Magnetostratigraphic Chronology of a Cenozoic Sequence From DSDP Site 274, Ross Sea, Antarctica

New paleomagnetic results from the late Eocene-Middle Miocene samples from Deep Sea Drilling Project Site 274, cored during Leg 28 on the continental rise off Victoria Land, Ross Sea, provide a chronostratigraphic framework for an existing paleoclimate archive during a key period of Antarctic climate and ice sheet evolution. Based on this new age model, the cored late Eocene-Middle Miocene sequence covers an interval of almost 20 Myr (from ∼35 to ∼15 Ma). Biostratigraphic constraints allow a number of possible correlations with the Geomagnetic Polarity Time Scale. Regardless of correlation, average interval sediment accumulation rates above 260 mbsf are ∼6 cm/kyr with the record punctuated by a number of unconformities. Below 260 mbsf (across the Eocene/Oligocene boundary) interval, sedimentation accumulation rates are closer to ∼1 cm/kyr. A major unconformity identified at ∼180 mbsf represents at least 9 Myr accounting for the late Oligocene and Early Miocene and represent non-deposition and/or erosion due to intensification of Antarctic Circumpolar Current activity. Significant fluctuations in grain size and magnetic properties observed above the unconformity at 180 mbsf, in the Early Miocene portion of this sedimentary record, reflect cyclical behavior in glacial advance and retreat from the continent. Similar glacial cyclicity has already been identified in other Miocene sequences recovered in drill cores from the Antarctic margin.


INTRODUCTION
Deep Sea Drilling Program (DSDP) Leg 28 ( Figure 1) was designed to explore the long-term climatic, biostratigraphic and geological history of Antarctica and its environments . Such geological records provide insights into modern and future climate sensitivity estimates, particularly for time periods characterized by the presence of continental ice sheets and a paleogeography similar to modern (e.g., Markwick, 2007;Farnsworth et al., 2019).
Earth's climate underwent a stepwise shift from greenhouse to icehouse conditions during the Cenozoic. Major ice sheets first appeared on Antarctica across the Eocene/Oligocene boundary coincident with the earliest Oligocene Oi-1 oxygen isotope event (1.0‰ δ 18 O increase at ca. 33.55 Ma; e.g., Miller et al., 1991;Zachos et al., 1996;Zachos et al., 2001;Miller, 2005a;Miller et al., 2005b;Francis et al., 2009;Florindo, 2017, Westerhold et al., 2020). While long thought to be associated with the early glaciation of Antarctica (Kennett, 1977), the Antarctic Circumpolar Current did not appear until the Late Oligocene (Lyle, et al., 2007). Throughout the Oligocene, Antarctic ice sheets experienced repeated cycles of advance and retreat, as indicated by large and rapid changes in global sea level (Kominz and Pekar, 2001) and δ 18 O variations in deep-ocean sediment (Pälike et al., 2006). Following the inception of the Antarctic Circumpolar Current, the Oligocene terminated with another transient global cooling event (Mi-1, at ca. 23 Ma), characterized by a ∼1‰ positive excursion in marine benthic foraminiferal δ 18 O records (e.g., Wilson et al., 2009;Beddow et al., 2016) and associated with a large-scale Antarctic ice sheet expansion at the Oligocene-Miocene transition. This event was dramatic in scale, with the East Antarctic Ice Sheet (EAIS) estimated to have grown from 50 to 125% of its present-day size (e.g., Pekar et al., 2006). It was the first of a series of similar "Mi-" oxygen isotope events that, while less extreme than Mi-1, were responsible for the progressive cooling and glaciation experienced by Antarctica during the Miocene. As in the Oligocene, however, the cooling during the Miocene was not monotonic, but characterized by swings between cooler and warmer climate (e.g., Miller et al., 1991;Zachos et al., 1997;Zachos et al., 2001;Lear et al., 2004;Holbourn et al., 2005;Shevenell et al., 2008;Holbourn et al., 2013;Holbourn et al., 2014;Holbourn et al., 2015;Lear et al., 2015;Holbourn et al., 2018;Liebrand et al., 2017;Westerhold et al., 2020).
After the initial onset of and recovery from Mi-1, the Early Miocene experienced relatively warm intervals. The warming cycles of the Early Miocene culminated in the Mid-Miocene Climatic Optimum (MMCO) at around 17-14 Ma, immediately followed by the Middle Miocene Climate Transition (MMCT) at 14.2-13.8 Ma (Shevenell et al., 2004). The transition, as recorded in deep-ocean δ 18 O variations, was a significant one: the MMCO was the warmest interval since the Eocene, while the MMCT, and the subsequent cooling during the rest of the Miocene, produced temperatures colder than any previously observed in the Cenozoic. During this time, there FIGURE 1 | Map of Ross Sea with schematic representation of present-day Antarctic Slope Current (dashed white line) and the continental shelf current flow (orange lines) (Ainley and Jacobs, 1981;Whitworth et al., 1995;Orsi and Wiederwohl, 2009). Location of DSDP Site 274 is indicated by a red star. DSDP 270, AND-2A, Eltanin 27-21 and CRP are also shown because they are reference cores in the area (respectively, Kulhanek et al., 2019Jovane et al., 2008;Roberts et al., 2013).
was a sizable EAIS, as inferred from δ 18 O, sea level curves and direct sedimentological evidence from the Antarctic continent itself (e.g., Lewis et al., 2007).
In this paper, we present the first magnetostratigraphic chronology of a late Eocene-Early Miocene record from a sediment core (DSDP Leg 28, Site 274) obtained from offshore Cape Adare in the Ross Sea Sector of the Antarctic continent in 1973. The new age model will allow further studies of this unique core to produce future paleoenvironmental reconstructions focusing on the discharge of sediments from the catchment areas in this region of the Ross Sea (e.g., Sagnotti, et al., 1998;Jovane et al., 2004;Jovane et al., 2019;Roberts et al., 2013). A graben-like structure, with relief of several hundred meters (Houtz and Meijer, 1970), lies to the southwest of the site and probably served as an effective barrier to the downslope transport of terrigenous material. Seafloor magnetic lineations in the area and detailed biostratigraphic studies suggested a late Eocene/ Early Oligocene age for the base of the sedimentary cover . However, paleomagnetic analysis has not been carried out on those cores until this study.

