Chapter 16. Ice Sheets and Climate: the Marine Geological Record

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CHAPTER ICE SHEETS AND CLIMATE: THE MARINE GEOLOGICAL RECORD 16 S. Passchier Montclair State University, Montclair, NJ, United States

16.1 PAST GLACIATIONS The Earth’s glacial record spans billions of years (Hambrey and Harland, 1981; Deynoux et al., 1994). The most recent icehouse phase is best preserved in the geological record and it has been studied intensely. High-latitude drilling has yielded ice-proximal evidence of polar glaciation through the investigation of stratigraphic records of past glacial environments (Hambrey et al., 1991; Strand et al., 2003; St. John, 2008). The presence of gravel-sized ice-rafted detritus (IRD) in the Arctic Ocean shows that Northern Hemisphere glaciers were present B46 Ma (St. John, 2008). In Antarctica, glaciation commenced on areas of high topography probably in the Late Eocene (Strand et al., 2003) with continental-scale glaciation expressed as progradation of continental margins and a shift in weathering regime near the EoceneÀOligocene boundary at B34 Ma (Hambrey et al., 1991; Ehrmann and Mackensen, 1992). The onset of polar glaciation in both hemispheres is tightly coupled to increases in heavier oxygen isotopes in the shells of benthic foraminifera (Fig. 16.1). The oxygen isotope geochemistry of tests of benthic foraminifera yield a proxy record of deep-sea temperature and ice volume with high temporal resolution using compilations of data from numerous drillholes collected by the Deep-Sea Drilling Project (DSDP), the Ocean Drilling Program (ODP), and the Integrated Ocean Drilling Program (IODP) (Zachos et al., 2001; Cramer et al., 2011; Veizer and Prokoph, 2015). The partitioning of the deep-sea cooling and ice volume signals in the deep-sea oxygen isotope records is a topic of active research using Ca/Mg ratios of foraminiferal calcite (Lear et al., 2000). Furthermore, observations using satellite data, coupled with observations at GPS stations on land and at sea, have allowed the development of sophisticated ice sheet and climate models that are applied to the ice sheetÀclimate interactions of the past (Marshall and Clark, 2002; DeConto and Pollard, 2003; Stocchi et al., 2013; Goldner et al., 2014). Correlations between direct evidence of glacial activity at high-latitude sites and the deep-sea stable isotope records allowed detailed investigations into ice sheets in the climate system, in particular for the past 5 million years (Lisiecki and Raymo, 2005). The pre-Pleistocene glacial record for the Arctic and the Antarctic, however, is still incomplete at high-latitude sites (St. John, 2008; Escutia et al., 2014).

FIGURE 16.1 Oxygen isotope compilations for benthic foraminifera. The main panel shows individual data points (grey dots), and five-point moving average (blue line) of measurements for the Cenozoic. EECO, Early Eocene Climatic Optimum; EOT, EoceneÀOligocene Transition; MCO, Miocene Climatic Optimum; Oi-1 and Mi-1 isotope shifts after Miller et al. (1987). For the inset of the Late Pliocene to present: blue line represents a 10-point moving average through the data. M2 and numbers are glacials and interglacials (following Shackleton and Opdyke, 1973; Lisiecki and Raymo, 2005). MPT, Mid-Pleistocene Transition, positioned between δ18O records with obliquity cycles below and short-eccentricity cycles above. Compilation from Veizer, J., Prokoph, A., 2015. Temperatures and oxygen isotopic composition of Phanerozoic oceans. Earth Sci. Rev. 146, 92–104, partially based on a previous compilation of Cramer, B.S., Miller, K.G., Barrett, P.J., Wright, J.D., 2011. Late Cretaceous–Neogene trends in deep ocean temperature and continental ice volume: reconciling records of benthic foraminiferal geochemistry (δ18O and Mg/Ca) with sea level history. J. Geophys. Res. (C12), 116. 16.2 ICE SHEETS IN THE EARTH SYSTEM 567

