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Late Oligocene obliquity-paced contourite sedimentation in the Wilkes Land margin of East Antarctica: implications for paleoceanographic and ice sheet configurations

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Abstract

The late Oligocene experienced atmospheric concentrations of CO2 between 400 and 750 ppm, which are within the IPCC projections for this century, assuming unabated CO2 emissions. However, Antarctic ice sheet and Southern Ocean paleoceanographic configurations during the late Oligocene are not well resolved, but are important to understand the influence of high-latitude Southern Hemisphere feedbacks on global climate under such CO2 scenarios. Here, we present late Oligocene (26–25 Ma) ice sheet and paleoceanographic reconstructions recorded in sediments recovered by IODP Site U1356, offshore of the Wilkes Land margin in East Antarctica. Our study, based on a combination of sediment facies analysis, physical properties, and geochemical parameters, shows that glacial and interglacial sediments are continuously reworked by bottom-currents, with maximum velocities occurring during the interglacial periods. Glacial sediments record poorly ventilated, low-oxygenation bottom water conditions, interpreted to represent a northward shift of westerly winds and surface oceanic fronts. During interglacial times, more oxygenated and ventilated conditions prevailed, which suggests enhanced mixing of the water masses with enhanced current velocities. Micritic limestone intervals within some of the interglacial facies represent warmer paleoclimatic conditions when less corrosive warmer northern component water (e.g. North Atlantic sourced deep water) had a greater influence on the site. The lack of iceberg rafted debris (IRD) throughout the studied interval contrasts with early Oligocene and post-Oligocene sections from Site U1356 and with late Oligocene strata from the Ross Sea (CRP and DSDP 270), which contain IRD and evidence for coastal sea ice and glaciers. These observations, supported by elevated paleotemperatures and the absence of sea-ice, suggest that between 26 and 25 Ma reduced glaciers or ice caps occupied the terrestrial lowlands of the Wilkes Land margin. Unlike today, the continental shelf was not over-deepened, and thus marine-based ice sheet expansion was likely limited to coastal regions. Combined, these data suggest that ice sheets in the Wilkes Subglacial Basin were largely land-based, and therefore retreated as a consequence of surface melt during late Oligocene, rather than direct ocean forcing and marine ice sheet instability processes as it did in younger past warm intervals. Spectral analysis on late Oligocene sediments from the eastern Wilkes Land margin show that the glacial-interglacial cyclicity and resulting displacements of the Southern Ocean frontal systems between 26–25 Ma were forced by obliquity.
1
Late Oligocene obliquity-paced contourite sedimentation in the Wilkes Land
margin of East Antarctica: implications for paleoceanographic and ice sheet
configurations
5
Keywords:
Late Oligocene
Paleoceanography
Antarctic Ice sheet
Contourites 10
Obliquity
Ariadna Salabarnada1, Carlota Escutia1, Ursula Röhl2, C. Hans Nelson1, Robert McKay3, Francisco F.
Jiménez-Espejo4, Peter K. Bijl5, Julian D. Hartman5, Minoru Ikehara6, Stephanie L. Strother9, Ulrich 15
Salzmann9, Dimitris Evangelinos1, Adrián López-Quirós1, José Abel Flores7, Francesca Sangiorgi5,
Henk Brinkhuis5, 8
1Instituto Andaluz de Ciencias de la Tierra, CSIC-Univ. de Granada, Armilla, 18100, Spain
20
2MARUM - Center for Marine Environmental Sciences, University of Bremen, Leobener Strasse 8, 28359
Bremen, Germany
3Antarctic Research Centre, Victoria University of Wellington, Wellington, 6140, New Zealand
4Department of Biogeochemistry, Japan Agency for Marine-Earth Science and Technology, Yokosuka,
Kanagawa, 237-0061, Japan 25
5 Department of Earth Sciences, Marine Palynology and Palaeoceanography, Faculty of Geosciences, Laboratory
of Palaeobotany and Palynology, Utrecht University, Heidelberglaan 2, 3584CS Utrecht, The Netherlands
6Center for Advanced Marine Core research, Kochi University, Nankoku, Kochi, 783-8502, Japan
7Department of Geology, University of Salamanca, Salamanca, 37008, Spain
8NIOZ, Royal Netherlands Institute for Sea Research, and Utrecht University, Landsdiep 4, 1797SZ ‘t Horntje, 30
Texel, The Netherlands
9Department of Geography, Faculty of Engineering and Environment, Northumbria University, Newcastle upon
Tyne NE1 8ST, UK
35
Correspondence to: Ariadna Salabarnada (a.salabarnada@csic.es)
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Abstract
The late Oligocene experienced atmospheric concentrations of CO2 between 400 and 750 ppm, which
are within the IPCC projections for this century, assuming unabated CO2 emissions. However, Antarctic 40
ice sheet and Southern Ocean paleoceanographic configurations during the late Oligocene are not well
resolved, but are important to understand the influence of high-latitude Southern Hemisphere feedbacks
on global climate under such CO2 scenarios. Here, we present late Oligocene (26-25 Ma) ice sheet and
paleoceanographic reconstructions recorded in sediments recovered by IODP Site U1356, offshore of
the Wilkes Land margin in East Antarctica. Our study, based on a combination of sediment facies 45
analysis, physical properties, and geochemical parameters, shows that glacial and interglacial sediments
are continuously reworked by bottom-currents, with maximum velocities occurring during the
interglacial periods. Glacial sediments record poorly ventilated, low-oxygenation bottom water
conditions, interpreted to represent a northward shift of westerly winds and surface oceanic fronts.
During interglacial times, more oxygenated and ventilated conditions prevailed, which suggests 50
enhanced mixing of the water masses with enhanced current velocities. Micritic limestone intervals
within some of the interglacial facies represent warmer paleoclimatic conditions when less corrosive
warmer northern component water (e.g. North Atlantic sourced deep water) had a greater influence on
the site. The lack of iceberg rafted debris (IRD) throughout the studied interval contrasts with early
Oligocene and post-Oligocene sections from Site U1356 and with late Oligocene strata from the Ross 55
Sea (CRP and DSDP 270), which contain IRD and evidence for coastal sea ice and glaciers. These
observations, supported by elevated paleotemperatures and the absence of sea-ice, suggest that between
26 and 25 Ma reduced glaciers or ice caps occupied the terrestrial lowlands of the Wilkes Land margin.
Unlike today, the continental shelf was not over-deepened, and thus marine-based ice sheet expansion
was likely limited to coastal regions. Combined, these data suggest that ice sheets in the Wilkes 60
Subglacial Basin were largely land-based, and therefore retreated as a consequence of surface melt
during late Oligocene, rather than direct ocean forcing and marine ice sheet instability processes as it
did in younger past warm intervals. Spectral analysis on late Oligocene sediments from the eastern
Wilkes Land margin show that the glacial-interglacial cyclicity and resulting displacements of the
Southern Ocean frontal systems between 26-25 Ma were forced by obliquity. 65
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1. Introduction
Today, ice sheets on Antarctica contain about 26.5 million cubic kilometres of ice, which has the 70
potential for raising global average sea level by 58 m, with the East Antarctic Ice Sheet constituting
53.3 m of this sea level equivalent (Fretwell et al., 2013). Satellite observations indicate significant rates
of change in most of the West Antarctic Ice Sheet (WAIS) and some sectors of the East Antarctic Ice
Sheet (EAIS). These include thinning at their seaward margins (Pritchard et al., 2012) and accelerating
ice shelves basal melt rates (Rignot et al., 2013). Given the uncertainties in projections of future ice 75
sheet melt, there has been a growing number of studies of sedimentary sections from the surrounding
margins of Antarctica targeting records of past warm intervals (i.e., high-CO2 and elevated temperature
climates) in order to better understand ice sheets and Southern Ocean configuration under these
conditions. For example, the early Pliocene (5-3 Ma) has been targeted because atmospheric CO2
concentrations were similar to today’s (400 ppmv) concentrations (Foster and Rohling, 2013; Zhang et 80
al., 2013). These studies have shown that early Pliocene Southern Ocean surface waters were much
warmer (i.e., between 2.5- > 4 ºC) than present and that the summer sea ice cover was greatly reduced,
or even absent (Bohaty and Hardwood, 1998; Whitehead and Bohaty, 2003; Escutia et al., 2009; Cook
et al., 2013). They also record the periodic collapse of both the WAIS and EAIS marine-based margins
(Naish et al., 2009; Pollard and DeConto, 2009; Cook et al., 2013; Reinardy et al., 2015; DeConto and 85
Pollard, 2016). Foster and Rohling (2013) demonstrated a sigmoidal relationship between eustatic sea-
level and atmospheric CO2 levels whereby sea levels stabilise at ~22+/-12m above present day level
between about 400 ppm and 650 ppm, suggesting loss of the Greenland Ice Sheet and the marine-based
West Antarctic Ice Sheet (+11 m s.l.e.). This implies that continental EAIS volumes remained relatively
stable during these times, but experienced mass loss of some (or all) its marine–based margins (19 m 90
s.l.e), relative to the present day. With CO2 concentrations at > 650 ppm they infer further increases in
sea level, suggesting this as a threshold for initiating retreat of the terrestrial margins of EAIS. With
sustained warming, CO2 concentrations of more than 650 ppmv are within the projections for this
century (Solomon, 2007; IPCC 2014). The last time the atmosphere is thought to have experienced CO2
concentrations above 650 ppmv was during the Oligocene (23.03-33.9 Ma), when CO2 values remained 95
between 400 to ~750-800 ppm (Pagani et al., 2005; Beerling and Royer, 2011; Zhang et al., 2013).