SITE LOCATION AND LITHOLOGY
For this study, half-core samples were obtained from the International Ocean Discovery Program (IODP) Gulf Coast Repository from 320 m below sea floor (mbsf) (close to the contact with basement) to 123 mbsf. This interval comprises silty claystone, diatom-detrital silty clay ooze and diatom-rich silty clay ). An unconformity was identified at 180.5 mbsf by Frakes (1975) and interpreted to represent a time interval between late Oligocene (Core 20) and Early-Middle Miocene (Core 19).

MATERIAL AND METHODS
For the present study, DSDP Site 274 working half core samples were provided by the IODP Gulf Coast Repository from which we collected 8 cm 3 cubic samples, using a non-magnetic double blade saw. The core was dry but preservation was excellent.
Two hundred and sixty samples were taken from the center of the working halves between 316 mbsf (Core 34) and 123 mbsf (Core 14), with a sample collected at least every meter (depending on the core recovery, drilling-induced disturbance and where sufficient material was available). Sediment samples from the interval between 320 and 180 mbsf are stiffer and more consolidated, while those from the interval between 180 and 123 mbsf are softer and less consolidated. Samples were oriented only with respect to vertical direction, therefore absolute paleomagnetic declinations could not be determined. Paleomagnetic analyses were performed at Instituto de Astronomia, Geofísica e Ciências Atmosféricas da Universidade de São Paulo (IAG USP) in a magnetically shielded room using a 2 G Enterprises magnetometer (Model 755-4K). Grain size, magnetic susceptibility and rock magnetic analyses were conducted at the Centro Oceanográfico de Registros Estratigráficos (CORE) at the Instituto Oceanográfico da Universidade de São Paulo (IO USP) using a Microtrac Bluewave-SIA, an AGICO MFK1-FA Kappabridge, and a Vibrating Sample Magnetometer (VSM) MicroMag 3,900 Princeton-Lake Shore Cryotronics, respectively.

