Aims & Scope

Interaction of primary/secondary cosmic rays with the atomic nuclei in the upper atmosphere/earth’s surface produces ¹⁰Be and other cosmogenic nuclides such as ³He, ¹⁴C, ²¹Ne, ²²Ne, ²⁶Al, ³⁶Cl, and ⁵³Mn (Lal and Peters, 1967). The production of ¹⁰Be usually occurs due to the spallation of Oxygen and Carbon by the cosmic ray particles. Nucleons, protons in the upper atmosphere, and neutrons at the earth's surface dominate cosmogenic nuclide production, whereas subsurface production is dominated by muons, which weakly interact with matter and retain energy for deeper penetration (Lal, 1988). Distinct processes produce two different varieties of ¹⁰Be: meteoric ¹⁰Be is produced in the upper atmosphere, whereas in-situ ¹⁰Be is produced on Earth’s surface when cosmic rays interact with oxygen atoms within the lattices of mineral grains (Quartz) (Lal and Peters, 1967). The meteoric ¹⁰Be concentration in sediments has wider applicability for numerous geophysical and geochemical studies due to its high sensitivity to earth surface processes and its application to quartz free lithologies (Willenbring and von Blanckenburg, 2010). Meteoric ¹⁰Be is a potential tool for estimating basin-wide erosion rate (von Blanckenburg et al., 2012), reconstructing paleo-secular variations of the earth’s magnetic field (Aldahan and Possnert, 2003; Carcaillet et al., 2004; Simon et al., 2020), and estimating the sea-ice extent (Frank et al., 1995) and glacial meltwater discharge (Simon et al., 2016a). When normalized to its stable counterpart ⁹Be, the ratio can circumvent the variations relating to particle size or retention behavior (von Blanckenburg et al., 2012) and can be applied to measure spatio-temporal variability of continental erosion rates (von Blanckenburg and Willenbring, 2014; von Blanckenburg et al., 2015) and record deep ocean circulation patterns (von Blanckenburg et al., 1996). Due to the long half-life of ¹⁰Be (1.39×10⁶ yr), this ratio can be applied to study ice sheet melting dynamics for more extended periods than could be achieved using other traditional particle reactive radioisotopes (t_1/2 ²³⁰Th = 7.5×10⁴ yr and t_1/2 ²³¹Pa = 3.2×10⁴ yr).

At present, there exist only a few works concerning the ¹⁰Be in Antarctic sediments. Frank et al. (1995) studied sediment redistribution and paleo-productivity changes in the Weddell Sea sediment record and documented an increase in ¹⁰Be concentration during the late Quaternary interglacials. In another study, Frank et al. (2002) studied quasi-conservative behavior of dissolved ¹⁰Be from depth profiles of Antarctic seawater and suggested that the ¹⁰Be concentrations in deeper water are less affected by the scavenging process or ice melt. Other similar studies around the Antarctic used Be isotopes to distinguish glacial/post-glacial sediment sources (Sjunneskog et al., 2007; Yokoyama et al., 2016), estimate variations in sea ice cover, and study amplitude and rhythm of geomagnetic field change (Raisbeck et al., 2006). Despite previous high-resolution studies, the present understanding of the variability of Be isotope concentrations in Antarctic sediments concerning paleoenvironmental changes has remained unclear, as the complex characteristics of glacial and glacio-marine sediments do not offer an uncomplicated interpretation of climatic or chronological signals (Sjunneskog et al., 2007). Although Be isotopes are subject to reworking but have different pathways than other cosmogenic radioisotopes (e.g., ¹⁴C) (Willenbring and von Blanckenburg, 2010), hence can compensate for uncertainties associated with chronology and can illustrate changes in earth surface processes.

In the present study, we analyze ¹⁰Be (meteoric), ⁹Be, and their ratio from a Ross Sea sediment core to enhance our previous understanding of paleoclimatic records from Antarctic continental shelf sediments. The prime focus of this study is to analyze the applicability of Be isotopes as a proxy to reconstruct Antarctic Ice sheet melting by pulsed warming intervals during mid to late Quaternary and to study the variation of ¹⁰Be flux during the geomagnetic reversals.

