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Sea Ice Changes in the Southwest Pacific Sector of the Southern Ocean During 1 the Last 140 , 000 Years 2 3

18 Sea ice expansion in the Southern Ocean is believed to have contributed to glacial-interglacial 19 atmospheric CO2 variability by inhibiting air-sea gas exchange and influencing the ocean’s 20 meridional overturning circulation. However, limited data on past sea ice coverage over the last 21 140 ka (a complete glacial cycle) have hindered our ability to link sea ice expansion to oceanic 22 processes that affect atmospheric CO2 concentration. Assessments of past sea ice coverage 23 using diatom assemblages have primarily focused on the Last Glacial Maximum (~21 ka) to 24 Holocene, with few quantitative reconstructions extending to the onset of glacial Termination II 25 (~135 ka). Here we provide new estimates of winter sea ice concentrations (wSIC) and summer 26 sea surface temperatures (sSSTs) for a full glacial-interglacial cycle from the southwestern 27 Pacific sector of the Southern Ocean using fossil diatom assemblages from deep-sea core 28 TAN1302-96 (59.09°S, 157.05°E, water depth 3099 m). We find that winter sea ice was 29 consolidated over the core site during the latter part of the penultimate glaciation, Marine 30 Isotope Stage (MIS) 6 (from at least 140 to 134 ka), when sSSTs were between ~1 and 1.5°C. 31 The winter sea ice edge then retreated rapidly as sSSTs increased during the transition into the 32 Last Interglacial Period (MIS 5e), reaching ~4.5°C by 125 ka. As the Earth entered the early 33 glacial stages, sSSTs began to decline around 112 ka, but winter sea ice largely remained absent 34 until ~65 ka during MIS 4, when it was sporadically present but unconsolidated (<40% wSIC). 35 WSIC and sSSTs reached their maximum concentration and coolest values by 24.5 ka, just prior 36 to the Last Glacial Maximum. Winter sea ice remained absent throughout the Holocene, while 37 SSSTs briefly exceeded modern values, reaching ~5°C by 11.4 ka, before decreasing to ~4°C and 38 stabilizing. The absence of sea ice coverage over the core site during the early glacial period 39 suggests that sea ice may not have been a major contributor to CO2 drawdown at this time. 40 During MIS 5d, we observe a weakening of meridional SST gradients between 42 ̊ to 59 ̊S 41 throughout the region, which may have contributed to early reductions in atmospheric CO2 42 concentrations through its impact on air-sea gas exchange. Sea ice expansion during MIS 4, 43 however, coincides with observed reductions in Antarctic Intermediate Water production and 44 https://doi.org/10.5194/cp-2021-107 Preprint. Discussion started: 17 August 2021 c © Author(s) 2021. CC BY 4.0 License.

processes that affect atmospheric CO2 concentration. Assessments of past sea ice coverage 23 using diatom assemblages have primarily focused on the Last Glacial Maximum (~21 ka) to 24 Holocene, with few quantitative reconstructions extending to the onset of glacial Termination II 25 (~135 ka). Here we provide new estimates of winter sea ice concentrations (wSIC) and summer 26 sea surface temperatures (sSSTs) for a full glacial-interglacial cycle from the southwestern 27 Pacific sector of the Southern Ocean using fossil diatom assemblages from deep-sea core 28 TAN1302-96 (59.09°S, 157.05°E, water depth 3099 m). We find that winter sea ice was 29 consolidated over the core site during the latter part of the penultimate glaciation, Marine 30 Isotope Stage (MIS) 6 (from at least 140 to 134 ka), when sSSTs were between ~1 and 1.5°C. 31 The winter sea ice edge then retreated rapidly as sSSTs increased during the transition into the 32 Last Interglacial Period (MIS 5e), reaching ~4.5°C by 125 ka. As the Earth entered the early 33 glacial stages, sSSTs began to decline around 112 ka, but winter sea ice largely remained absent 34 until ~65 ka during MIS 4, when it was sporadically present but unconsolidated (<40% wSIC).