Grain Size
Samples were prepared by crushing ∼0.2 g of bulk sediment in an agate mortar. The crushed samples were further disaggregated using 10% hydrogen peroxide to remove organic matter. After 48 h of reaction, the samples were passed through 2 mm sieves and dried in an oven. After drying, sodium pyrophosphate was added to prevent the particles from flocculating. The Microtrac Bluewave equipment was used to measure the particles by laser diffraction in water, with their size being calculated according to the modified Mie theory for non-spherical particles (Chýlek et al., 1976). To establish the basic statistical parameters of particle size, as well as particle size distribution, the GRADISTAT v. 4.0 program, developed by Blott and Pye (2001), was used to establish the basic statistical parameters of particle size, as well as particle size distribution.

Mineral Magnetic Properties
Rock magnetic analyses were conducted on 21 selected samples distributed along the studied interval to estimate downcore variations in the composition, concentration, and grain size of magnetic minerals. These analyses included volume normalized low-field magnetic susceptibility (κ), hysteresis loops, with maximum fields of 0.5 T, isothermal remanent magnetization (IRM) imparted in progressively (100 acquisition steps) to saturation (SIRM), or to the maximum attainable direct field of 1.8 T. IRM curves were subdivided into discrete coercivity components, as in Robertson and France (1994), with cumulative log-Gaussian (CLG) functions, using the spreadsheet developed by Kruiver et al. (2001) and Heslop et al. (2002).
In addition, we investigated the temperature dependence of magnetic susceptibility using a MFK1-FA equipped with CS4 furnace (AGICO) Kappabridge, to discriminate ferromagnetic mineralogy. Samples were measured up to a maximum temperature of 700°C in an argon atmosphere, in order to minimize thermochemical alteration.

Grain Size
The lithology of the studied interval is predominantly clay. The average grain size is 11.16 phi (0.435 μm), while the mode is 11.17 phi (0.432 μm), with a standard deviation of 0.33. The coarsest average grain size (9.87 phi; 1.065 μm) is found at 158.09 mbsf. The finest average grain size (12.41 phi; 0.181 μm) was observed at 162.11 mbsf. Several samples have multiple peaks in the grain size distribution (Figure 2). A drop-stone, 2 cm in diameter, was found at 319.74 mbsf, and additional ice rafted debris (IRD) particles greater than 5 mm in diameter were identified at 258.5 and 247.5 mbsf. Frakes (1975) also reported the presence of a coarse fraction in Cores 19 and 18 (180.5 and 161.5 mbsf, respectively) interpreted to reflect ice rafting events.
We observed a lower proportion of clay and a higher proportion of sand in the interval between 180 and 130 mbsf (Cores 19-14, respectively), where the variability of the average grain size is greater and ranges between 10 and 12 phi (0.98-0.24 μm). Between 180 and 160 mbsf (Cores 19-18, respectively) the average grain size is finer between 11 and 12 phi (0.48-0.24 μm) and coarsens between 160 and 130 mbsf  to values ranging between 9.7 and 11.7 phi (1.20-0.30 μm). The average grain size is finer in the lower interval of the core (between ∼300 and ∼180 mbsf; Cores 34-20) with values ranging between 11 and 11.5 phi (0.48-0.34 μm), with a slight coarsening in average grain size in the interval ∼320-∼304 mbsf (Cores 34-33). Abrupt changes in grain size at ∼262, ∼240 and ∼224 mbsf suggest the possibility of minor unconformities in the sedimentary succession.