Ross Ice Shelf occupies a large part of the Ross Sea embayment, fed by eastern glaciers and western ice streams (Sjunneskog et al., 2007). Together with sea ice, the ice shelf controls the ocean water circulation in decadal to millennial time scales, e.g., the ice shelf exceeded the continental shelf region during the Last Glacial Maxima (Domack et al., 1999; Ship et al., 1999), and its retreat during the post-glacial warm period resulted in renewed exposure to open ocean conditions (Sjunneskog et al., 2007; Jeong et al., 2018). As a consequence of the advance/retreat of the Antarctic ice shelf during glacial/interglacial periods, depositional environments in the Ross Sea have altered. Unlike the meteoric ¹⁰Be, the ⁹Be is primarily terrestrially sourced. Hence the Be isotope ratio variability in glaciomarine sequences will depend on (1) atmospheric flux, (2) sediment pathways and settling time, and (3) syn-depositional reworking. These Antarctic shelf sediment pathways are subject to change during different climactic episodes, thus proving them as a tracer to study climate-induced glaciomarine sediment dynamics.

The Ross Sea is an important aspect of global thermohaline circulation as it is one of the three primary sources of Antarctic Bottom Water (AABW). Flowing into the abyssal ocean, it provides cooling and ventilation of the deep ocean (Orsi et al., 2001). Jacobs et al. (2011) suggests that the AABW in the Ross Sea has recently freshened due to melt water influx from Amundsen and Bellingshausen Seas. Recent studies have suggested that upwelling of Circumpolar Deep Water (CDW) have accelerated the ice melting in the Ross Sea (Pritchard et al., 2012; Thompson et al., 2018), as it is relatively warmer (>0℃).

The Ross Sea contains the most abundant ice volume, which drains the East Antarctic Ice Sheet (EAIS) and the West Antarctic Ice Sheet (WAIS) to the Ross Ice Shelf (RIS; Fig. 1a). The RIS advanced northward during the Last Glacial Maximum (LGM), covering most of the continental shelf of the Ross Sea with glacial sediments (Anderson et al., 2014). The Joides Trough, which is on the outermost area of the continental shelf, was the major ice drainage of EAIS, which transported glacial sediments during the LGM (Fig. 1b) (Anderson and Bartek, 1992; Ship et al., 1999; Harris et al., 2001; Livingstone et al., 2012; Anderson et al., 2014; Halberstadt et al., 2016; Park et al., 2018). The glacial geomorphological features collected from the seafloors indicate that the Joides Trough was the front of the grounding line, and nearby Mawson and Pennell Bank show pre-LGM ¹⁴C datings of carbonates (Taviani et al., 1993; Ship et al., 1999; Anderson et al., 2014).

Fig. 1. 
Study area. (a) Map showing the location of Ross Sea, Weddell Sea, Ross Ice Shelf (RIS), East and West Antarctic Ice Sheet (b) Detailed physiographic map of Ross Sea (c) Bathymetry (in meters) map of Ross Sea showing the location of LC42 core (d) Lithostratigraphy of LC42 (Adapted from Ohneiser et al., 2019). Ohneiser et al. (2019) classified the LC42 core into three lithologically distinct units: Unit I (0-1.55 m) is dominated by gravelly muddy sand with interbeds of fine to medium sand and silt; Unit II (1.55-3.15 m) comprising of gravelly muddy sand, with interbeds of laminated and bioturbated silt; Unit III (3.15-11.78 m) consisting of clast free interbeds of laminated, sandy silt, bioturbated silt and gravelly muddy sands.

The Central Basin is just in front of the Joides Trough, where the glaciomarine process could not affect during the glacial periods, and we expect the pure sea ice variance performance between the permanent sea ice coverage or seasonally opened. The basin is surrounded by the Hallett Ridge (W), abrupt gradient change from the Joides Trough (S), Iselin Bank (E), and the mouth to the Southern Ocean (N) (Fig. 1c).