Geological records of heavy isotope values ~2.5 ‰ and far field sea level records from passive margins
during the Oligocene suggest that, following the continental-wide expansion of ice during the Eocene-
Oligocene transition that culminated at the Oi-1 event (33.6 Ma), the Antarctic ice cover was at least 100
~50 % of the current volume (e.g., Kominz and Pekar, 2001; Zachos et al., 2001; Coxall et al., 2005;
Pekar et al., 2006; Liebrand et al., 2011, 2017; Mudelsee et al., 2014). The early part of the Oligocene
records a significant δ18O decreasing slope with high-latitude sites exhibiting a strong
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deglaciation/warming that persisted until ~32 Ma (Mudelsee et al., 2014). This was followed by
seemingly stable conditions on Antarctica as evidenced by minimal δ18O and Mg/Ca changes (Billups 105
and Schrag, 2003; Lear et al., 2004; Mudelsee et al., 2014). A slight glaciation/cooling is recorded
before ~27 to 28 Ma, which was followed by an up to 1 ‰ long-term decrease in the δ18O isotope
records that was interpreted to result from the deglaciation of large parts of the Antarctic ice sheets
during a significant warming trend in the late Oligocene (27-26 Ma) (Zachos et al., 2001a).
Nevertheless, there are marked differences between the late Oligocene low δ18O values recorded in 110
Pacific, Indian and Atlantic Ocean sites (e.g., Pälike et al., 2006; Cramer et al., 2009; Liebrand et al.,
2011; Mudelsee et al., 2014; Hauptvogel et al., 2017), and the sustained high δ18O values recorded in
Southern Ocean sites (Pekar et al., 2006; Mudelsee et al., 2014). High δ18O values in the Southern
Ocean sediments are in agreement with the ice proximal record recovered by the Cape Roberts Project
(CRP) in the Ross Sea, which show the existence of glaciers/ice sheets at sea level (Barrett et al., 2007; 115
Hauptvogel et al., 2017). Based on the study of the isotopic record in sediments from the Atlantic, the
Indian and the equatorial Pacific, Pekar et al. (2006) explained the conundrum of a glaciated Antarctica,
and varying intrabasinal δ18O values with the coeval existence of two deep-water masses, one sourced
from Antarctica and another, warmer bottom-water, sourced from lower latitudes. Superimposed on the
above long-term swings in the δ18O Oligocene record, fluctuations on timescales shorter than several 120
Myr were identified in the high-resolution record from ODP 1218 (Pälike et al., 2006). These
fluctuations in periods of 405 kyr and 1.2 Myr are related to Earth’s orbital variations in eccentricity
and obliquity, respectively and have been referred as the short-term “heartbeat” of the Oligocene
climate (Pälike et al., 2006). Oligocene records close to Antarctica are needed to better resolve
Antarctic ice sheet and paleoceanographic configurations and variations at different timescales and 125
under scenarios of increasing atmospheric CO2 values and δ18O records, which imply a climatic
warming and/or ice volume loss.
Integrated Ocean Drilling Program (IODP) Expedition 318 drilled a transect of sites across the eastern
Wilkes Land margin at the seaward termination of the Wilkes Subglacial Basin (WSB) (Escutia et al., 130
2011; Escutia et al., 2014) (Fig. 1). Good recovery (78.2 %) of late Oligocene (26-25 Ma) sediments
from Site U1356 between 689.4 and 641.4 meters below sea floor (mbsf) provides an opportunity to
study ice-sheet and ocean configurations during the late Oligocene and to relate them with other
Antarctic and global records. In this paper, we present a new glacial-interglacial sedimentation and
paleoceanographic model for the distal glacio-marine record of the Wilkes Land margin constructed on 135
the basis of sedimentological data (visual core description, facies analysis, computed tomography
images, and high-resolution scanning electron microscopy images), selected physical properties data
(magnetic susceptibility), and X-ray fluorescence data (XRF). We also provide insights into the
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configuration of the ice sheet in this sector of the east Antarctic margin and evidence for orbital forcing
of the glacio-marine glacial-interglacial sedimentation at Site U1356. 140
2. Materials and Methods
2.1 Site U1356 description
145
Site U1356 (63º 18.6138’S, 135º 59.9376’E) is located at 3992 m water depth in front of the glaciated
margin of the eastern Wilkes Land Coast of East Antarctica, and penetrated 1006 meters into the flank
of a levee deposit in the transition between the lower continental rise and the abyssal plain (Escutia et
al., 2011; Fig. 1). Overall recovery was 35% with sediments dated between the early Eocene and
Pliocene, but several intervals provide good stratigraphic control (Escutia et al., 2011; Tauxe et al., 150
2012). The Oligocene section was recovered between 895 and 434 mbsf, Cores U1356-95R-3 83 cm to
U1356-46R. Our study focuses on the relatively high-recovery (78.2 %) interval within the late
Oligocene, which spans from 689.4 to 641.4 mbsf (Cores U1356-72R to -68R). The sediments from this
interval are part of shipboard lithostratigraphic Unit V, which is characterized by light greenish-grey,
strongly bioturbated claystones and micritic limestones interbedded with dark brown, sparsely 155
bioturbated, parallel- and ripple-laminated claystones with minor cross-laminated interbeds (Escutia et
al., 2011). The bioturbated and calcareous claystones and limestones were broadly interpreted to
represent pelagic sedimentation superimposed on the background hemipelagic sedimentary input
(Escutia et al., 2011). The laminated claystones and ripple cross-laminated sandstones were interpreted
to likely result from variations in bottom current strength and fine-grained terrigenous supply (Escutia 160
et al., 2011). In addition, a notable absence of Ice Rafted Debris (IRD) in this interval relative to
underlying and overlying strata was also recorded.
Today, Site U1356 lies close to the Southern Boundary of the Antarctic Circumpolar Current, near the
Antarctic Divergence at ~63ºS (Orsi, 1995; Bindoff, 2000) (Fig. 1). However, the paleolatitude of Site 165
U1356 was around 58.5±2.5ºS (van Hinsbergen et al., 2015) during the late Oligocene, more northerly
than today. Scher et al. (2008, 2015) reconstructed the position of the early Oligocene Antarctic
Divergence to be located around 60ºS (Fig. 1), based on the distribution of terrigenous and biogenic
(calcareous and siliceous microfossils) sedimentation, Nd isotopes, and Al/Ti ratios through a core
transect across Australian-Antarctic basin in the Southern Ocean. According to these interpretations Site 170
U1356 lay far to the north of the Antarctic Divergence zone, and was closer to the Polar Front, during
the Oligocene.
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2.2 Age Model
175
The age model for Site U1356 was established on the basis of the magnetostratigraphic datums
constrained by marine diatom, radiolaria, calcareous nannoplankton and dinocyst biostratigraphic
control (Escutia et al., 2011; Tauxe et al., 2012; Bijl et al., in press). The late Oligocene interval
contains three magnetostratigraphic datums (Table 1): 1) Chron C8n.1n (o) between 643.70 and 643.65
mbsf (U1356-68R-2); 2) C8n.2n (y) between 652.60 and 652.55 mbsf (U1356-69R-2), and 3) C8n.2n 180
(o) between 679.90 and 678.06 mbsf (U1356-71R). For this study, the age model by Tauxe et al.
(2012), which was calibrated to the Gradstein 2004 Geological Time Scale, has been updated to the
Geological Time Scale of Gradstein et al. (2012). Based on this calibration, the age of sediments
between 678.98 and 643.37 mbsf is 25.99 and 25.26 Ma, respectively (Fig. 2; Table 1).
185
2.3 Facies Analyses
Lithofacies are determined on the basis of detailed visual logging of the core during a visit to the IODP-
Gulf Coast Repository (GCR), expanding on the lower resolution descriptions in Escutia et al. (2011).
For this analysis, we logged the lithology, sedimentary texture (i.e., shape, size and distribution of
particles) and structures with a focus on the contacts between the beds and on bioturbation in cores 190
expanding from 896 to 392 mbsf (Cores U1356-95R to -42R) (see Supplementary material S1 Fig. S1,
S2). Physical properties data were measured during IODP Exp. 318 using the Whole-Round
Multisensor Logger. Magnetic susceptibility measurements were taken at 2.5 cm intervals, and Natural
gamma radiation (NGR) was measured every 10 cm (Escutia et al., 2011). In this paper, we focus on the
interval between 689.4 and 641.4 mbsf that comprise cores 72R to 68 R (Fig. 2). 195
X-ray Computed Tomography scans (CT-scans) measure changes in density and allow for analysis of
fine-scale stratigraphic changes and internal structures of sedimentary deposits in a non-destructive
manner (e.g., Duliu, 1999; St-Onge and Long, 2009; Van Daele et al., 2014; Fouinat et al., 2017). To
further characterize the different facies in our cores, selected intervals of Core U1356-71R-6 (678.11 to 200
676.91 mbsf) and Core U1356-71R-2 (672.8 to 671.35 mbsf) were CT-scanned at the Kochi Core
Center (KCC) (Japan). For this, we used the GE Medical systems LightSpeed Ultra 16. 2D scout
shooting conditions at 120Kv with 100mA, and 3D Helical image with 120Kv and 100mA and
FOV=22.0. Image spatial resolution consisted of 0.42 mm/pixel with 0.625 mm of slice thickness
(voxel spatial resolution of 0.42 x 0.42 x 0.625 mm). 205
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The type and composition of biogenic and terrigenous particles, particle size, and morphology of each
lithofacies was characterized with a high-resolution scanning electron microscope (HRSEM) at the
Centro de Instrumentación Científica (University of Granada, Spain).
210
2.4 X-Ray Fluorescence (XRF) analyses
Non-destructive X-ray fluorescence (XRF) core scanning measurements were collected every 2 cm
down-core over a 1 cm2 area with split size of 10 mm, a current of 0.2 mA (Al - Fe) and 1.5 mA (all
other elements) respectively, and a sampling time of 20 seconds, directly at the split core surface of the
archive half with XRF Core Scanner III at the MARUM – Center for Marine Environmental Sciences, 215
University of Bremen, Germany. Prior to the scanning, cores were thermally equilibrated to room
temperature, the surface was cleaned, flattened, and covered with 4 µm thin SPEXCerti Prep Ultralene1
foil to protect the sensor and prevent contamination during the scanning procedure. Scans were
collected during three separate runs using generator settings of 10 kV for the elements Al, Si, S, K, Ca,
Ti, Mn, Fe; 30 kV for elements such as Br, Rb, Zr, Mo, Pb; and 50 kV for Ba. The here reported data 220
have been acquired by a Canberra X-PIPS Silicon Drift Detector (SDD; Model SXD 15C-150-500) with
150eV X-ray resolution, the Canberra Digital Spectrum Analyzer DAS 1000 and an Oxford Instruments
100W Neptune X-ray tube with rhodium (Rh) target material. Raw data spectra were processed by the
Analysis of X-ray spectra by Iterative Least square software (WIN AXIL) package from Canberra
Eurisys. 225
This non-destructive method yields element intensities on the surface of split sediment cores and
provides statistically significant data for major and minor elements (Richter et al., 2006; O’Regan et al.