Magnetic Properties
Magnetic susceptibility (κ) values range widely between 1.8 × 10 −6 and 92.9 × 10 −6 SI with a mean value of 10.9 × 10 −6 SI ( Figure 2). Below the unconformity at ∼180 mbsf, the highest κ values occur in samples between 280 and 210 mbsf and are associated with IRD intervals with coarser sediments and larger clasts. Above 180 mbsf, κ gradually increases reaching peak values at 140 mbsf.
IRM acquisition curves obtained for the samples were separated into four coercivity components ( Figure 3). Both samples (at 218.92 and 285.41 mbsf) show a component with . Those components are assumed to represent low-coercivity SD magnetite and high-coercivity SD magnetite, probably of biogenic origin (Egli, 2004;Jovane et al., 2012;Yamazaki , 2012;Heslop et al., 2004;Rodelli et al., 2019). The component characterized by mean coercivities of 100 and 280 mT (dispersion parameter 0.15 and 20, respectively), are interpreted as hematite (Kruiver et al., 2001;Abrajevitch et al., 2009;Rodelli et al., 2019). The overall contribution to the IRM in sample 218.92 mbsf from the SD magnetite is 72, and 9% is linked instead to detrital magnetite. Hematite only represents 19% of the total contribution. IRM acquisition curves for a sample at 285.41 mbsf, typical of samples with lower κ, was fitted with only four components ( Figure 3E): the low coercivity SD magnetite component is missing, and the hard SD magnetite is less abundant, corresponding to only 64% of the magnetic signal, and the detrital magnetite is more abundant (18%). Hematite is slightly less abundant (18%). A κ-T curve for sample at 218.92 mbsf ( Figure 3C) shows increasing values up to about 330°C, which can be interpreted as maghemite (Rodelli et al., 2018;Rodelli et al., 2019), possibly related to partial oxidation of magnetite. This is followed by a peak at about 520°C and a major inflection at 579°C, calculated with the two-tangent method of Grommé et al. (1969), and consistent with chemically pure magnetite (Thompson and Oldfield, 1986). A κ-T curve for sample at 285.41 mbsf ( Figure 3F) has a sharp drop at 565°C consistent with magnetite probably mixed with titano-magnetite (Ozima and Larson, 1970). The cooling curve for sample at 285.41 mbsf has higher overall κ than the heating curve, which indicates production of new magnetic phases during heating while sample at 218.92 mbsf maintain similar values.

Demagnetization Behavior
NRM intensities (Figure 2A) for the DSDP 274 succession range widely between 9.36 × 10 −2 and 6.58 × 10 −5 A/m about a mean of 4.72 × 10 −3 A/m. At the base of the sequence (320-310 mbsf), values are higher than for the rest of the sequence below the unconformity, then increase again above 180 mbsf. Glaciomarine sediments at the base are coarser and contain larger clasts that can increase intensities and lead to non-reliable paleomagnetic behavior (e.g., unstable remanence, low inclination or multi-component magnetizations). For these reasons, samples at the base were excluded from magnetostratigraphic interpretations.
Some samples exhibit a low coercivity horizontal component of remanence ( Figure 4) interpreted to be a drilling induced radial overprint related to strong magnetic fields in the cuttingshoe and/or core barrel (e.g., Acton et al., 2002;Florindo et al., 2003). In most cases, this is removed by AF demagnetization in a peak field of 20 mT, but in some cases the horizontal overprint is harder and persists to higher coercivities (>80 mT; Figure 4D) Frontiers in Earth Science | www.frontiersin.org December 2020 | Volume 8 | Article 563453 6 and prevents the isolation of a ChRM. In a few samples, characterized by the lowest NRM intensities, this overprint could not be removed by AF demagnetization. We cannot rule out that drying may have further hardened the radial overprint.
The ChRM is identified in most samples, regardless of polarity, between AFs of 20 and 70 mT (Figure 4). Approximately half of the ChRMs are directed to the origin of vector component plots ( Figures 4A,C) and half follow a trajectory toward a high coercivity component not removed by peak AFs of 100 mT ( Figures 4F,G). The high coercivity component of samples containing a small amount of hematite is assumed to be a secondary chemical remanent magnetism and unanchored ChRM components are reported.