35
WSIC and sSSTs reached their maximum concentration and coolest values by 24.5 ka, just prior 36 to the Last Glacial Maximum. Winter sea ice remained absent throughout the Holocene, while 37 SSSTs briefly exceeded modern values, reaching ~5°C by 11.4 ka, before decreasing to ~4°C and 38 stabilizing. The absence of sea ice coverage over the core site during the early glacial period 39 suggests that sea ice may not have been a major contributor to CO2 drawdown at this time.

40
During MIS 5d, we observe a weakening of meridional SST gradients between 42˚ to 59˚S 41 throughout the region, which may have contributed to early reductions in atmospheric CO2 42 concentrations through its impact on air-sea gas exchange. Sea ice expansion during MIS 4, 43 however, coincides with observed reductions in Antarctic Intermediate Water production and 44 2.0 Methods 104

Study Site & Age Determination 105
We reconstruct diatom-based wSIC and sSST using marine sediment core TAN1302-96 106 benthic foraminifera (yellow circles). Metadata for these cores are provided in Table A1. ice extents, respectively (Reynolds et al., 2002;2007). 119 120 The age model for TAN1302-96 (Figures 2 and 3) was based on a combination of 121 radiocarbon dating of mixed foraminiferal assemblages, and stable oxygen isotope stratigraphy 122 on Neogloboquadrina pachyderma (180-250 µm). Seven accelerator mass spectrometry (AMS) 123 14 C samples were collected (Table A1 in   TAN1302-96 was sampled every 3-4 cm throughout the core except between 130-180 158 cm, where samples were collected every 10 cm due to limited availability of sample materials 159 (Table A3 in Appendix A). Diatom slide preparation followed two procedures. The first approach 160 approximated the methods outlined in Renberg (1990), while the second followed the protocol 161 outlined in Warnock & Scherer (2014). To ensure there were no biases between preparation 162 techniques, results from each technique were first visually compared followed by a comparison 163 of sample means (see Figure B1 in Appendix B). No biases in the data were observed between 164 methods. 165 The first procedure was conducted at Victoria University of Wellington and Simon Fraser 166 University on samples every 10 cm throughout the core. Sediment samples contained high 167 concentrations of diatoms with little carbonaceous or terrigenous materials, so no dissolving 168 aids were used. Instead, approximately 50 mg of sediment was weighed, placed into a 50 ml 169 centrifuge tube, and topped up with 40 ml of deionized water. Samples were then manually 170 shaken to disaggregate sediment, followed by a 10-second mechanical stir using a vortex 171 machine. Samples were then left to settle for 25 seconds. 0.25 mL of the solution was then 172 pipetted onto a microscope slide from a consistent depth, where it was left to dry overnight. 173 Once the sample had dried, coverslips were permanently mounted to the slide using Permount, 174 a high refractive index mountant. Slides were redone if they contained too many diatoms and 175 identification was not possible, or if they contained too few diatoms (generally <40 specimens 176 per transect). Sediment sample weight was adjusted to achieve the desired dilution. 177 The second procedure was conducted at Colgate University on samples every 3-4 cm 178 throughout the core. Oven-dried samples were placed into a 20 ml vial with 1-2 ml of 10% H2O2 179 and left to react for up to several days, followed by a brief (2-3 second) ultrasonic bath to 180 disaggregate samples. The diatom solution was then added into a settling chamber, where 181 microscope coverslips were placed on stages to collect settling diatoms. The chamber was 182 gradually emptied through an attached spigot, and samples were evaporated overnight. Cover 183 slips were permanently mounted onto the slides with Norland Optical Adhesive #61, a 184 mounting medium with a high refractive index. 185 Diatom identification was conducted at Simon Fraser University using a Leica Leitz 186 DMBRE light microscope using standard microscopy techniques. Following transverses, a 187 minimum of 300 individual diatoms were identified at 1000x magnification from each sample 188 throughout the core. Individuals were counted towards the total only if they represented at 189 least one-half of the specimen so that fragmented diatoms were not counted twice.