We acquired sediment core from the Central Basin during the 2015 R/V ARAON cruise of the Korea Polar Research Institute (KOPRI). We collected the RS15-LC42 in 2084 m depth (length: 11.75 m; 71° 49.40′ S, 178° 34.76′ E), and the magnetic polarity of this core exhibited normal-to-reversed magnetism at 8.26 ± 0.03 m (0.781 Ma) (Ohneiser et al., 2019). Assuming constant sedimentation rate and the top sample as of modern age, the age of the core was linearly determined using the geomagnetic reversal marker (Ohneiser et al., 2019).

The lithostratigraphy of the core was described in detail by Ohneiser et al. (2019). We followed the physical and chemical treatment procedures outlined by the Korea University Geochronology Laboratory, Seoul, to extract and measure meteoric ¹⁰Be from the clay rich marine sediments (Bourles et al., 1989; Jeong et al., 2018).

A total of 55 samples were taken for beryllium isotope analyses, roughly at every 30 cm. The samples were dried and grinded for homogenization. Over 200 μm fraction was separated, and only the <200 μm fraction was selected to remove size dependency related to chemical adsorption of beryllium isotope. After oxidization for removing organic materials, 0.4 mg of ⁹Be carrier (~1000 ppm) was added to 0.5 g of sediment samples before beryllium extraction procedures. The ¹⁰Be was extracted through a step-wise leaching procedure using a 0.04 M hydroxylamine in 25% acetic acid solution (Bourles et al., 1989), and its ⁹Be concentration was measured by Inductively Coupled Plasma Atomic Emission Spectrometry (ICP-AES) and Mass Spectrometry (ICP-MS) at Korea Basic Science Institute (KBSI). The relative error (of 1 standard deviation) for ⁹Be measurement is <±2.5%. ¹⁰Be analyses were performed using 6 MV AMS at the Korea Institute of Science and Technology (KIST) (Lee et al., 2011; Rhee et al., 2017). The blank for the beryllium carrier averaged is less than 6.5×10^-15.

The meteoric ¹⁰Be and ⁹Be data for the core taken from the Ross Sea (LC42) is shown in Table 1. The ¹⁰Be concentration (Decay corrected) varies from 5.29±0.05×10⁸ to 0.214×10⁸ atoms/g, with an average value of 1.52±0.02×10⁸ atoms/g. The ¹⁰Be concentration depicts clear contrasts between different time intervals during the past 1 Ma. Between 1 to 0.8 Ma there is a steady decrease in ¹⁰Be concentration (average = 1.14±0.02×10⁸ atoms/g), which increases sharply from 0.8 to 0.7 Ma (average= 1.96±0.03×10⁸ atoms/g). Thereafter, the ¹⁰Be concentration remains approximately constant up to ~0.43 Ma, with an average value of 0.8±0.01×10⁸ atoms/g, which is nearly half of the whole core's average value. After 0.43 Ma, the ¹⁰Be concentration increases steadily, showing maximum value at the top of the core. The average value of ¹⁰Be concentration during the post 0.43 Ma period is 1.85±0.02×10⁸ atoms/g, nearly doubling the average value of the preceding interval. The ⁹Be concentration does not show much fluctuation throughout the core. The relative variation of average ⁹Be concentration during these time intervals is within 10% of the average value of the core (2.01±0.03×10¹⁶ atoms/g), except an anomalously higher value corresponding to 451 cm depth (Age: 425 Ma; ⁹Be concentration: 11×10¹⁶ atoms/g). The downcore variation of the ¹⁰Be/⁹Be ratio (Be ratio, decay corrected) mimics the variation of ¹⁰Be concentration, with an average value of 0.9±0.01×10^-8 for the whole core. The interval between 1 to 0.8 Ma shows a gradual decrease of the ratio (average = 0.7±0.01×10^-8), followed by a sharp increase (0.8-.7 Ma, average = 1±0.01×10^-8). The most noticeable feature of the ¹⁰Be in this interval is a peak overpassing the average value of all samples. The Be ratio remains approximately steady till 0.43 Ma (average = 0.5×10^-8) and further increases towards the top of the core. The average value of Be ratio after 0.43 Ma (1.25±0.01×10^-8) doubles the preceding interval's average value. The time since the mid-Pleistocene can be divided into several climatic episodes based on Be isotope data from the present study and paleoclimatic inferences drawn from previous works. The glacial (light blue)/interglacial (yellow) periods corresponding to the increase/decrease of ¹⁰Be concentration are shown in Fig. 2.