Magnetozones
Of the total of 287 samples, 65.51% (188) were used to determine the magnetic polarity zonation, with discrete magnetozones defined as an interval where at least two consecutive samples show polarities different from adjacent intervals. The other 99 samples were discarded according to the following criteria. Samples were discarded from the final analysis if they met the following conditions: 1) samples with a radial overprint ( Figure 4D); 2) single samples showing a ChRM with an inclination of opposite polarity to that of the previous and following samples; 3) samples with anomalous low inclination (shallower than ±25°) were also not considered ( Figure 4G); 4) samples containing IRD which do not demagnetize due to the presence of high-coercivity minerals and 5) samples with overprints (drilling overprint (Fuller et al., 1998) or coresplitting overprint (Wilson et al., 2000;Florindo et al., 2001) that could not be removed by low field AF demagnetization.
Mean ChRM inclinations were computed separately for normal and reverse polarity populations using a maximum likelihood method (Arason and Levi, 2010) as implemented in the PuffinPlot application (Lurcock and Florindo, 2019). The normal polarity intervals have a mean inclination of −55.01°(N 56, α 95 6.67°, θ 63 27.03°), the reverse polarity intervals have a mean of 68.3°(N 132, α95 4.45°, θ 63 27.8°) ( Figure 5). Some samples have lower inclination than the expected Geocentric Axial Dipole (GAD) inclination (±64.5°at 20 Ma and ±63.9°at 40 Ma; Torsvik et al., 2012) for this period and for this site latitude. A post depositional tilt of ∼5°the strata is FIGURE 5 | Histograms of paleomagnetic inclinations from discrete samples for DSDP Site 274 after AF demagnetization. Mean ChRM inclinations (shown in blue), and associated α 95 , were computed separately for normal and reverse polarity populations using a maximum likelihood method (Arason and Levi, 2010). The expected inclinations for the GAD are shown in red for this period and for the site latitude (Torsvik et al., 2012).
Frontiers in Earth Science | www.frontiersin.org December 2020 | Volume 8 | Article 563453 8 implied by the difference in normal and reversed polarity inclinations.
We recognize eight normal and eight reversed magnetozones ( Figure 6). The base of magnetozone N1 is not identified due to an interval of undefined polarity at the base of the core with most samples characterized by radial overprints. The upper part of magnetozone N1 is well defined by four normal polarity samples. Samples 313.21 (normal polarity) and 311.4 (reversed polarity) mbsf define the boundary of magnetozone N1 with R1. Between 312.3 and 302.38 mbsf magnetozone R1 is clearly identified, although it includes several samples with noisy data and others that do not demagnetize. Magnetozone N2 is well defined by three normal polarity samples between 302.82 (top of magnetozone R1) and 297.72 mbsf, which is the base of magnetozone R2. Between 298.4 and 280.09 mbsf, magnetozone R2 includes a long series of samples with reversed polarity with only occasional samples that do not demagnetize. The boundary between magnetozones N3 and R3 lies within an 8-m-interval where polarity could not be determined due to samples in which the radial overprints do not demagnetize. The interval between ca. 270 and 241.08 mbsf includes two intervals of well-defined reversed polarity: magnetozones R3 and R4, with a long interval magnetozone N4 where polarity is unclear with several samples not demagnetizing and others with persistent overprints. Magnetozone N5 is defined at its base between 241.08 and 240.49 mbsf and at its top between samples 236.37 and 235 mbsf. Magnetozones R5, N5 and N4 contain several samples that do not fully demagnetize. The boundary between magnetozones R5 and N6 is not well defined due to an interval including some samples that do not demagnetize and are noisy and have radial overprints. Magnetozone N6 is represented by four samples but also contains some samples that do not demagnetize. The boundary between magnetozones N6 and R6 is at 219.83 mbsf; at the midpoint between samples of opposite polarity at 220.74 and 218.92 mbsf. Magnetozone N7 includes only two samples and is separated from magnetozone R6 by an interval of questionable polarity conditions containing several noisy samples that do not demagnetize and have radial overprints. The interval of the core above 210 mbsf includes a well-defined N-R-N-R pattern of polarity, with the boundary between magnetozones N7 and R7 at 206.70 mbsf between samples 206.94 and 206.47 mbsf. Magnetozone R7 includes several noisy samples that do not demagnetize and have radial overprints. The boundary between magnetozones R7 and N8 is at 175.87 mbsf midway between a reversed sample at 177.53 mbsf and a normal sample at 174.22 mbsf. The boundary between magnetozone N8 and R8 is identified at 160.65 mbsf between samples at 162.11 and 159.20 mbsf.
Although most of the magnetozone boundaries were identified, some of the boundaries, particularly in the lower interval of the core, could not be located precisely because of intervals with uncertain polarity due to strong radial overprints, samples that do not demagnetize and others which were truncated by sedimentary unconformities identified by abrupt grain size changes. The boundaries between magnetozones N6 and R6, R4 and N5, and within N4 occur within sedimentary unconformities.