Modern Analog Technique 211
Past wSIC and sSSTs (January to March) were estimated for TAN1302-96 and 212 recalculated for SO136-111 by applying the Modern Analog Technique (MAT) to the fossil 213 diatom assemblages, as outlined in Crosta et al. (1998;. The MAT reference database 214 used for this analysis is comprised of 249 modern core top samples (analogs) located primarily 215 in the Atlantic and Indian sectors from ~40°S to the Antarctic coast. The age of the core tops 216 included in the reference database have been assessed through radiocarbon and/or isotope 217 stratigraphy when possible. Core tops were visually evaluated for selective diatom dissolution, 218 so it is believed that sub-modern assemblages contain well-preserved and unbiased specimens. 219 Modern summer SSTs and wSIC were interpolated from the reference core locations using a 220 to approximately 0% over the same time periods (Figure 4). Reconstructed sSSTs continued to 255 rise slightly throughout MIS 5e, reaching a maximum value of 4.6°C at 118 ka, after which they 256 declined throughout MIS 5. During this period of sSST decline, winter sea ice was largely absent, 257 punctuated by brief periods during which sea ice was present but unconsolidated (wSIC of 258 14.7% and 17% at 105 and 90 ka, respectively). During MIS 4 (71 to 57 ka), sSSTs cooled to 259 between roughly 1°C and 3°C, and sea ice expanded to 36%, such that it was present but 260 unconsolidated for intervals of a few thousand years. SSSTs increased slightly from 1.5°C at 61 261 ka (during MIS 4) to ~2.5°C at 50 ka (during MIS 3), followed by a general cooling trend into MIS 262 2. Sea ice appears to have been largely absent during MIS 3 (57 to 29 ka), although sampling 263 resolution is low, but increased rapidly to 48% cover during MIS 2 where winter sea ice was 264 consolidated over the core site. During MIS 2, sSSTs cooled to a minimum of <1°C at 24.5 ka. 265 After 18 ka, the site rapidly transitioned from cool, ice-covered conditions to warmer, ice-free 266 winter conditions during the early deglaciation. This warming was interrupted by a brief cooling 267 estimates than the MAT (e.g., ~50% wSIC at 23 ka while the MAT produced ~37% for the same 314 time period); however, we believe it is unlikely that statistical approaches alone could explain a 315 larger difference (i.e., 50%) between E27-23 and TAN1302-96. 316 The second possible explanation is through lateral sediment redistribution and focusing 317 by the ACC. We estimated sediment focusing for E27-23 using 230 Th data from Bradtmiller et al. 318 (2009) together with dry bulk density estimated using calcium carbonate content (Froelich, 319 1991). Both sedimentation rates and focusing factors for the E27-23 are relatively high (~35 320 cm/ka and max=26, respectively) during the LGM and Holocene, which could influence the 321 reliability of wSIC and sSST estimation (see Figure B2 in Appendix B). Several peaks in focusing 322 occurring around 16, 12, and 3 ka appear to closely correspond to periods of peak wSIC (~67%, 323 ~54%, and ~35%, respectively), suggesting a possible link. Lateral redistribution could artificially 324 increase or decrease relative abundances of some diatom groups, which could lead to over or 325 under estimations of sea ice coverage. Thorium analysis for TAN1302-96 is beyond the scope of 326 this study; however, future work could help address this uncertainty. 327 Although we are unable to identify the specific cause of the differences, we do suggest 328 using caution when interpreting the exact magnitude of sea ice expansion in this region, and to 329 consider the results from all cores when drawing conclusions of regional sea ice history.