Fig. 2. 
(a) Global average surface temperature (GAST) changes based on deep sea oxygen isotopes over the last 2 Ma (Snyder, 2016). Our EAIS deglaciations coincide with the early Pleistocene warm period (pre-mid-Pleistocene transition, pre-MPT; 45) and the late Pleistocene warm interglacial periods (post-mid-Bruhnes transition event, post-MBE; 51, 56). (b) Climate proxy changes during the last 1 Ma. (1) GAST changes. (2) The temperature gradient between the interior EAIS temperature history (East Antarctica Dome C, EDC3; 47) and present temperature. (3) Surface accumulation rate, in centimeters of ice equivalent per year (EDC3; 57). (4) and (5) meteoric ¹⁰Be concentrations and ¹⁰Be/⁹Be ratios along Long Core 42, collected during the 2015 Ross Sea cruise (this study). The sky-blue boxes represent periods of meteoric ¹⁰Be depletion, which indicates more and thick sea ice coverage over the ocean surface due to a decrease in temperature that hindered the direct deposition of meteoric ¹⁰Be to the seafloor. The yellow boxes indicate seasonally opened periods. The brown box represents a weakened Geomagnetic Dipole Moment (GDM, XXI) period (Matuyama- Brunhes Boundary; MBB) that is < 20% of the averaged GDM and produced higher meteoric ¹⁰Be than the other periods (Simon et al., 2016a). The ¹⁰Be peaks during the period highlighted by the brown box should be interpreted with caution.

Although using a single geomagnetic reversal event to establish the chronology of the core may impose age uncertainties on short-term climatic events, the longer climatic divisions will remain less affected. The Be isotope data from the present study shows changes over sub-millennial time scales, representing the variability of geomagnetic intensity and climate change over broader time intervals (Fig. 2). Thus, the present study encompassing many sub-millennial episodes of climatic variability during the past 1 Ma will not be affected by the age uncertainties.

5.1 Relationship with paleomagnetic changes

The Late Matuyama Chron witnessed an increase in the geomagnetic field intensity with significant peaks at around 870 ka, 910 ka, and 960 ka (Valet and Meynadier, 1993; Meynadier et al., 1994; Kok and Tauxe, 1999; Dinarès-Turell et al., 2002). These episodes of intensity maximas (spanning about 50 ka) were separated by periods of lower intensity. The ¹⁰Be and ¹⁰Be/⁹Be ratios show a gradual decrease during the late Matuyama Chron. Although the coarse sample resolution in the present study does not depict high-resolution geomagnetic changes during this period, a decreasing trend towards the Matuyama-Brunhes boundary (MBB) suggests a relative increase in paleointensity. A significant increase in the ¹⁰Be and ¹⁰Be/⁹Be values is observed during 0.8-0.7 Ma. No significant correlation is found between the Be isotopic record and other available paleoclimatic proxies during this period, suggesting a possible paleointensity variation (Ohneiser et al., 2019). The Be overproduction episode observed during this period is synchronous with U-Pb zircon age of MBB (772.7±7.2 ka) from Byakubi-E tephra (coupled with Oxygen isotope chronology) (Suganuma et al., 2015), ¹⁰Be overproduction observed in equatorial Pacific ocean (Simon et al., 2016b), the age of MBB obtained from ⁴⁰Ar/³⁹Ar chronology of North Atlantic marine cores (mean age of 773.1 ka) (Channell et al., 2010) and the peak ¹⁰Be flux (during 777-766 ka) observed in EPICA (European Project for Ice Coring in Antarctica) Dome C ice core (Raisbeck et al., 2006). This near-doubling of ¹⁰Be production during the MIS 19 resembles overproduction episodes associated with the largest geomagnetic reversals or collapse of geomagnetic dipole during long-term reversals (Raisbeck et al., 2006). Since there is no previous ¹⁰Be overproduction episode closely associated with the MBB overproduction, our results contradict the frequently reported pre- MBB precursor event (Hartl and Tauxe, 1996; Singer et al., 2005; Raisbeck et al., 2006).