Biostratigraphic Data
No further biostratigraphic information has been reported for the 316-123 mbsf interval (Cores 34-14) since the original biostratigraphic first and last occurrence information provided by the Shipboard Report (Table 1; . Here, we update those occurrences accounting for more recent age and occurrence data from other Southern Ocean drill holes in order to correlate our magnetic polarity stratigraphy with the Geomagnetic Polarity Time Scale (GPTS) (Figure 7).
Calcareous nannofossils are present only in isolated horizons and burrows in the interval covered by , but only in Core 21 (199.5-190 mbsf), Introduction, 93 cm and Material and Methods, 70 cm, was the assemblage sufficient for an age determination (Chiasmolithus altus assemblage) (Burns, 1975;Bukry, 1975 , Table 1). Fioroni et al. (2012), in a Southern Ocean calibration, attributed the Chiasmolithus altus biozone to the time interval from chrons C12r to C8n, which was later confirmed by Kulhanek et al. (2019). Dictyococcites scrippsae (middle Eocene to Upper Oligocene; Bukry and Percival, 1971), Coccolithus pelagicus (Danian to present; Wallich, 1877) and Reticulofenestra species (Ypresian to present, Fioroni et al., 2015) were also found in the same samples.
On the basis of radiolarian assemblages and their comparison with foraminiferal and calcareous nannofossil assemblages, Chen (1975) hypothesized a hiatus between cores 20 (190.0-180.5 mbsf) and 19 (180.5-171.0 mbsf). This interpretation seems to be supported by his data, with the Miocene and younger portion of the sequence being well constrained based on radiolarian marker species which are well calibrated, and an older section starting with Core 20 (sample 20R-3, 40-50 cm) which is substantially different, older and much more uniform in its radiolarian assemblage ( Table 1). The LADs recorded in Core 20 include Larcopyle FIGURE 7 | Age vs. depth plot with correlation of DSDP Site 274 polarity zonation to the GPTS of Gradstein et al. (2012). Hiatuses defined from the sedimentological data and biostratigraphical datum are represented following the legend. Magnetostratigraphic model A is in dark red; models B and E are in red; models C and D are in orange.  , and Calocyclas (?) semipolita (30.48-?33.06 Ma) which, while clearly identifying this core as substantially older than Core 19, make it difficult to assign it a well-constrained age.