Other Potential Contributors to Early Glacial CO2 Variability 375
The changes observed in wSIC and sSST from TAN1302-96 suggests that sea ice 376 expansion was likely not extensive enough early in the glacial cycle for a sea ice capping effect The first possible explanation considers that not all sea ice has the same capacity to 384 facilitate or inhibit air-sea gas exchange. We previously suggested that because sea ice was not 385 at its maximum extent during MIS 5d, the contribution of sea ice on CO2 sequestration would 386 likely not be at its maximum extent either. However, this assumes that there is a linear 387 relationship between sea ice coverage and CO2 sequestration potential. We know that different 388 sea ice properties, such as thickness and temperature, determine overall porosity, with thicker 389 and colder sea ice being less porous and more effective at reducing air-sea gas exchange 390 compared to thinner and warmer sea ice (Delille et al., 2014). It is therefore possible that if 391 modest sea ice advances took place closer to the Antarctic continent (and were therefore not 392 captured by TAN1302-96), they may have been more effective at reducing CO2 outgassing, 393 either by experiencing some type of reorganization or consolidation, or through a change in 394 properties such as temperature or thickness. It is also possible that sea ice coverage over some 395 regions leads to more effective capping, while in other regions sea ice growth contributes only 396 to marginal reductions in air-sea gas exchange. This, theoretically, could point to a non-linear 397 response between sea ice expansion and CO2 sequestration potential, and could link modest 398 sea ice growth around the Antarctic continent to the ~35 ppm initial CO2 drawdown event. 399 While this is theoretical and cannot be adequately addressed in this analysis, it is worthy of 400 deeper consideration. 401 The second possible explanation involves changes in the global overturning circulation.

Sea Ice Expansion and Ocean Circulation 447
Although the TAN1302-96 wSIC record suggests that sea ice was largely absent at the 448 core site until the mid-glacial (~65 ka), the observed changes in sea ice throughout the glacial-449 corresponding to lower %AAIW ( Figure 5). This means that during cooler periods, the AAIW-471 UCDW interface shoaled, reducing the total volume of AAIW and indirectly causing an 472 expansion of UCDW (Ronge et al., 2015). 473 Our comparison between %AAIW and regional wSIC estimates suggest a strong link 474 between the two ( Figure 5). Specifically, we observe that sea ice expansion occurs during time 475 periods when AAIW has shoaled and UCDW has expanded (i.e., %AAIW is low). In contrast, 476 during periods of low wSIC and warmer summer sea surface temperatures (e.g., MIS 5e), 477 %AAIW is high. This correlation supports the idea that increased concentrations of regional sea 478 ice resulted in a substantial summer freshwater flux into the AAIW source region. This regional 479 freshening likely promoted a shallower subduction of AAIW and a corresponding volumetric 480 expansion of UCDW, which can be seen by the isotopic offset of the δ 13 C values between the 481 reference cores, and also by the increased carbonate dissolution in MD97-2120 during glacial ice coverage is minimal and freshwater input from summer sea ice melt is low. This lack of 487 freshwater input allows AAIW to subduct to deeper depths and bath core MD97-2120, 488 capturing the higher δ 13 C signature of the overlying AAIW waters. The AAIW-UCDW interface 489 (red dashed line) is located beneath MD97-2120. CO2 outgassing is occurring as carbon-rich 490 Circumpolar Deep Waters upwell near Antarctica. B) Depicts glacial conditions where sea ice 491 expansion has occurred beyond TAN1302-96, increasing brine rejection, and stabilizing the 492 water column. As a result of the increased sea ice growth, subsequent summer melt increases 493 the freshwater flux into the AAIW source region and increases AAIW buoyancy. This buoyancy 494 gain shoals the AAIW-UCDW interface above core MD97-2120, causing the core site to be 495 bathed in low δ 13 C UCDW. The shoaling of AAIW causes an indirect expansion of CDW, 496 increasing the glacial carbon stocks of the deep ocean while sea ice reduces CO2 outgassing via 497 the capping mechanism. 498 499 In addition to its influence on regional freshwater forcing and AAIW reductions, these 500 and observed proxy data, and fit well within the hypothesis that mid-glacial CO2 variability was 520 primarily the result of a more sluggish overturning circulation (Kohfeld & Chase, 2017). 521 522