2010, Wilhelms-Dick et al., 2012). Detailed bulk-chemical composition records acquired by XRF core
scanning allows accurate determination of sedimentological changes as well as assessment of the
contribution of the various components in the biogenic and lithogenic fraction of the marine sediments 230
(Croudace et al., 2006). The data are given as element intensities in total counts. The light elements Al,
Si, and K show large element variations (intra-element variations of 1 order of magnitude or more, Fig.
2). Similar variations have been previously described in sediment cores to indicate substantial analytical
deviations due to physical sedimentary properties (i.e. Tjallingii and Röhl et al., 2007; Hennekam and
de Lange 2012). Accordingly, for this study we have concentrated our interpretations on Al, Si and K 235
values from the XRF analyses in discrete samples (see below). As Titanium (Ti) is restricted to the
terrigeneous phase in sediments and is inert to diagenetic processes (Calvert and Pedersen, 2007), we
utilized Ti to normalize other chemical elements for the terrigenous fraction. Linear correlation (R
Pearson) above standardised values has been done in order to find statistical relationships among the
variables. 240
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In addition, we conducted measurements of a total of 50 major and minor trace elements in 25 discrete
sediment samples collected at 0.4 and 1 m spacing to determine their chemical composition. For this,
we used a Pioneer-Bruker X-Ray Fluorescence (XRF) spectrometer S4 at the Instituto Andaluz de
Ciencias de la Tierra (CSIC-UGR) in Spain, equipped with a Rh tube (60 kV, 150 mA) using internal 245
standards. The samples were prepared in a Vulcan 4Mfusion machine and the analyses performed using
a standard-less spectrum sweep with the Spectraplus software.
2.5 Spectral Analyses
We selected key environmental indicators from XRF core scanner data and elemental ratios (i.e., Zr/Ba, 250
Ba, Zr/Ti, Ca/Ti, MS) to conduct spectral analyses on the data from the interval between 689.4 to 641.4
mbsf (Cores U1356-72R to -68R). We performed evolutionary spectral and harmonic analysis on each
dataset using Astrochron toolkit on the R software (Meyers, 2014). Detailed methodology is provided as
supplementary information following the Astrochron code of Wanlu Fu et al. (2016). This method
allows the detection of non-stationary spectra variability within the time series. The time series were 255
analysed in the depth scale and then anchored to the obliquity solution (Laskar 2004) to transform them
to an age scale, with the basis of the already resolved age model. The Evolutionary Average Spectral
Misfit method was then used to resolve unevenly sampled series and changing sedimentation rates
(Meyers et al., 2012). This method is used to test a range of plausible timescales and simultaneously
evaluate the reliability of the presence of astronomical cycles. The eccentricity, obliquity and precession 260
target periods were determined from La04 (Laskar et al., 2004) using the interval from 25.0 26.4 Ma
(Supplemental material S2).
3. Results
3.1 Sedimentary facies 265
The revised Oligocene facies log (Fig S1, S2), includes the high-recovery interval between 689.4 and
641.4 mbsf (Fig. 2). The integration of our lithofacies analyses, with physical properties (MS), CT-sans
and HRSEM analyses characterize an alternation between two main facies (Facies 1 and 2) (Figs. 2, 3,
4). These two facies were already identified shipboard but the interpretation of these facies was limited.
Consequently, our analyses allow us to more comprehensively characterize the facies, and to provide a 270
more rigorous interpretation about the depositional environments and the processes involved in their
development.
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Facies 1 (F1) consists of slightly bioturbated greenish claystones with sparse (Fig. 3a) to common
laminations (Figs. 2, 3a-f; Table 2). Laminae, as described on shipboard, vary from 0.1 to 1 cm thick 275
and, based on non-quantitative smear slide observations, are composed of well-sorted silt to fine sand
size quartz grains (Escutia et al., 2011). Laminations can be planar, wavy, with ripple-cross lamination
structures (Escutia et al., 2011), and show faint internal truncation surfaces, mud offshoots, and internal
erosional surfaces (Fig. 3a-f). HRSEM analyses of the claystones show that the matrix is composed of
clay-size particles and clay minerals (Fig. 3g, i). In addition, they show rare calcareous nannofossils that 280
are partially dissolved (Fig. 3g, i). Authigenic carbonate crystals are also identified (Fig. 3i).
Bioturbation in F1 is scarce, ichnofossils in the sediments are dominated mainly by Chondrites Fig. 3d).
CT-scans also show the presence of Skolithos, with their vertical thin tubes filled with high-density
material suggesting they are pyritized (Fig. 3b). Pyrite was also observed in shipboard smear slides in
small abundances from the laminated facies in the studied interval (Escutia et al., 2011). Magnetic 285
susceptibility values within the laminated facies are low, between 40-70 MS instrumental units (iu),
with higher values when silt laminations are more abundant (Figs. 2, 4). Natural Gamma Ray (NGR) is
anti-correlated with MS, with high values in F1 varying between 50 - 65 counts per second (cps) (Fig.
2).
290
Facies 2 (F2) is composed by light greenish grey strongly bioturbated claystones and silty claystones
(Figs, 2, 3; Table 2) with variable carbonate content varying between 5-16% based on our XRF
analyses. No primary structures are preserved due to the pervasive bioturbation (Fig. 3a-c). Burrows are
backfilled with homogeneous coarse material (silt/fine sand). Different types of ichnofossils are present
with Planolites and Zoophycos being the most abundant (Fig. 3a, b). HRSEM images show: 1) silt-size 295
grains containing quartz grains with conchoidal fractures in the corners and impact marks on the crystal
faces, indicative of high-energy environments; and 2) biogenic carbonate consisting of moderately- to
poorly-preserved coccoliths, which exhibit dissolution of their borders, and to a minor degree detrital
carbonate grains (Fig. 3h - j). A total of 13 carbonate-rich layers have been observed within the studied
interval F2, and they range in thickness from 10 to 110 cm. F2 CT-scans images show an increase in 300
density (i.e., gradation towards lighter colours in the scan) towards the top of each bioturbated interval
(Fig. 3b). MS values are higher in F2 compared to F1. Values vary from 50-150 instrumental units (iu)
and exhibit an inverse grading or a bigradational-like morphology (Fig. 2, 4), while NGR is inversely
correlated with minimum values occurring in F2 (between 35-55 cps) (Fig. 2).
305
Contacts between the two facies are sharp and apparently non-erosive, with minimal omission surfaces
or lags (Figs. 3,4). However, when bioturbation is present, gradual contacts in the transition from F1 to
F2 also occur (Fig. 3b). Both sharp and transitional contacts are well imaged on the MS plots (Fig. 2).
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In addition, the CT-scan images confirm the shipboard visual absence of outsized clasts and coarse
sands grains (>0.5 mm) in F1 and F2. 310
3.2 Geochemistry
Down-core changes in the log ratios of various elements have been plotted against the facies log (Figs.
2, 4). In addition, in order determine geochemical element associations we performed a Pearson
correlation coefficient analysis of major elements on the whole XRF-scanner dataset (Table 3). This 315
analysis highlights two main groups that are used as proxies for terrigenous (i.e., Zr, Ti, Rb, Ba) vs
biogenic (i.e., Ca = carbonate) sedimentation.
Titanium (Ti), Zirconium (Zr), and Rubidium (Rb) are primarily derived from terrigenous sources,
where Ti represents the background terrigenous input. During sediment transport Zr, Rb and Ti tend to 320
become concentrated in particular grain-size fractions due to the varying resistance of the minerals in
which these elements principally occur. Zr tends to become more concentrated in fine sand and coarse
silt fractions, Ti in somewhat finer fractions and Rb principally in the clay-sized fraction due to their
typical mineralogical association and their natural presence in the different grain size categories
(Veldkamp and Kroonenberg 1993; Dypvik and Harris 2001). The lack of correlation between Zr and Ti 325
(Fig. 2; Table 3) implies that they are settled in different minerals and processes. The Zr/Rb ratio has
been applied as a sediment grain-size proxy in marine records (Schneider, et al., 1997; Dypvik and
Harris 2001; Croudace et al., 2006; Campagne et al., 2015). Zr/Al has been interpreted as an indicator
for the accumulation of heavy minerals due to bottom currents (Bahr et al., 2014). In our cores, Zr/Rb
and Zr/Ti ratios have a near identical variability downcore (Fig. 2). We utilize the high-amplitude Zr/Ti 330
signal in our records as indicator for larger grain-size and current velocity (Fig. 2). The Zr/Ti ratio
varies between 0.1 and 1 and exhibits maximum values within F2 showing an increasing upwards or
bigradational patterns (Fig. 2). Although minimum Zr values (cps) are found in F1, laminations within
this facies are also characterized by elevated Zr values similar to those in F2 (Figs. 3, 4; Table 3). The
Zr/Ti pattern is positively correlated with magnetic susceptibility throughout the studied interval (Fig. 335
2).
The Zr/Ti, Zr/Rb and Zr/Ba ratios co-vary characterizing the laminations within F1 and the alternation
between F1 and F2 by defining the contacts between them (Figs. 2, 4). They also mark the coarsening
upwards or bigradational tendency in F2 (Fig. 4). Of the three ratios, the Zr/Ba ratio is the one that 340
highlights these patterns best (Figs. 2, 4).
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Barium (Ba) is present in marine sediments mainly in detrital plagioclase crystals and in the form of
barite (BaSO4; Tribovillard et al., 2006). In the studied sediments, Ba and Ti have a correlation factor of
r2=0.66 (Table 3), which is taken to indicate that Barium is predominantly present as a constituent of the 345
continental terrigenous fraction and/or that biogenic barite was sorted by bottom currents. Ba has
maximum values (10,000 total counts) at the base of F1 and decreases upwards in a saw-tooth pattern,
reaching minimum concentrations within F2 (5,000 total counts) (Fig. 2; Table 3). The detrital fraction
of Ba in the open ocean has been used in other studies as a tracer of shelf waters (Moore and Dymond,
1991; Abrahamsen et al., 2009; Roeske, 2011) and Ba record also is affected by current intensity in 350
other depositional contourite systems (Bahr et al., 2014) preventing his use as paleoproductivity proxy
in environments dominated by contour currents.