Correlation to the Geomagnetic Polarity Time Scale
Correlation of the paleomagnetic polarity zonation with the GPTS (Gradstein et al., 2012) provides an age interpretation between the upper Eocene and the Middle Miocene (from Chron C15n to C5Br) (Figure 7).
Biostratigraphic data identify a major unconformity at ∼180 mbsf  separating Early/Late Oligocene strata below from Early/Middle Miocene strata above. More recent constraints on age data ( Table 1) restrict the interval of core beneath the unconformity to the late Eocene-Early Oligocene and above the unconformity to the late Early-Middle Miocene (Figure 7).
Beneath the unconformity, magnetozones N1 to R7 are correlated to chrons C15-C9 as follows (Figure 7): Between 315 and 281 mbsf the normal magnetozones N1 and N2 are correlated with chrons C15n and C13n, respectively, and reversed magnetozones R2 and R2 with chrons C13r and C12r, respectively. This correlation implies a sedimentation rate of ∼1 cm/kyr for the late Eocene/Early Oligocene interval of the core with the Eocene/Oligocene boundary at ∼304 mbsf between Core 33 and Core 32. Two correlations are considered between 280 and 180 mbsf (magnetozones N3-R7; Figure 7). The younger correlates magnetozones N3. N5, R5, N6, R6, N7 and R7 with chrons C12n, C11n.2n, C11n.1r, C10n.2n, C10n.1r, C10n.1n and C9r (model D), respectively and implies that the last appearances of radiolarians Cyclampterium milowi and Calocyclas semipolita are almost one million year younger than their currently identified last occurrences. The older correlation links magnetozones N5, R5, N6, R6, N7 and R7 with chrons C12n, C11r, C11n.2n, C11n.1r, C11n.1n and C10r (model E) and has a smaller lag and consequently more consistent with the radiolarian late occurrence data but leaves magnetozone N3 with no corresponding chron in the GPTS. While we expect to see magnetozones truncated and potentially missing from the DSDP Site 274 record, magnetozone N3 is represented by several samples and multiple normal polarity chrons, and we would not expect to see an additional magnetic polarity interval not identified in the GPTS. Regardless, we observe that the sediment accumulation rate in the mid Oligocene interval of the core (280-180 mbsf) is more than four times that of the lower Eocene-Oligocene boundary interval, with 25% of the time interval accounted for by distributed unconformities as identified by marked grain size changes (Figures 2, 7).
Our new age model implies that the unconformity at 180 mbsf accounts for between 9 Myr at its shortest and 11 Myr at its longest, which is significantly shorter than the >15 Myr identified from biostratigraphy only in the original age depth model of .
According to our correlation and calculation of the sedimentary accumulation rate, the studied section encompasses ∼20 Myr. For the interval 180-123 mbsf (cores 19-14), we produced the following estimates of mean sedimentation rates for the two most likely age reconstructions: (model A) 2.0 cm/kyr (C5B-C5C), (model B) 5.1 cm/kyr (C5C-C5D) and (model C) 3.3 cm/kyr (C5D-C5E). The interval between 320 and 180 mbsf (cores 34-20) has a mean sediment accumulation rate of 1.5 cm/kyr (model D) and 1.4 cm/kyr (model E) with a mean of 0.8 cm/kyr between chrons C12n and C15n. The highest sedimentation rate is observed during C5Cn (following model B) with 5.1 cm/kyr and lowest sediment accumulation rate is observed during Chron C12n 0.7 cm/kyr.

