (Figs. 10,
Paleotopographic control on the deposition of Pagoda Fm and Mackellar
Fm has been noted by authors throughout the TAB (Lindsay 1970b; Barrett
1972; Isbell et al. 1997a, 2008c; Cornamusini et al. 2017). In previous
studies, ice-ﬂow directions (usually striae on bedrock or clast pavements)
have often been combined with other transport directions in the Pagoda Fm
to infer a generalized transport direction. However, recent work in modern,
high-relief, glaciated landscapes has shown that the relationship between
glacier ﬂow directions, other transport directions, and topography can be
used to infer the thickness of the glacier relative to the magnitude of relief
on the landscape (Landvik et al. 2014). In other words, whether or not a
glacier ‘‘follows’’ the underlying topography is a function of the glacier’s
thickness. Therefore, indicators of glacier ﬂow should be considered
separately from other transport directions.
The ice-ﬂow directions below the Pagoda Fm in the Shackleton Glacier
region are oriented generally northwest to southeast at Mt. Butters and Mt.
Munson (Fig. 11D). Though none of the striae observed during this study
had unidirectional indicators, previous workers in this area have found
glacially carved features in the basement underlying the Pagoda Fm that
show glacier ﬂow was basin-ward (toward the south) or along the basin
axis (toward the southeast) (Appendix A). These uniform ﬂow directions
on both a paleotopographic high (MM-17) and paleotopographic low (MB-
17 and MBSE-17) suggest that the glacier, when it created these striae, was
sufﬁciently thick to ‘‘ overtop’’ the pre-existing topography in the
Shackleton Glacier Area. Based on the difference in Pagoda Fm thickness,
the local relief between site MB-17 (Mt. Butters) and MM-17 (Mt.
Munson) was at least 85 m, and the onlapping of the Mackellar Fm onto
basement in this area suggests that localized relief may have exceeded 100
m (Seegers 1996; Isbell et al. 1997a; Seegers-Szablewski and Isbell 1998).
This scale of relief is on the scale of large hills. The topographic
prominence of subglacial features on the scale of 100 m is considered
negligible in studies of modern ice-sheet margins (e.g., Lindb ¨ack and
Pettersson 2015; Cooper et al. 2019), but would likely perturb or redirect
the ﬂow of relatively thin glaciers.
This discussion is all to say that the thickness of the glacier that created
these striae more likely than not greatly exceeded the thickness of local
topographic relief (~100 m), and that ﬂow was most likely toward the
center of the TAB. This inference suggests that the glacier was more likely
an ice cap or ice sheet than an alpine glacier.
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Glacial Systems Tracts
In this study, the sequence stratigraphy of the Pagoda Fm in the
Shackleton Glacier region can be considered only at Mt. Butters, because
that is the only location where the authors were able to measure complete
sections. Since the Mt. Butters locations are basin-marginal successions
(Fig. 11), this analysis of glacial sequence stratigraphy should not be
considered applicable to basinal successions of the Pagoda Fm. The
Pagoda Fm at Mt. Butters is unique in the TAB because it contains only a
single glacial sequence as deﬁned by Powell and Cooper (2002) and
Rosenblume and Powell (2019) (Fig. 12). The sequence described in this
paper is bounded at its base by a surface of glacial erosion (deﬁned by
striae on bedrock and the deformation of underlying lacustrine sediments
(see Isbell et al. 2001) and at its top by an iceberg termination surface
(deﬁned by the ﬁnal outsized clast in the section). Since the Pagoda Fm
in this location is overlain by the nonglacial Mackellar Fm, there is no
true maximum retreat surface beyond the iceberg termination surface
(i.e., the last dropstone in the lower Mackellar Fm). Most of the
succession likely represents a glacial-retreat systems tract, though there is
likely some fraction of the massive diamictite facies above the erosional
surface that is more likely to have been subglacially deposited and would
therefore represent a glacial-maximum systems tract. The transition
FIG. 12.—Glacial sequence stratigrapahy of the
Mt. Butters sections, after Powell and Cooper
(2002) and Rosenblume and Powell (2019). The
depositional systems are deﬁned as N, nonglacial;
D, glacier distal; P, glacier proximal; and I, ice
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between those two systems tracts would be deﬁned by a grounding-line
This sequence is most consistent with Rosenblume and Powell (2019)’s
Type I ‘‘idealized glacial sequence,’’ which is a model developed to reﬂect
a sedimentation sequence deposited during the retreat of a relatively warm
subpolar glacier with a glacial erosion surface as a lower sequence
boundary and sufﬁcient meltwater for the development of grounding-line
fans. This sequence model was developed based on upper Miocene
sediments from the Ross Sea region of Antarctica (Rosenblume and Powell
2019), whose climatic and geology were likely reasonably similar to the
TAB during the Permian. The Type I idealized sequence is interpreted to
represent a dynamic climatic glacial system with very high erosion rates
and debris ﬂuxes, which is consistent with the depositional model
presented for the Pagoda Fm in this paper.