Yamazaki (1999) documented high-resolution relative paleointensity variations during the Brunhes Chron from sediment cores of the Pacific Ocean. The results showed an increase in intensity after the MBB till the Mid-Bruhnes Event (MBE) (during the pre-MBE period). In contrast, a gradual decrease in intensity is observed during the post MBE period. A significant drop in intensity was observed at about 30-50, 110-120, 180-200, 280-300, 380-410, and 520-550 ka. The Be isotopic records from our study agree with the relative paleointensity records from the previous studies. The ¹⁰Be and ¹⁰Be/⁹Be ratios are approximately constant during the pre-MBE period and show a gradual increase after the MBE. Although the sampling resolution for the present study was high, geomagnetic reversals/excursions during the Brunhes Chron could be identified based on relative variation in the ¹⁰Be and ¹⁰Be/⁹Be records (Fig. 3). Be overproduction episodes are observed during 25 ka (Mono Lake excursion, dated 24-39 ka by Kent et al. (2002)), 114 ka (Blake excursion, dated 110-120 ka by Macrì et al. (2005)), 171 and 284 ka (Biwa I and Biwa II, respectively, dated 187-192 ka and 292-298 ka by Worm, (1997)), 322 ka (Calabrian Ridge 1, dated at 315-325 ka by Langereis et al. (1997)), and 378 ka, (Levantine dated at 360-370 ka by Langereis et al. (1997)).

Fig. 3. 
Relative variation of ¹⁰Be concentration and ¹⁰Be/⁹Be ratio during different geomagnetic reversals/excursions.

The ¹⁰Be concentration in the Antarctic open marine condition is about 1.5×10⁹ atoms/g (Sjunneskog et al., 2007). However, the LC42 record exhibit concentration of a magnitude of less than one degree throughout the core. The concluding points from the present study are:

1) The downcore variation of the ¹⁰Be and ¹⁰Be/⁹Be in the Ross Sea sediment core illustrates glacial-interglacial transitions during the mid to late Pleistocene period. The relatively high/low concentration of ¹⁰Be indicates relatively open marine/sea ice-dominated conditions, respectively.

2) The gradual decrease in the ¹⁰Be concentration from 1 to 0.8 Ma suggests relatively extended glacial conditions persisted during the late Matuyama Chron. This study shows that the MIS 22 was the coldest period during the late Matuyama Chron, which corroborates previous studies showing an abrupt increase in the ice sheet volume at the onset of MIS 22. The progressively increasing trend observed in the ¹⁰Be concentration after 430 Ka suggests that the post-MBE period was relatively warmer than the pre-MBE period.

3) The downcore variation of the ¹⁰Be and ¹⁰Be/⁹Be ratio shows some degree of correlation with the geomagnetic reversals/excursions. No precursor event associated with the MBB reversal contradicts previous investigations showing a pre-MBB overproduction episode closely associated with the MBB overproduction episode.

4) This study shows that further investigations in marginal environments of RIS may be useful to document high-resolution glacial-interglacial transitions, as sediment characteristics at this boundary between the ice shelf and open marine environment is sensitive to small changes in the amplitudes of climate cyclicity.