Variations in Ca, Mn, and Sr are strongly intercorrelated (Fig. 2) with r2>0.87 (Table 3). Biogenic
calcite precipitated by coccoliths and foraminifera have greater Sr concentration than inorganically 355
precipitated calcite or dolomite (Hodell et al., 2008). The positive Ca and Sr correlation could therefore
potentially be used to differentiate between terrigenous Ca sources (e.g. feldspars and clays) and
biogenic carbonates (e.g. Richter et al., 2006, Foubert and Henriet, 2009, Rothwell and Croudace,
2015). Based on these observations, we interpret that Ca in our sediments is mainly of biogenic origin
(CaCO3). This interpretation is supported by HRSEM images taken from carbonate-rich intervals of F2, 360
which show abundant coccoliths (Fig. 3d). Peaks in Ca in our record (Fig. 2) coincide with the
carbonate-rich layers listed in the previous section. Additional peaks in the record may indicate
carbonate-rich layers that we have been unable to identify visually.
In order to estimate the CaCO3 content continuously throughout the studied interval we use a calibration 365
(r2U1356=0.81) between natural logarithm (ln) of Ca/Ti ratio (ln(Ca/Ti)) from the XRF core scanner data
and the XRF discrete CaCO3 measurements (weight %) from Site U1356 as applied in other studies
(Zachos et al., 2004; Liebrand et al., 2016) (Fig. 5). “CaCO3 est.” is used throughout the text to refer to
carbonate content estimated by ln(Ca/Ti) ratio. CaCO3 est. concentrations are generally low (between 0-
16%). Carbonates are mostly present in F2, varying between 5-16 %, although small contents (from 0 to 370
5 %) can be seen in the intervals of F1 with scarce laminations (Fig. 4). CaCO3 est. peaks in some
intervals have a particular morphology producing a double peak in the beginning and/or the end of
bioturbated F2 (Figs. 2, 4).
Mn(II) is soluble under anoxic conditions and precipitates as Mn(IV) oxyhydroxides under oxidising 375
conditions (Tribovillard et al., 2006). Manganese is frequently remobilized to the sedimentary pore
fluids under reducing conditions. Dissolved Mn can thus migrate in the sedimentary column and
(re)precipitate when oxic conditions are encountered (Calvert and Pedersen, 1996). As such, large Mn
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enrichments primarily reflect changing oxygen levels at the sediment–water interface (Jaccard et al.,
2016). The strong-correlated peaks of Mn and Ca (Fig. 2; Table 3) suggest that at least some of the Mn 380
is present in the studied interval as Mn carbonates and/or Mn oxyhydroxides under oxic sediment-water
interphase (Calvert et al., 1996; and Calvert and Pedersen, 1996; Tribovillard et al., 2006; Calvert and
Pedersen 2007).
Br/Ti has been previously used as an indicator of organic matter in sediments (e.g., Agnihotri et al., 385
2008; Ziegler et al., 2008; Bahr et al., 2014). Br/Ti in our record shows generally low values (Fig. 2)
most likely as the organic matter content in both facies types is relatively low (<0.5 %, Escutia et al.,
2011). However, it exhibits some variability (0.01 to 0.05 Br/Ti ratio) within the two facies with higher
ratio values in F1. Darker coloured sediments in F1 are in agreement with these higher Br/Ti values
inside F1. 390
In addition to the elemental analyses of the XRF-scanned data, we use the detrital Al/Ti ratio in discrete
XRF bulk sediment samples to reflect changes in terrigenous provenance (Kuhn and Diekmann, 2002;
Scher et al., 2015). Al/Ti ratio varies between 17-21, with the highest values found within F1 and the
lowest in F2 (Fig. 2). 395
3.4 Spectral analysis
To detect periodical signals, spectral analysis of time series was performed on the Zr/Ba and other
elemental proxies (i.e., Ba, Zr/Ti, CaCO3, Magnetic Susceptibility) using Astrochron R software
(Meyers, 2014) (Figs. 6; S3-10). 400
Multiple-taper spectral analysis (MTM) in Zr/Ba show a clear and statistically significant (>90%)
cyclicity every 2m (0.5 cycles/m), and at 4.67m (0.21 cycles/m), and less significant one (>80%) at 1m
(0.94 cycles/m) (Fig. S3). On the basis of a linearly calculated sedimentation rate between the two
extreme tie-points (Table 1), we obtained a sedimentation rate of approximately 5 cm/kyr. Within this 405
sedimentation rate, the 0.5 cycles/m peak corresponds to the 41-kyr obliquity frequency; and the 0.21
and 0.94 cycles/m to the 95 and 21-kyr shorter eccentricity periods and precession frequencies,
respectively.
After initial analysis, we ran an Evolutive Harmonic Analysis (EHA) (Astrochron (Meyers, 2014)) with 410
3 data tapers for the untuned Zr/Ba in depth domain with 2 cm resolution (Fig. S3). The statistical
significance of spectral peaks was tested relative to the null hypothesis of a robust red noise
background, AR(1) modelling of median smoothing, at a confidence level of 95% (Mann and Lees,
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1996). Despite a short core gap in the middle of the time series, obliquity (40 kyr) dominates throughout
the time series (Fig. 6). The sedimentation rates obtained by this method vary between 4.6 and 5.4 415
cm/kyr for the studied section, similar to those obtained with linearly calculated sedimentation rates.
Additionally, the Nyquist frequency for Zr/Ba data is 1 m-1 (0.5 kyr), which implies the site is sampled
sufficiently to resolve precessional scale variations however, core gaps prevent identification of long
eccentricity cycles (Fig. S6).
420
Apart from obliquity, spectral analysis of the tuned age model reveals an alignment of the eccentricity
and precession bands (Fig. 6, S8). For example, a marked cyclicity at the obliquity periods of 41 Kyr is
seen at Ba and Zr/Ti (99% confidence) and also eccentricity at 100 kyr, and precession at 20kyr (95%
confidence) (Fig. S9). We also observe coherent power above the 90% significance level at ~54 and
~29 ky periods, which are secondary components of obliquity. The anchored age model provides an 425
unprecedented 500 yr resolution (2.5 cm sampling) of the data during the Late Oligocene. Orbital
frequencies were tested in each core section individually in the Zr/Ba dataset in the depth scale in order
to assure that cyclicity is not an artefact related to the gaps in the series (Fig. S10).
4. Discussion 430
4.1 Glacial and interglacial contourite sedimentation off Wilkes Land
Alternations between laminated glacial deposits and hemipelagic deposits have been previously
reported to characterize Pleistocene and Pliocene glacial-interglacial continental rise sedimentation,
respectively, on this sector of the Wilkes Land margin (Escutia et al., 2003; Patterson et al., 2014).
Gravity flows, mainly turbidity flows are the dominant process during glacial times resulting in 435
laminated deposits. Interglacial sedimentation is dominated by hemipelagic deposition with higher opal
and biogenic content (Escutia et al., 2003). Erosion and re-deposition of fine-grained sediment by
bottom contour currents has also been reported as another important process during Pleistocene
interglacials (Escutia et al., 2002; Escutia et al., 2003).
440
The depositional setting on the continental rise was however different during the late Oligocene. The
stratigraphic evolution of the region testifies the progradation of the continental shelf taking place after
continental ice sheet build-up during the Eocene-Oligocene Transition (EOT, 34 Ma) (Eittreim et al.,
1995; Escutia et al., 1997; Escutia et al., 2005), which resulted in: 1) seismic and sedimentary facies
becoming more proximal up-section (Hayes and Frakes, 1975; Escutia et el., 2000; Escutia et al., 2005; 445
Escutia et al., 2014), and 2) high sedimentation rates during the Oligocene (Escutia et al., 2011; Tauxe
et al., 2012). In this context, the late Oligocene sediments from Site U1356 record distal continental rise
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deposition in an incipient/low-relief levee of a channel-levee complex. As progradation continued, a
complex network of well-developed channels and high-relief levee systems developed on the
continental rise (Escutia et al., 2000) from the latest Oligocene onwards. 450
Laminated claystones (F1) from Site U1356 were originally interpreted by the shipboard science team
to have formed during glacial times relating to variations in bottom current strength and fine-grained
terrigenous supply. Conversely, the bioturbated claystones and micritic limestones (F2) were interpreted
to result from mostly hemipelagic sedimentation during interglacial times (Escutia et al., 2011). In this 455
study, we have further characterized these facies on the basis of sedimentological data (visual core
description, facies analysis, CT-scans, HRSEM), physical properties (magnetic susceptibility, NGR);
and geochemical data (X-ray Fluorescence-XRF), which allow us to construct a sedimentation model
for the depositional setting of Site U1356 during the late Oligocene that is dominated by bottom-current
reworking of both, glacial and interglacial deposits. 460
Laminated, fossil-barren, glaciogenic deposits consistent with those of Facies F1 have been observed on
younger sedimentary sections from other polar margins and interpreted as contour current modified
turbidite deposits and as muddy contourites (Anderson et al., 1979; Mackensen et al., 1989; Grobe and
Mackensen, 1992; Pudsey, 1992; Gilbert et al., 1998; Pudsey and Howe, 1998; Pudsey and 465
Camerlenghi, 1998; Anderson, 1999; Williams and Handwerger, 2005; Lucchi and Rebesco, 2007,
Escutia et al., 2009). This particular type of glaciogenic contourite facies is associated with glacimarine
deposition during glacial times, and has been interpreted to result from unusual, climate-related,
environmental conditions of suppressed primary productivity and oxygen-poor deep-waters (Lucchi and
Rebesco, 2007). Despite being sparse, the occurrence of bioturbation in our sediments, which slightly 470
affects both claystones and silt laminations, indicates slow and continuous sedimentation, which is not
consistent with instantaneous turbidite deposition. It is however consistent with distal overbank fine-
grained sediments being entrained by bottom-currents. Silt layer sedimentary structures similar to those
described by Rebesco et al. (2008, 2014) indicate that there is current reworking of the sediments. For
example, silt layers can be continuous or discontinuous with wavy and irregular morphologies, and 475
within layers, sedimentary structures such as cross-laminations are common (Fig. 3c-f). Within the
cross laminae, mud offshoots and internal erosional surfaces are distinctive features of fluctuating
currents where successive traction and suspension events are super-imposed, indicating bottom-currents
sedimentation as the principal process for the F1 laminated claystones (Shanmugam et al., 1993; Stow
et al., 2002). Based on these observations, we interpret F1 as glacial laminated muddy contourites 480
following the classification of Stow and Faugères (2008). The F1 sedimentary structures suggest
bottom-currents with fluctuating intensities, that result in laminations and internal structures forming
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during peak current velocities (Lucchi and Rebesco, 2007; Martín-Chivelet et al. 2008; Rebesco et al.,
2014).