This study ﬁnds that most of sediments in the Pagoda Fm in the
Shackleton Glacier region were deposited during the retreat of a temperate
to ‘‘mild’’ subpolar glacier. The key indicators for this retreat include the
presence of grounding-line fan systems, ample evidence for plume
sedimentation and rapid deposition, as well as abundant glacially
transported clasts with a wide size range made of local basement
lithologies. This glacier had a subaqueous terminus during the deposition
of the Pagoda Fm in the Shackleton Glacier area, either in a marine or a
One of the ultimate goals of the study of LPIA glaciogenic sediments is
to infer the ‘‘type’’ and distribution of glaciation experienced in any given
basin. Glacier ‘‘type’’ typically refers to the glacier’s thermal regime and its
size (i.e., ice sheet, ice cap, or ice ﬁeld). Such characteristics of glaciers are
controlled by many factors but generally tie back into climate and geologic
setting. Glaciogenic sedimentary deposits are often used to infer the
thermal regime of their parent glacier (Dowdeswell et al. 2016; Kurjanski
et al. 2020). Recent studies all agree that glaciogenic deposits in the TAB
are most likely the result of transport and deposition by temperate or
‘‘mild’’ subpolar glaciers, largely because grounding-line fans are common
in the Pagoda Fm and its equivalents, which suggest that the TAB glaciers
released an abundance of meltwater (Isbell et al. 2008c; Isbell 2010; Koch
and Isbell 2013; Cornamusini et al. 2017).
The presence of subaqueous fan deposits interspersed in the massive
diamictite facies of the Pagoda Fm in the Shackleton Glacier area suggests
that the glacier had an active, persistent, and organized subglacial
hydrologic system, and that during its retreat the glacial margin was
stationary for at least a few years at a time to create grounding-line fans
(Cowan and Powell 1991; Hunter et al. 1996; Dowdeswell et al. 2015,
2016). These inferences are also supported by the succession’s stratigraphy,
which is characteristic of a ‘‘ Type I’’ (‘‘mild’’ subpolar glacier with
grounding-line-fan development) glacial systems tract (Rosenblume and
Powell 2019). Whether the glacier responsible for the deposition of the
Pagoda Fm in the Shackleton Glacier Region had more of a ‘‘ mild
subpolar’’ thermal regime, similar to modern glaciers in eastern Svalbard,
or more of a truly ‘‘temperate’’ thermal regime, like modern glaciers of
southern Alaska, is difﬁcult to discern. To clarify likely thermal regime,
additional glaciogenic successions and nonglacial paleoclimate indicators
in the TAB should be examined. In either case, glaciers with mild thermal
regimes and developed subglacial hydrological systems are far and away
the more proliﬁc producers and transporters of sediments, and that the
sedimentary record is therefore biased towards them. The presence of
deposits from temperate or mild subpolar glaciers does not mean the
thermal regime of the glaciers were never cold-based or that there was not
lateral variation in glacier thermal regime.