DISCUSSION
We present a new magnetostratigraphic chronology for an upper Eocene to Middle Miocene sedimentary section cored during DSDP Leg 28 (Site 274) on the continental rise off Victoria Land, Ross Sea. We confirm a significant hiatus between chrons C9r and C5Cr/C5Dr/C5Er, i.e., from the earliest Chattian to the mid Burdigalian (Figure 7), and its duration is significantly shorter than previous estimates.
Our new age model also identifies a four-fold increase in sedimentation rate on the Antarctic margin after the Eocene/ Oligocene boundary (Figure 7). Previous drilling in McMurdo Sound (southern Victoria Land) also identified a significant early Oligocene unconformity ) and a significant increase in sediment accumulation rates across the Eocene/ Oligocene boundary Wilson et al., 2000;Hannah et al., 2001). The increase in sedimentation rate has been linked to the onset of Antarctic ice sheet growth across the Eocene/Oligocene boundary (Salamy and Zachos, 1999;Galeotti et al., 2016). We confirm those findings across the wider Antarctic shelf and slope. The unconformity at 180 mbsf, is coincident with other significant Late Oligocene unconformities reported from the Ross Embayment (Florindo et al., 2005) but extends into the mid Miocene. We relate the long unconformity to be linked to the early evolution of the Antarctic Circumpolar Current which was established in the Late Oligocene and intensified into the Miocene after the growth of Antarctic ice sheets at the Eocene-Oligocene Boundary (Barker et al., 2007;Lyle et al., 2007;Jovane and Verosub, 2011). Sedimentation did not begin again at DSDP Site 274 until the Antarctic Polar Front Frontiers in Earth Science | www.frontiersin.org December 2020 | Volume 8 | Article 563453 migrated northward coincident with Mid Miocene Glacial intensification (Nelson and Cook, 2001). The unconformity at ca. 180 mbsf could also be partially related to the formation of the Adare Basin, which accommodates the sediments from the West Antarctica Rift System and forms a barrier for the arrival of sediments to the site (Granot et al., 2010(Granot et al., , 2013. Regardless of the exact explanation of the unconformity, these insights are invaluable because they demonstrate that the Early Miocene is characterized by long-term fluctuations of the Antarctic Ice Sheets (Levy et al., 2016;Levy et al., 2019;Jovane et al., 2019;Evangelinos et al., 2020) and that, prior to the Middle Miocene Climatic Optimum (∼15 Ma), there was a period of climatic instability modulated by long-term variations with phases of bottom current intensification which either eroded or did not allow the deposition of sediments in the deep parts of the basin.

CONCLUSION
We provided new paleomagnetic results from the late Eocene-Middle Miocene samples from DSDP Site 274, cored during Leg 28 on the continental rise off Victoria Land, Ross Sea.
-Based on our new age model, the cored late Eocene-Middle Miocene sequence covers an interval of almost 20 Myr, from ∼35 to ∼15 Ma. -We confirm and better define a major unconformity identified at ∼180 mbsf which represents at least 9 Myr from the earliest Chattian to the mid Burdigalian. It may represent non-deposition and/or erosion due to intensification of Antarctic Circumpolar Current activity. -Significant fluctuations in grain size and magnetic properties are observed above the unconformity (above 180 mbsf) interpreted as cyclical behavior in glacial advance and retreat from the continent. This new age model will allow further studies of this unique core to reconstruct paleoenvironmental studies focusing on the discharge of sediments from the catchment areas in this region of the Ross Sea.

DATA AVAILABILITY STATEMENT
The raw data supporting the conclusions of this article will be made available by the authors, without undue reservation.

AUTHOR CONTRIBUTIONS
LJ: paleomagnetic measurements, data analysis, discussion and writing; FF: discussion and writing; DR: paleomagnetic measurements, data analysis, and writing; SL: grain size measurements, data analysis, and writing; MH: paleomagnetic measurements, data analysis, and writing; GC: radiolaria and writing; GW: data analysis, discussion and writing.

ACKNOWLEDGMENTS
We thank the editor Hagay Amit and the reviewers for useful comments and corrections which greatly improved the manuscript. The Deep Sea Drilling Project samples and data were provided by the International Ocean Discovery Program (sample requests: 043328-IODP and 069232-IODP). The samples were collected at the UC Davis Paleomagnetism Laboratory as part of NSF grant OPP 06-17194 to K. L. Verosub. LJ, MBH and DR acknowledge funding from Fundação de Amparo á Pesquisa do Estado de São Paulo (FAPESP) grant 2016/24946-9 and 2018/17061-6. LJ is also financed by the Coordenação de Aperfeiçoamento de Pessoal de Nível Superior -Brasil (CAPES)-Finance Code 001. The data are available from the authors upon request.