485
Bioturbated sediments in F2 were previously interpreted as interglacial hemipelagic deposits (Escutia et
al., 2011). In this study, we interpret F2 as hemipelagic and overbank deposits reworked by bottom-
currents. The coarser grain-size in F2 compared to F1 (silty-clay matrix as seen in HRSEM Fig. 3g-j),
the distribution of heavy minerals as indicated by the Zr/Ba, and the elevated values of the magnetic
susceptibility record with a bigradational pattern within the facies (Figs. 2,4), support the notion that 490
interglacial sediments of F2 have been heavily modified by bottom currents. Hemipelagic sediments are
expected to be homogeneous in terms of grain-size and grading is not expected. Current winnowing of
hemipelagic deposits and removal of the fine-grained fraction can produce the higher accumulation of
heavy (indicated by the Zr) and ferromagnetic (indicated by MS) minerals observed in F2 compared to
F1 (Fig. 2; Table 2). Bi-gradational trends have been previously described in contourite sediments and 495
interpreted to record an increase followed by decrease in the current velocities (e.g., Martín-Chivelet et
al., 2008). The bi-gradational patterns in the Zr/Ba and MS plots (Figs. 2,4) are therefore interpreted to
depict a constant and smooth increase followed by a decrease in current velocity with little gradual
changes in flow strength. In addition, the presence of grains of quartz with conchoidal fractures and
reworked coccolitospheres with signs of dissolution (Fig., 3h,j) support the reworking of background 500
hemipelagic and turbidite overbank sediments by bottom currents in a high-energy environment
(Damiani et al., 2006). Following the classification by Stow and Faugères (2008), we interpret that F2
has more silty massive contourites resulting from higher and more constant bottom current velocity
compared to F1.
505
Transitions between the F1 and F2 facies are characterized by glacial-to-interglacial contacts that may
be sharp or diffuse due to bioturbation, and characterized by a gradual change in physical and
geochemical sediment parameters (Figs., 3, 4; Table 3). Interglacial-to-glacial contacts (F2 to F1), on
the other hand, are characterized by an apparently non-erosional sharp lithological boundary. The sharp
lithological boundaries between interglacial to glacial transitions can be explained by maximum current 510
intensities achieved at the end of the interglacials (Shanmugam, 2008; Rebesco et al., 2014).
4.2. Ice sheet configuration during the warm late Oligocene
Early Oligocene sediments from Site U1356 contain outsized clasts interpreted as ice rafted debris
(IRD) (Escutia et al., 2011). In addition, dinocyst assemblages indicate the presence of sea ice (Houben 515
et al., 2013). Based on this, the site should have been within the reach of icebergs calving from an
expanded ice sheet grounded at the coast or beyond in the late Oligocene. This is supported by
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Pliocene-Pleistocene sedimentary sections in adjacent continental rise sites containing IRD (Escutia et
al., 2011; Patterson et al., 2014). Thus, the lack of IRD in our studied interval is taken to indicate the
relative absence of marine-terminating ice sheets at the nearby margin. 520
The interpretation of smaller ice sheets and partly ice-free margins is in agreement with the absence of
sea ice species Selenopemphix antarctica and common to abundant gonyaulacoid phototrophic
dinocysts, which suggest warm-temperate surface waters (Bijl et al., submitted, this volume). Overall 525
elevated sea surface temperatures (i.e., 10-15 ºC) based on biomarker sea surface temperatures (TEX86
data in Hartmann et al., submitted, this volume) support a sea ice-free scenario during the late
Oligocene. Furthermore, the presence of in situ terrestrial palynomorphs suggests that during the late
Oligocene margins nearby were in part free of ice sheets and covered by a cool-temperate vegetation
with trees and shrubs (Strother et al., in prep). 530
These observations are consistent with the iceberg survivability modelling in the Southern Ocean during
the warm Pliocene intervals, which shows the distance that icebergs could travel before melting was
significantly reduced (Cook et al., 2014). Warm Pliocene seasonal temperatures up to 6°C warmer than
today during interglacials and prolonged Pliocene warm intervals have been reported in the Ross Sea 535
(e.g., Naish et al., 2009; McKay et al., 2012) and other locations around Antarctica (Whitehead and
Bohaty, 2003; Whitehead et al., 2005; Escutia et al., 2009; Bart and Iwai, 2012). Contrary to what we
observe in our late Oligocene record, during the warm Pliocene abundant IRD were delivered to
adjacent continental rise sites (Escutia et al., 2011; Patterson et al., 2014). This was interpreted by Cook
et al. (2017) to suggest that a considerable number of icebergs (iceberg armadas) had to be produced in 540
order to reach the site under these warm conditions. We argue that the lack of IRD delivery to site
U1356 during the late Oligocene likely results from the different paleotopographic setting between the
Oligocene and the Pliocene. Paleotopographic reconstructions from 34 Ma ago (Wilson et al., 2012) and
the early Miocene (Gasson et al., 2016), show the Wilkes Subglacial Basin (WSB) to be an area of
lowlands and shallow seas in contrast to the over-deepened marine basin that it is today (Fretwell et al., 545
2013). This difference is important, as an ice sheet grounded on an overdeepened continental shelf can
experience marine ice sheet instability, a runaway process relating to ice sheet retreat across a reverse
slope continental shelf (Weertman 1974), which is proposed to be a driver for Pliocene collapse of the
WSB (Cook et al., 2013). This paleotopographic configuration would have precluded widespread
marine ice sheet instability during the Oligocene. Conversely, a shallower continental shelf allows for 550
the potential expansion of grounded ice sheets into the marine margin during warmer-than-present
climates (Wilson et al. 2012), and thus direct records are required to assess the climate threshold for
such an advance.
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In contrast to the distal U1356 Wilkes Land margin record, the Ross Sea Embayment ice proximal 555
sediments obtained by the Cape Roberts Project (CRP) contain Oligocene to Early Miocene
palynomorphs, foraminifera and clay assemblages that point to a progressive decrease in fresh
meltwater, cooling and intensifying glacial conditions (Leckie and Webb, 1983; Hannah et al., 2000;
Hannah et al., 2001; Raine and Askin, 2001; Thorn, 2001; Ehrmann et al., 2005; Barrett, 2007).
Therefore, the coastal CRP sediment record does not support a significant loss of ice or warming during 560
the late Oligocene (Barrett, 2007), as has been suggested by compilations of deep-sea benthic δ18O data
(Zachos et al., 2001). Moreover, sediments recovered at Deep Sea Drilling Project (DSDP) Site 270 on
the mid continental shelf of the Ross Sea contain IRD and pollen assemblages that provide evidence for
the coexistence of ice masses and vegetation through the Oligocene (Kemp and Barrett, 1975). The high
sedimentation rates during the late Oligocene-early Miocene at Site 270 were interpreted to reflect 565
turbid plumes of glaciomarine sediments derived from polythermal-style glaciers or ice sheets that were
calving into an open Ross Sea, without an ice shelf (Kemp and Barrett, 1975). In addition, seismic data
indicate that during the late-mid Oligocene widespread expansion of a marine-based ice sheet onto the
outer Ross Sea shelf did not take place but instead glaciers and ice caps drained from local highs and
advanced only into shallow marine areas, rather than whole-scale marine ice sheet advance (Brancolini 570
et al., 1995; DeSantis et al., 1995; Bart and De Santis, 2012).
Combined, this evidence suggests that late Oligocene marine-terminating glaciers, ice caps or ice sheets
persisted along the Transantarctic Mountain front in the Ross Sea, but not in the eastern Wilkes Land
margin. This suggests an ice sheet with a similar configuration as modelled for Miocene topographies 575
with CO2 scenarios of 500-840 ppm (Gasson et al., 2016) (Fig. 7). This is also supported by vegetation
reconstructions derived from fossil pollen from both margins, which indicate for the middle Miocene
and Late Oligocene higher terrestrial temperatures and more tree taxa at Wilkes Land (Salzmann et al.,
2016; Sangiorgi et al., 2017; Strother et al., in prep) than the Ross Sea (Askin and Raine, 2000; Prebble
et al., 2006). We therefore suggest that during the late Oligocene both vegetation and glaciers or ice 580
caps coexisted in the lowlands of the WSB, and that the ice did not extend significantly to the coast.
4.3 Paleoceanographic implications
Sediment physical properties and geochemical signatures of F1 and F2 are here related to changes in
bottom water-sediment interphase oxygenation/ventilation during successive glacial and interglacial 585
periods (Table 2). These changes are linked to shifts in water-masses driven by a north-south
displacement of the position of the westerlies, and associated changes in the intensity of frontal mixing
or location of the Polar Front and Antarctic Divergence (Fig. 7). Based on these observations, we
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propose a model to explain the interpreted changes in bottom-water conditions at Site U1356 during
successive glacial and interglacial times (Fig. 7). 590
4.3.1. Glacial paleoceanographic configuration
The Chondrites-like bioturbation with pyrite infilling the tubes of Skolithos within F1 (Fig. 3b, d), has
previously been reported to characterize low-oxygen conditions at the water-sediment interphase
(Bromley and Ekdale, 1984). In addition, pyritized diatoms are present throughout the Oligocene
section of this site, but are found preferentially inside F1. The presence of pyritized diatoms was 595
interpreted during Expedition 318 to indicate a prolific production and syn-sedimentary diagenesis in a
restricted circulation (low oxygen) environment, mainly during glacial periods (Escutia et al., 2011).