The unidirectional orientation of subglacial ﬂow indicators, along with
evidence to topographic relief exceeding ~100 m in the Shackleton
Glacier region, suggest that glacier thickness also exceeded 100 m during
its maximum. Such observations do not allow an inference of maximum
possible ice thickness. However, we can infer that the glacier was more
likely part of an ice sheet or ice cap than an ice ﬁeld or alpine glacier. The
glacier also may have thinned substantially during its retreat and
subsequent deposition of the Pagoda Fm in this area.
Throughout the TAB, Asselian–Sakmarian glaciogenic depositional
environments were locally and regionally variable in part due to the
inherent complexity of glacial processes and preservation potential, but
also due to the topography of the pre-existing landscape. Most studies of
glaciogenic rocks in the TAB note up to several hundred meters of
topographic relief on the underlying basement that is not wholly ‘‘ﬁlled’’
by glacial sediments and continued to inﬂuence postglacial sediment
deposition (Isbell et al. 1997a). Studies of both modern (e.g., Lawson
(1979)) and ancient (e.g., Cornamusini et al. (2017)) glacial sedimentary
systems have observed how topographic effects can result in seemingly
contradictory ﬂow directions. This study is another example. Transport
directions at all of the sites in this study appear contradictory, unless
paleotopography and processes that created each feature are considered
(Fig. 11). For example, at Mt. Butters the ﬂow directions of Permian
glaciers in the TAB appears to be from north to south, while the transport
directions of gravity-driven deposits and ﬂow directions in the grounding-
line fan are perpendicular to that (Fig. 11D, E). If averaged together, these
ﬂow directions would imply general transport toward the southwest, when
the most likely scenario was that the mass- and current-transported deposits
followed a paleotopographic slope and the glacier did not.
The Pagoda Fm in the Shackleton Glacier region likely represents a
single glacial–interglacial cycle, stratigraphically represented by a single
glacial sequence. In this context, the phrase ‘‘glacial–interglacial cycle’’
refers to the advance of a glacier into the basin and its whole retreat out
of the basin, which is stratigraphically deﬁned by a surface of glacial
retreat. This is not to say that the position of the glacier margin did not
ﬂuctuate during that cycle, but that only one grounded erosional surface
is present in the study area and only one surface of glacier retreat (Powell
and Cooper 2002; Rosenblume and Powell 2019). No instances of a
grounded readvance over any of the sections examined in this study were
observed beyond the basal erosional surfaces. As previously discussed,
the processes and environments responsible for the extensive deposition
of the massive, glaciogenic diamictite facies were likely diverse and
largely subaqueous. The preservation of such sediments is most probable
if they were deposited during the ﬁnal retreat phase with no subsequent
advances over the area (Kurjanski et al. 2020). This is especially true in a
basin-marginal position in a basin like the TAB that was trough-shaped
and not exposed to open-marine conditions. Evidence for glacier
readvance above the basal erosional surface of the Pagoda Fm does
exist at other Pagoda Fm outcrops in the TAB, including in the
Beardmore Sub-basin (Lindsay 1970a; Miller 1989; Koch and Isbell
2013). This evidence typically occurs in thicker, basinal facies
association which are likely areas with higher accommodation than the
The ﬂow directions and facies analyses from this study strongly support
the hypothesis that an ice center was positioned inboard of the
Panthalassan margin of Antarctica (an area that is now Marie Byrd Land)
during the Asselian–Sakmarian (Fig. 2B, ice center ‘‘q’’ ). The presence of
such an ice sheet is a relatively new hypothesis that was ﬁrst proposed by
Isbell (2010) and Isbell et al. (2008c) based on transport directions in
South Victoria Land. In recent publications, this proposed ice center on the
Panthalassan side of the TAB has been inconsistently included (Fielding et
al. 2008c; Isbell et al. 2012; Monta˜
nez and Poulsen 2013) and excluded
(Fielding et al. 2010; Craddock et al. 2019) from LPIA ice-center
reconstructions. Evidence from this study and Isbell (2010) shows that an
ice center should be included on the non-cratonic edge of the TAB in
reconstructions including the Gzhelian–Sakmarian, an interval also
EARLY PERMIAN SOUTH POLAR GLACIATIONJSR 631
by USGS Library user
on 18 June 2021
referred to as ‘‘Event 5’’ (L ´opez-Gamund´ı et al. 2021), and Australian ‘‘P1’’
(Fielding et al. 2008b). Ice-ﬂow directions elsewhere in the TAB clearly
indicate that glaciers also ﬂowed into the TAB off of the East Antarctic
Craton and along the TAB’s basin axis toward the Wisconsin and Ohio
ranges (Fig. 2B, ice center ‘‘r’’ ). The multiple ice centers contributing to
sedimentation in the TAB may not have been synchronous in their
advances and retreats throughout the LPIA. This additional evidence for an
ice center over Marie Byrd Land is important because it helps to explicate
the hypothesis that glaciation during the LPIA consisted of asynchronous,
discrete ice centers and not a single, large ice sheet centered over
Antarctica (Isbell et al. 2012; Monta ˜
nez and Poulsen 2013) (Fig. 2B).