523
This study presents new wSIC and sSST estimates from marine core TAN1302-96, 524 located in the southwestern Pacific sector of the Southern Ocean. We find that the wSIC 525 remained low during the early glacial cycle (130 to 70 ka), expanded during the middle glacial 526 cycle (~65 ka), and reached its maximum just prior to the LGM (~24.5 ka). These results largely 527 agree with nearby core SO136-111 but display some differences in wSIC magnitude with E27-528 23. This discrepancy may be explained by differences in statistical applications and/or lateral 529 sediment redistribution, although more analysis is required to determine the exact cause(s). 530 The lack of changes in sSSTs and the absence of sea ice over the core site during the 531 early glacial suggests that the sea ice capping mechanism and corresponding surface 532 stratification in this region is an unlikely cause for early CO2 drawdown, and that alternative 533 hypotheses should be considered when evaluating the mechanism(s) responsible for the initial 534 drawdown. More specifically, we consider the impact of changes in SST gradients between ~40˚ 535 to 60˚S and support the idea that changes in air-sea disequilibrium associated with reduced 536 sea-surface temperature gradients could be a potential mechanism that contributed to early 537 glacial reductions in atmospheric CO2 concentrations (Khatiwala et al., 2019). Another key 538 consideration is the potentially non-linear response between sea ice expansion and CO2 539 sequestration potential (i.e., that not all sea ice is equal in its capacity to sequester carbon). 540 More analyses are required to adequately address this. 541 We also observe a strong link between regional sea ice concentrations and vertical 542 fluctuations in the AAIW-UCDW interface. Regional sea ice expansion appears to coincide with 543 the shoaling of AAIW, likely due to the freshwater flux from summer sea ice melt increasing 544 buoyancy in the AAIW formation region. Furthermore, major sea ice expansion and AAIW 545 shoaling occurs during the middle of the glacial cycle and is coincident with previously 546 recognized shoaling in AMOC and mid-glacial atmospheric CO2 reductions, suggesting a 547 mechanistic link between sea ice and ocean circulation. 548 This paper has focused exclusively on sea ice as a driver of physical changes, but we 549 recognize that these changes in sea ice will be accompanied by multiple processes that interact 550 and compete with each other. Marzocchi & Jansen (2019) note that teasing apart the individual 551 components of CO2 fluctuations is complicated because of interactions between sea ice 552 capping, air-sea disequilibrium, AABW formation rates, and the biological pump. We recognize 553 that these processes may not act independently, and as such, hope to contribute new data to 554 help advance our collective understanding of the role of sea ice on influencing atmospheric CO2 555 variability on a glacial-interglacial time scale.    The first potential cause for the observed differences between TAN1302-96 and E27-23 582 wSIC estimates is through the cumulative effects of different laboratory protocols. While it is 583 difficult to determine precisely how much different laboratory protocols could influence the 584 results, we cannot exclude this explanation as a possible contributor to differences in wSIC. 585 The second potential cause for differences in wSIC estimates between E27-23 and 586 TAN1302-96 are differences in counting and identification methods. We believe this is an 587 unlikely cause for the differences observed between E27-23 and TAN1302-96 primarily because 588 of the magnitude of counting discrepancies required to cause a difference of 50% wSIC 589 estimates between the two cores. The close coupling of wSIC estimates between TAN1302-96 590 and SO136-111 over the entire glacial-interglacial cycle supports that a fundamental issue 591 relating to taxonomic identification and/or methodology is an unlikely explanation for the 592 observed wSIC differences. 593 Finally, the fourth potential cause of differing wSIC estimates is selective diatom 594 preservation (e.g., Pichon et al., 1999;Ragueneau et al., 2000). The similarities between 595 TAN1302-96 and SO136-111 wSIC estimates, along with independent indicators in cores E27-23 596 and TAN1302-96, suggest that this is unlikely. For E27-23, Bradtmiller et al. (2009) used the 597 consistent relationship between 231 Pa/ 230Th ratios and opal fluxes to suggest that dissolution 598 remained relatively constant between the LGM and Holocene periods. In TAN1302-96, we 599 assigned a semi-quantitative diatom preservation value between 1 (extreme dissolution) and 4 600 (virtually perfect preservation) for each counted specimen. The average preservation of 601 diatoms for the entire core was 3.38 ± 0.13, with no observed bias based on sedimentation rate 602 or MIS. This assessment, although semi-qualitative, suggests that preservation remained 603 relatively constant (and good) throughout TAN1302-96, and is therefore unlikely to cause large 604 differences in wSIC between the two cores.