Reducing conditions in the sediment also help to preserve primary sedimentary structures of the silt
layers in F1 because bioturbation is limited. Higher amounts of organic matter in F1 compared to F2 are
suggested by increased values of the Br/Ti ratio (Fig. 2). This higher organic content most likely results 600
from oxygen depletion in the water-sediment interphase, which creates a poorly ventilated environment
with near reducing conditions where pyrite has been able to precipitate (Tribovillard et al., 2006). In
spite of this, no total oxygen depletion is observed, and is supported by palynomorphs preservation
inside F1 (Bijl et al., submitted, this volume).
605
High MS values result from stronger bottom currents deposition and/or increased terrigenous input
(e.g., Pudsey et al., 2000; Hepp et al., 2007). In our record, low MS values are found in F1 (Fig.4; Table
2). Low MS values around Antarctica have been attributed to MS dissolution caused by dilution and/or
primary diagenesis effects on the sediments due to the higher concentration in organic matter or to
changing redox conditions (Korff et al., 2016). Several authors have postulated that oxygen-depleted 610
Antarctic Bottom Water (AABW) occupying the abyssal zones of the oceans can change the redox
conditions in the sediment, trapping and preserving dissolved and particulate organic matter and,
consequently reducing and dissolving both, biogenic and detrital magnetite (Florindo et al., 2003; Hepp
et al., 2009; Korff et al., 2016). At present, Site U1356 is influenced by AABW forming in the adjacent
Wilkes Land shelf (Orsi et al., 1999; Fukamachi et al., 2000) and in the Ross Sea spilling over to the 615
Wilkes Land continental shelf (Fukamachi et al., 2010) (Fig. 1). Our records indicate a reduced
continental ice-sheet in the eastern Wilkes Land margin, likely not reaching the coastline, and reduced
sea ice presence compared to today (Bijl et al., submitted, this volume). Under these conditions, bottom
water formation and downwelling can still occur (with or without presence of sea ice) as a result of
density contrasts related to seasonal changes in surface water temperature and salinity (Huber and 620
Sloan, 2001; Otto-Bliesner et al., 2002). Moreover, stable Nd isotopic composition in Eocene-
Oligocene sediments from Site U1356 is consistent with modern day formation of bottom water from
Adélie Land, as reported by Huck et al (2017).
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Our evidence above points to deposition of F1 during glacial cycles under poorly-ventilated, low-625
oxygenation conditions at the water-sediment interface (Fig. 7a). We postulate, that during glacial
periods, westerly winds and surface oceanic fronts migrate towards the equator, generating a more
stratified ocean and reduced upwelling closer to the margin, with sporadic and fluctuating currents (Fig.
7a). Records of the Last Glacial Maximum show that this northward migration results in a weakening of
the upwelling of the Circumpolar Deep Water (CDW) (Govin et al., 2009), increasing stratification and 630
reduced mixing of water masses also due to an enhanced sea ice formation, not seen during the late
Oligocene.
4.3.2. Interglacial paleoceanographic configurations
635
The higher degree of bioturbation in F2 with no primary structures preserved and the ichnofacies
association (i.e., Planolites and Zoophycos), suggest a more oxygenated environment in comparison
with F1. This is supported by the covariance of Mn and CaCO3 est. (Fig 4) where Mn enrichments
indicate the redox state conditions at the sediment-water interface (Calvert and Pedersen, 2007). More
oxygenated conditions during interglacial periods can be achieved under a more ventilated and mixed 640
water masses, with enhanced current velocities. We interpret that F2 had enhanced current velocities
based on coarser grain size, and the increased accumulation of heavy and ferromagnetic minerals as
indicated by the high values of the Zr/Ti ratio and MS within F2 (Figs. 2,4). The bigradational pattern
of the Zr/Ba and the MS (Fig. 4) is also interpreted to record an increase followed by a decrease in
current velocities within F2. 645
The intervals of micritic limestone within F2 have calcareous nannofossils preserved (Fig 3d). The
productivity of calcareous nannofossils and the later preservation of these coccoliths in the sediment
indicate specific geochemical conditions enabling carbonate deposition and preservation. Although
today nannoplankton is abundant in surface waters at the Antarctic Divergence (Eynaud et al., 1999), 650
these rarely deposit on the deep ocean floor because of corrosive bottom waters, which dissolve
calcareous rain. A number of studies in other areas of the Antarctic margin have correlated the presence
of calcareous nannofossils during the Oligocene with the presence of temperate north component water
masses (NADW-like) that intrude close to the Antarctic continent and influences the Southern Ocean
during the late Oligocene (e.g., Nelson & Cook, 2001; Pekar et al., 2006; Villa & Persico, 2006; Scher 655
and Martin, 2008) and during more recent times such as the Quaternary (Diekman, 2007, Villa et al.,
2012). In addition, Pleistocene sedimentary records of past warm interglacial events in Antarctica also
have reported enhanced NADW production (e.g. interglacial event MIS11 from M. S. Poli et al., 2000,
Kemp et al., 2010; DeCesare et al., 2013).
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20
660
During interglacials, our records point to more oxygenated and ventilated conditions suggesting
enhanced mixing of the water masses (Fig. 7b-c). We postulate that during interglacials westerly winds
and the Polar Front are shifted south and become more aligned. Under these conditions, upwelling of
deep waters is promoted, facilitating the mixing and oxygenation of surface waters that form the
precursor to bottom water. Such a process would also generate increased geostrophic current velocities 665
of bottom water mass as evinced by the coarser grain size and heavy mineral concentrations in the
bioturbated F2 facies. During interglacials, bottom water formation is likely warmer and less saline due
to enhanced freshwater runoff from surface and subglacial melt of the continental ice sheet. This may
allow this less dense water mass to occupy shallower depths in abyssal to intermediate ocean, and
promote more vigorous mixing with oxygenated CDW (Fig. 7b). During warmer interglacials, the 670
influence of more northern-sourced water masses, relative to Antarctic-sourced, could enable carbonate
productivity as seen in the interglacial facies with coccolitosphere remains (Fig. 7c). This is also
reinforced by several interpretations that document a late Oligocene increase in the influence of North
Component Water (e.g. NADW-like) in the Southern Ocean (Billups et al., 2002; Pekar et al., 2006;
Villa and Persico, 2006; Scher and Martin, 2008; Liebrand et al., 2011). These data are also in 675
agreement with the δ13C global isotopes oscillations between 26 and 25 Ma (Cramer et al., 2009), that
suggest low values for an AABW and high δ13C values for a NADW, that may represent the different
oceanic primary production and ventilation rates, as proposed in this work. In addition, δ13C records on
the Atlantic show systematic offsets to lower values toward a North Atlantic signal for most of the late
Oligocene to early Miocene. These data suggest the influence of two distinct deep-water sources: cooler 680
southern component water and warmer northern component water (Billups et al., 2002; Pekar et al.,
2006; Liebrand et al., 2011). The observed carbonate-rich facies suggest an increased influence of
warmer northern component waters over the site at least in 13 occasions between 26 and 25 Ma.
4.4 Orbital forcing and Glacial and Interglacial cyclicity 685
The first spectral analysis on late Oligocene sediments from the eastern Wilkes Land margin at Site
U1356 shows that glacial-interglacial cycles, resulting in changes in the oceanic configuration off
Wilkes Land are paced with variations in Earth’s orbit and seasonal insolation. Although the data is
somewhat noisy due to gaps in our record, it clearly shows that the glacial-interglacial cyclicity (every 2
m or 41 kyr) discussed above has a persistent obliquity pacing throughout the studied late Oligocene 690
interval (26-25 Ma) in the Wilkes Land. Consequently, this obliquity-paced cyclicity modulates the
amount of deep-water production in the Southern Ocean, exerts a major control on oceanic
configuration and current strength. Bottom current velocity fluctuations and ventilation of bottom
sediments respond to the forcings applied by the strength of the Southern Hemisphere westerlies, the
Clim. Past Discuss., https://doi.org/10.5194/cp-2017-152
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21
position of the PF respect to the site, and consequently by the water mass occupying the bottom of the 695
basin at each time. In addition to obliquity, our record captures de precession frequencies, suggesting a
late Oligocene dynamic ice sheet.
East Antarctic ice volume fluctuations at orbital periodicities in the obliquity band in the Wilkes Land
margin have been previously reported from early warm Pliocene (3-5 Ma) sediments obtained from Site 700
U1361 (Patterson et al., 2014). In the Ross Sea, cyclicity in sediments collected by the CRP from the
late Oligocene, the late Miocene and the early warm Pliocene period was also paced by obliquity (Naish
et al., 2001; McKay et al., 2009; Naish et al., 2009). Similar orbital variability in the deep-water
circulation patterns have also been inferred to have occurred with the growth of the EAIS during the
middle Miocene between 15.5 to 12.5 Ma (Hall et al., 2003). In addition, other studies have linked 705
changes in Atlantic meridional overturning (Lisiecki et al., 2008; Scher et al., 2015) and Antarctic
circumpolar ocean circulation (Toggweiler et al., 2008) to obliquity forcing. An interglacial mechanism
has been proposed whereby the southward expansion of westerly winds and associated Ekman transport
is compensated for by enhanced upwelling of warmer, CO2-rich CDW (Toggweiler et al., 2008), which
also promotes atmospheric warming. In the equatorial Pacific, Pälike et al. (2006) also report strong 710
obliquity in the benthic δ13C isotopic record between 26-25 Myr, implying that changes in the carbon
cycle (pacing glacial /interglacial periods) are triggered in the high southern latitudes and transferred to
the global deep-ocean through the bottom water masses.