Inferences made from glaciogenic and glacially inﬂuenced sediments
(‘‘near-ﬁeld’’ records) can be tied to global and ‘‘far-ﬁeld’’ records of
climate change from the ~80 Myr of the LPIA to approach a holistic
understanding of the effects that the onset, duration, and ultimate collapse
of a global icehouse-inﬂuenced Earth, both geologically and biologically
nez et al. 2007; Rygel et al. 2008; Soreghan et al. 2019).
The Pagoda Fm in the Shackleton Glacier region is glaciogenic and
was deposited in a basin-marginal subaqueous setting, by a glacier with a
temperate or mild subpolar thermal regime. The dominant lithology in the
Pagoda Fm here is massive, sandy, clast-poor diamictite. The depositional
processes governing these diamictites were subaqueous glacial processes
in a marine or brackish setting; likely a combination of mass transport,
iceberg rain-out, iceberg scouring, plume sedimentation, and subglacial
till deposition. Current-transported sands and stratiﬁed diamictites in the
Pagoda Fm were deposited as part of grounding-line fan systems. In the
Shackleton Glacier region, all glaciogenic sediments in the Pagoda Fm
were likely deposited during the retreat phase of a single glacial
sequence. The transport directions and thicknesses of strata along the
TAB margin were strongly controlled by topographic relief on the
underlying erosional surface. Glacier ﬂow directions (towards the south
and southeast) and trends in Pagoda Fm thicknesses in the Shackleton
Glacier Area support the hypothesis that an ice center was present toward
the Panthalassan margin of East Antarctica (Marie Byrd Land) during the
Supplemental appendices are available from SEPM’s Data Archive: https://
This work would have been impossible without the hard work of all the
people who made the 2017–2018 Shackleton Deep Field Camp such a success,
including the talented people of the National Science Foundation, Antarctic
Support Contract, Ken Borek Air, Petroleum Helicopters, Inc., New York Air
National Guard, and the U.S. Air Force. Special thanks are owed to Edith Taylor
and Rudolf Serbet for aiding in the ﬁeld planning, Danny Uhlmann and Ted
Grosgebauer for keeping us from tumbling off of cliffs, and to Patty Ryberg,
Rudolf Serbert, Brian Atkinson, and Erik Gulbranson for their companionship
and cooking in the deep ﬁeld. Thanks also to Kate Pauls and Eduardo Rosa for
their feedback on the manuscript. Funding for this research came from National
Science Foundation OPP-1443557, EAR-1729219, and OISE-1559231 grants,
the University of Wisconsin–Milwaukee Graduate Fellowships programs, the
P.E.O. Scholar Awards Program, The American Federation of Mineralogical
Societies, The Wisconsin Geological Society, University of Wisconsin–
Milwaukee (RGI grant), and the University of Wisconsin–Milwaukee
Department of Geosciences. And ﬁnally, thanks are due to Fernando Vesely
and an anonymous reviewer for their thoughtful comments and suggestions that
improved this manuscript.
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