Conclusions
Our study provides new insights regarding Antarctic ice sheet and paleoceanographic configurations 715
that prevailed in the eastern Wilkes Land margin between 26 and 25 Ma. Sediments at IODP Site
U1356 during this interval are characterized by the alternation between two main facies (F1 and F2),
that are dominated by reworking of glacial-interglacial deposits by bottom-currents with varying
intensities. Claystones with silty laminations (F1) are interpreted to represent fluctuating bottom current
intensities during glacial periods. Massive bioturbated silty clays and micritic limestones (F2) are 720
interpreted as interglacial deposits and record maximum velocities of bottom-currents at this site. The
lack of iceberg rafted debris (IRD) and sea ice and warm sea surface temperatures indicators throughout
the studied interval contrasts with early Oligocene and younger sections from Site U1356 and with late
Oligocene sediments from the Ross Sea (CRP and DSDP 270), which contain IRD and evidence for sea
ice and ice at or near the coast. Based on these observations, we postulate that reduced glaciers or ice 725
caps occupied the lowlands of the Wilkes Subglacial Basin between 26 and 25 Ma and that iceberg
calving was only a background process during this time due to the lack of marine terminating ice sheets,
with ablation of the ice sheet was largely controlled by melt processes rather than iceberg calving.
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22
Glacial sediments record poorly ventilated, low-oxygenation conditions at the water-sediment interface 730
that we postulate result when westerly winds and surface oceanic fronts migrate towards the equator
and overturning was reduced near the Antarctic margin. During interglacial times, more oxygenated and
better-ventilated conditions are inferred to have prevailed which would act to enhance mixing of the
water masses with increased current velocities. We postulate that during interglacials, westerly winds
shifted south and became more aligned with the Antarctic Divergence and Polar Fronts, which 735
promoted upwelling of deep waters and facilitated the mixing and oxygenation of bottom waters.
Micritic limestone intervals within interglacial F2, record warmer paleoclimatic conditions when a
greater relative proportion of warm north component waters reached the site allowing the preservation
of carbonate. Preservation of carbonate in some F2 intervals supports previous paleoceanographic
studies that consider at least a two-layer ocean with an Antarctic Bottom Water (undersaturated with 740
respect to calcium carbonate), and a warmer Northern Component Water (NADW-like) to reconcile
intra-basinal differences in δ18O values (Pekar et al., 2006). Based on the number of carbonate-rich
layers, warmer NADW-like waters reached the site at least 13 times during the studied interval.
Spectral analysis on late Oligocene sediments from the eastern Wilkes Land margin reveal that glacial-745
interglacial paleoceanographic changes during the late Oligocene are regulated primarily by obliquity,
although frequencies in the eccentricity and precession band are also recorded. However, as we do not
have a measure of ice dynamics during this time (e.g. ice rafted debris), the orbital response of
terrestrial ice remains ambiguous, beyond what is inferred from the deep-sea isotope record.
750
Our record shows that during under the high CO2 values of the late Oligocene (i.e., from ~750 ppm to
400 ppm), ice sheets had retreated to their terrestrial margins, with ice sheet mass loss dominated by
surface melt processes. It also indicates a slowdown of the southern limb of overturning circulation,
with the increased presence of North Component Deep waters influencing the preservation of carbonate
in this sector of the eastern Wilkes Land margin. 755
Acknowledgements
This research used samples and data provided by the Integrated Ocean Drilling Program (IODP). The IODP is sponsored by
the US National Science Foundation (NSF) and participating countries under the management of Joint Oceanographic 760
Institutions, Inc. Funding for this research is provided by the Spanish Ministerio de Economía y Competitividad (Grant CTM
2011-24079 and CTM2014-60451-C2-1-P) co-funded by the European Union through FEDER funds and the Deutsche
Forschungsgemeinschaft (DFG) (RO 1113/6 to UR). We thank the staff onboard Exp. 318, at GCR and BCR for assistance
in core shipping and handling, and Vera Lukies (MARUM) for technical support with XRF core scanning. PKB, FS and
JDH acknowledges funding through NWO polar programme grant no 866.10.110; PKB acknowledges funding through 765
Clim. Past Discuss., https://doi.org/10.5194/cp-2017-152
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Discussion started: 5 December 2017
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23
NWO-VENI grant no 863.13.002. US acknowledges funding received from the Natural Environment Research Council
(NERC grant NE/H000984/1).
Author contributions 770
CE and AS designed the research. PKB, JH, FS and HB, provided insights regarding biomarker-based sea surface
temperatures and sea ice conditions based on dinocysts. UR provided XRF core-scanning data. FJJE and UR provided
geochemical input. CHN and RM provided input with sedimentary and facies interpretations. MI provided the CT-Scans
data. JAF provided input in the paleoceanographic interpretations. DE and ALQ provided Antarctic overview and
petrographic input. SR and US provided palynology insights. AS integrated, cross-validated and compiled the data, and 775
wrote the paper with input from all co-authors.
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Tables 1-3:
1170
Table 1: Age model by Tauxe et al., (2012) and transformed ages to GPTS 2012
Core Section
Site U1356
Exp. 318
Top depth
(mbsf)
Bottom depth
(mbsf)
Depth used (m)
GPTS 2004
(Myr)
(Tauxe et al.,
2012)
GPTS 2012
(Myr)
Chron
68R-2
643.10
643.65
643.37
25.444
25.260
C8n.1n (o)
69R-2
652.55
652.60
652.57
25.492
25.300
C8n.2n (y)
71R-6
678.06
679.90
678.98
26.154
25.990
C8n.2n (o)
Table 2: Types of facies differentiated by physical, geochemical, and biological character and their interpretation in
terms of sedimentary processes and paleoclimate.!1175
!
Facies 2 (F2)
Lithological description
Bioturbated green claystones with thin silt
laminae with planar and cross-bedded
laminations
Highly bioturbated, thicker pale-brown, silty-
claystones
Contacts
Top
Gradual, bioturbated
Sharp
Bottom
Sharp
Gradual, bioturbated
Bioturbation
Sparse bioturbation. Primary structures
preserved
Strong bioturbated. Massive. No primary
structures preserved
Nannos
Barren to rare
Barren to variable abundance and preservation
IRD
No
No
Magnetic susceptibility
(MS)
Low in claystones and high in silty
laminations
High
XRF-Scanner
elements
concentration
Zr
Low in claystones and high in silty
laminations
High, (max. values on top)
Ba
High, (max. values on bottom)
Low
Ca
No
Variable, low to high
Formation process
Bottom currents of fluctuating intensities
Bottom currents with higher velocity and
constant flux
Facies interpretation
Cold periods. Supply of terrigenous by
density current flows, reworked by bottom
currents.
Well-oxygenated deep-sea sedimentation.
Warm periods with reworking of sediments by
bottom currents
!
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31
Table 3: R Pearson Linear correlation between XRF-scanner elements. 1180
MS
S
Ca
Ti
Mn
Fe
Br
Rb
Zr
Sr
S
-0.214
Ca
0.226
-0.122
Ti
-0.212
0.620
-0.290
Mn
0.151
-0.121
0.858
-0.246
Fe
0.0419
0.0449
-0.396
0.510
-0.324
Br
-0.297
0.111
-0.438
0.118
-0.363
0.056
Rb
-0.282
0.036
-0.576
0.286
-0.489
0.455
0.493
Zr
0.480
-0.164
-0.036
-0.099
-0.058
-0.055
0.102
0.067
Sr
0.186
0.006
0.871
-0.074
0.677
-0.345
-0.303
-0.515
0.040
Ba
-0.290
0.339
-0.234
0.662
-0.210
0.354
0.343
0.402
0.018
0.039
!
!
1185
1190
1195
1200
1205
1210
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32
121
5
122
0
122
5
123
0
123
5
124
0
Fig. 1: Location of IODP 318 Site U1356 (Escutia et al., 2010) on the Adélie coast continental rise. Bed
124
5
topography from IBSCO2 (Arndt, JE et al., 2013). Schematic position of the different water masses at present
and locations of Antarctic Bottom Water formation (Orsi, 1995) are indicated.) The position of the Oligocene
Polar Front (Scher et al., 2015) is also shown. ASF: Antarctic Slope Front; SB: Southern Boundary; SACCF:
Southern Antarctic Counter Current Front; ARB: Alie Rift Block.
125
0
O
l
i
g
o
c
e
n
e
P
o
l
a
r
F
ro
n
t
U135 6
P
o
l
a
r
F
r
o
n
t
a
l
Z
o
n
e
U135 6
Dibble
Glacier
Ross
Ice Shelf
160º 15 14
130º E
60º S
70º S
WILKES LANDWILKES LAND
ADÉLIE
COAST
ADÉLIE
COAST
WILKES
SUBGLACIAL BASIN
WILKES
SUBGLACIAL BASIN
ROSS SEA
GYRE
SACCF
SB
ASF
0500
Km
ARB
DSDP Site 269
DSDP Site 270 CRP
IODP 318 Site
Polar Fronts
Bottom water
formation
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33
125
5
126
0
126
5
127
0
127
5
128
0
128
5
Fig. 2: Magnetic susceptibility (MS), natural gamma radiation (NGR), and selected X -Ray Fluorescence (XRF)
data (in total counts) and elemental rations plotted plot against the new detailed U1356 facies log between 689.4
and 641.4 mbsf.
129
0
Bioturbated
Brown silty-calystone
Green claystone with thin
Silt laminae
10
0
20
30
40
50
60
70
80
90
100
110
120
130
140
150
160
170
180
190
200
210
220
230
240
250
260
270
280
290
300
310
320
330
340
350
360
370
380
390
610
620
630
640
650
660
670
680
690
700
710
720
730
740
750
760
770
780
400
410
420
430
440
450
460
470
480
490
500
510
520
530
540
550
560
570
580
590
600
^ĞĐƟŽŶϮ^ĞĐƟŽŶϯ
^ĞĐƟŽŶ
^ĞĐƟŽŶϭ^ĞĐƟŽŶϰ^ĞĐƟŽŶϱ^ĞĐƟŽŶϲ
25,444 Ma
PM
68
10
0
20
30
40
50
60
70
80
90
100
110
120
130
140
150
160
170
180
190
200
210
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240
250
260
270
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340
350
360
370
380
390
610
620
630
640
650
660
670
680
690
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710
720
730
740
750
760
770
780
790
800
400
410
420
430
440
450
460
470
480
490
500
510
520
530
540
550
560
570
580
590
600
^ĞĐƟŽŶϮ^ĞĐƟŽŶϯ
^ĞĐƟŽŶ
^ĞĐƟŽŶϭ^ĞĐƟŽŶϰ^ĞĐƟŽŶϱ^ĞĐƟŽŶϲ
69
10
0
20
30
40
50
60
70
80
90
100
110
120
130
140
150
160
170
180
190
200
210
220
230
240
250
260
270
280
290
300
310
320
330
340
350
360
370
380
390
400
410
420
430
440
450
460
470
480
490
500
510
520
530
540
^ĞĐƟŽŶϮ^ĞĐƟŽŶϯ
^ĞĐƟŽŶ
^ĞĐƟŽŶϭ^ĞĐƟŽŶϰ
25,492 Ma
PM
70
10
0
20
30
40
50
60
70
80
90
100
110
120
130
140
150
160
170
180
190
200
210
220
230
240
250
260
270
280
290
300
310
320
330
340
350
360
370
380
390
610
620
630
640
650
660
670
680
690
700
710
720
730
740
750
760
770
780
790
800
400
410
420
430
440
450
460
470
480
490
500
510
520
530
540
550
560
570
580
590
600
^ĞĐƟŽŶϮ^ĞĐƟŽŶϯ
^ĞĐƟŽŶ
^ĞĐƟŽŶϭ^ĞĐƟŽŶϰ^ĞĐƟŽŶϱ^ĞĐƟŽŶϲ
71
10
0
20
30
40
50
60
70
80
90
100
110
120
130
140
150
160
170
180
190
200
210
220
230
240
250
260
270
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330
340
350
360
370
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390
610
620
630
640
650
660
670
680
690
700
710
720
730
740
750
760
770
780
790
800
400
410
420
430
440
450
460
470
480
490
500
510
520
530
540
550
560
570
580
590
600
^ĞĐƟŽŶϮ^ĞĐƟŽŶϯ
^ĞĐƟŽŶ
^ĞĐƟŽŶϭ^ĞĐƟŽŶϰ^ĞĐƟŽŶϱ^ĞĐƟŽŶϲ
72
Core #
Facies Log
Zr/Ba
13
Zr/Ti Ba/Ti
Br/Ti
40
Facies 1 (F1) No recovery
Facies 2 (F2)/F2 with CaCO3
25.26 Ma
25.99 Ma
Paleomag.
Tie points
25.3 Ma
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34
129
5
130
0
130
5
131
0
131
5
Fig. 3: Detailed images, CT-scans and HRSEM from Facies 1 (F1) and Facies 2 (F2). (a) Example of F1 taken
from core 71R-2 119/146 cm, showing faint laminations (fl) and bioturbation by Planolite s (p) (b) CT-scan 3D
132
0
image of the same core interval, note the pyritized burrows (py). (c) Example of F2 taken from core 72R -1 18/53
cm). (d-f) Close-ups of laminations from Facies 1: ripples (r), planar lamination (pl), and faint laminations (fl),
with mud offshoots (mo). (d) Chondrite s (Ch) bioturbation inside F1. (g) HRSEM image of F1 (68R-4-86/88 cm)
with detritic aspect and a mudstone clay matrix, Quartz grains (Qz), diagenetic calcite (arrows), and dissolved
coccoliths (circles); (h) HRSEM image of F2 (71R-2 140/142 cm) silt sized matrix and reworked calcareous
132
5
nannofossils, and conchoidal quartz grain (C-Qz); (i) Detail of dissolved coccoliths and diagenetic calcite
mineral; (j) Detail of a dissolved and reworked calcareous nannofossils and a fractured conchoidal quartz (C-Qz).
133
0
(a) (c)(b)
Density
+-
71R-2 119/146 cm
FACIES 2 (F2) FACIES 1 (F1)
FACIES 2 (F2) FACIES 1 (F1)
CT-SCAN
IMAGE
CORE
PHOTO
CORE
PHOTO
72R-1 18/53 cm
72R-5 47 cm
(e)
(d) 71R-CC 12/19 cm
(f) 69R-2 58/67 cm
Ch
p
p
(g) 68R-4 86/88 cm F1
C-Qz
Qz
Qz
Qz
(i)
Clay
Clay
(h) 71R-2 140/142 cm F2
30 ȝm
40 ȝm
Clay
pl
pl
py
fl
r
mo
fl
fl
Facies 1 (F1)
Facies 2 (F2)
Dissolved nannos
Diagenetic calcite 30 ȝm
(j) 71R-5 70/72 cm F2
C-Qz
(i) 68R-4 86/88 cm F1
10 ȝm
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35
133
5
134
0
134
5
135
0
135
5
136
0
Fig. 4: Detailed Facies characterization of two representative sections using: (a) Interpreted facies F1 and F2; a
high-resolution digital image of the core sections (b), facies log (c), Magnetic susceptibility (MS) (d), XRF Zr/Ba
136
5
ratio (e), and XRF calcium counts (f).
MSt Sd G
684.0
684.4
684.8
685.2
Bigradational grading
Facies 1 - GLACIALFacies 1 - GLACIAL Facies 2 - INTERGLACIAL
Depth
(mbsf)
U1356
72R-4 Zr/Ba Zr/Ba
MSt Sd G
652.4
652.8
653.2
653.6
Facies 1 - GLACIAL GLACIAL Facies 2 - INTERGLACIALIG
Depth
(mbsf)
U1356
69R-2
Inverse Bigradational grading
grading
(a) (b) (c) (d) (e) (a) (b) (c) (d) (e)(f) (f)
10
CaCO3 est. (%)
Ca (total counts)
MS
(instr. units)
Bioturbated
Brown silty-calystone
Facies 1 (F1) - GLACIAL
Facies 2 (F2) - INTERGLACIAL
Green claystone with thin
Silt laminae
Silt laminae, with cross
and/or parallel laminations
Peak current
velocity
CaCO3 est. (%)
Ca (total counts)
MS
(instr. units)
6
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36
137
0
137
5
138
0
138
5
139
0
139
5
140
0
Fig. 5: Linear correlation between CaO% (discrete XRF) and ln(Ca/Ti) (XRF scanner) values in order to estimate
carbonate contents (CaCO3 est. %).
140
5
y = 3.9003x - 0.9449
R = 0.81448
-2
0
2
4
6
8
10
12
14
1 2 3
CaCO3 estimates (%)
CaCO3 est. (%)
Lineal (CaO (%))
ln(Ca/Ti)
CaO (%)
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37
141
0
141
5
142
0
142
5
143
0
143
5
Fig. 6: Spectral analysis results of the Zr/Ba obliquity tuned and anchored data. (a) Zr/Ba ratio tuned with
144
0
Astrochron (Meyers, 2014) and anchored to the top of the C8n.2n (o) chron. (b) EHA and (c) MTM spectral
analysis on Zr/Ba tuned data. EHA normalized power with 300-kyr window with 3DPSS tapers. (d) EHA
amplitude for the eccentricity-obliquity-precession ETP solution (Laskar et al., 2004) calculated for the same
period of time with with 3DPSS tapers and 200-kyr window.
144
5
0.00 0.02 0.04 0.06
26000 25800 25600 25400
0.0 0.5 1.0 1.5 2.0
0.00 0.01 0.02 0.03 0.04 0.05 0.01 0.03 0.050.06
26000 25900 25800 25700 25600 25500 25400
0.0 0.2 0.4 0.6 0.8 1.0
Time (Kyr)
Frequency (cycles/Kyr) Frequency (cycles/Kyr)
0.5 2.0
26200 26000 25800 25600 25400 25200
(a)
(b)
(c)
(d)
Anchored
Zr/Ba
Normalized power for the
anchored U1356 Zr/Ba data
Amplitude for the
eccentricity-obliquity-precession
ETP solution
100
Kyr
400
Kyr
40
Kyr
23
Kyr
0.00 0.01 0.02 0.03 0.04 0.05 0.06
0.0000 0.0008
MTM P ow er (bl ac k); AR 1 t (r ed) ; 9 0%CL , 95%CL , 99%C L (dotte d)
Frequency
Linear Power
O1
O1
O1
P
P1
E1?
E2-3
E
E2-3
E2-3
P2
O1
O1
O1
P
P1
E1?
E2-3
E
E2-3
E2-3
P2
P
P
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38
Fig. 7: Paleoceanographic reconstructions based on our interpretations for Facies 1 and 2. (a) Modelled ice
thickness for the mid-Miocene ice sheet by Gasson et al., (2016). (b) Glacial periods with low obliquity 1450
configuration. Westerlies and Polar Front (PF) move northwards. There is enhanced proto-AABW formation.
Low ventilation conditions occur at the ocean/sediment interface and mixing of waters masses is diminished.
Bottom currents are weak and fluctuating, producing laminated sediments. (b) Interglacials occur during high
obliquity configuration. Westerlies and the PF move southwards, close to the Site U1356. Proto-AABW
formation is reduced. Intrusions of CDW/NADW-like reach southernmost positions. (c) During warm 1455
Interglacials, NADW-like is enhanced and CaCO3 sedimentation is more abundant. (b,c) Bottom water
ventilation and upwelling are more vigorous, with stronger bottom currents that result in fully bioturbated and
silty-sized sediments.
Westerlies migrate
North
Glacials
0
1
2
Depth (km)
S N
CO2
DSDP
269A
Site U1356
Actual position
Late Oligocene paleoposition 61º S-60º S63º S -61º S
61º S-60º S63º S -61º S
61º S-60º S63º S -61º S
Wilkes
Land
Adélie
Rift Block
CDW
e. Oligocene
AD/PF
formation
Weste rlie s
South
S N
DSDP
269A
0
1
2
Depth (km)
Warm Interglacials
CO2
Site U1356
Actual position
Late Oligocene paleoposition
Wilkes
Land
Adélie
Rift Block
CDW NADW
AD/PF ?
Weste rlie s
South
0
1
2
Depth (km)
S N
Interglacials
CO2
DSDP
269A
Site U1356
Actual position
Late Oligocene paleoposition
Wilkes
Land
Adélie
Rift Block
CDW NADW
AD/PF ?
AABW
Mixing intensity
Glacial drift
Interglacial drift
Warm inte rgla cial
drift
Upwelling
Air-sea
CO2
exchange
CDW / NADW
Productivity
(a) (b)
(c)
(d)
U1356
U1356
Paleoceanographic conguration of Wilkes Land region during the
Late Warm Oligocene (~26-25 Ma)
Gasson et al., 2016
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