16.2 ICE SHEETS IN THE EARTH SYSTEM 16.2.1 ICE VOLUME AND SURFACE TEMPERATURE Ice sheets grow when snow accumulation exceeds melt and dynamic ice loss through calving. Due to the dome-shaped geometry of ice sheets, as ice grows above the equilibrium line altitude, a larger area becomes available for snow accumulation. As an ice sheet increases in areal extent it also reflects more incoming solar radiation and decreases heat absorption by the Earth’s surface. The ice-elevation and ice-albedo feedbacks allow ice sheets to grow relatively quickly early on in the glaciation (Fig. 16.2). Snow accumulation and dynamic ice mass loss are controlled by surface and shallow ocean temperature. Evaporation is enhanced from water bodies with higher surface temperatures and the larger capacity of warmer air to hold water vapour facilitates higher rates of precipitation. For marine-grounded glaciers and ice sheets, higher snowfall rates, however, may result in increased dynamic ice mass loss due to steeper surface gradients and greater ice thickness near the grounding line (Ivins, 2009). In Greenland calving has accelerated for some large outlet glaciers in recent years. As ice sheets advance into marine environments on continental shelves, ice shelves can form that buttress ice drainage. However, ice-shelf instability along the Antarctic Peninsula (Fig. 16.3) suggests that temperature thresholds exist for the survival of ice shelves (Cook et al., 2005). The release of the buttressing effects of ice shelves under continued warming is already resulting in increased dynamic ice mass loss in some drainage basins (Wouters et al., 2015). Today most of the dynamic ice mass loss in Antarctica, however, is not through iceberg calving. Instead, dynamic thinning is increased where ocean temperatures are high (Pritchard et al., 2009), and where ice mass loss appears to be dominated by basal melt rather than enhanced calving (Rignot et al., 2013).

FIGURE 16.2 Ice sheets in the Earth system. (A) Onset of ice growth as ice-elevation feedback dominates and with small crustal response due to small mass and short time of ice development. (B) Snow accumulation exceeds ablation and large ice mass forms and depresses the crust over time. (C) Ablation exceeds snow accumulation: marine ice-sheet instability due to increased thickness at the ice sheet grounding line, landward overdeepening, and marine incursion. Speed at which the grounding line may restabilize depends on crustal thickness and mantle viscosity, which affect the rate of crustal rebound. ELA, equilibrium line altitude. 568 CHAPTER 16 ICE SHEETS AND CLIMATE

FIGURE 16.3 Map of Antarctic subglacial topography overlain with ice surface elevation contours in red. Grey dots represent drillholes with pre-Holocene Antarctic palaeoclimate records acquired by deep drilling drilled by the Deep-Sea Drilling Project (DSDP), Integrated Ocean Drilling Programs (ODP and IODP), and the Antarctic Drilling Program (ANDRILL). (Ryan, W.B., Carbotte, S.M., Coplan, J.O., O’Hara, S., Melkonian, A., Arko, R., et al., 2009. Global multi-resolution topography synthesis. Geochem. Geophys. Geosyst. 10 (3)).

Warming results in a positive surface mass balance only as long as increased rates of snow accumulation are not counteracted by larger dynamic ice mass loss through melt and calving. Surface melt rates vary from year to year. In Greenland snow accumulation is currently offset by both surface melt and enhanced calving. In contrast, studies of ice cores in continental Antarctica show increased long-term snow accumulation in response to regional warming in the near absence of surface melt. Modelling studies, however, suggest that a large proportion of the mass gain from increased snow accumulation expected in a warming scenario is counteracted by increased dynamic 16.2 ICE SHEETS IN THE EARTH SYSTEM 569

ice mass loss (Winkelmann et al., 2012). Thus, despite the increased snow accumulation, the Antarctic ice sheet is not gaining mass. These complex interactions of atmospheric, oceanographic, and glaciological processes, however, only partially force globally distributed sea level changes. Solid Earth processes also play a significant role (Fig. 16.2).

16.2.2 ICE SHEETS AND RELATIVE SEA LEVEL Sea level is defined in several ways. Mean sea level is determined by the geoid, a hypothetical surface that is influenced only by the Earth’s gravitational attraction and rotation, whereas eustatic sea level has a fixed point within Earth’s interior as a reference point (e.g., the centre of the Earth). However, sea level is generally measured locally at tide gauges, or as strandlines of past sea levels. What is measured is relative sea level: the sea level with respect to a local datum. Ice sheets influence relative sea level in multiple ways. The most important effects ice caps or ice sheets have on sea level are related to the water they hold, and their mass: (1) the water in ice can change from the solid to the liquid phase; (2) the mass of ice deforms the solid earth; and (3) the mass of ice exerts a gravitational attraction on the ocean surface, similar to how the Moon and the Sun produce tides. The magnitude of each of these effects varies across the globe and it matters where ice was at what time. In the near field, near an ice sheet or ice cap, the processes of solid Earth deformation and gravitational pull dominate the relative sea level change. Observations of palaeo strandlines along the coast of the Bothnian Bay, near the former centre of the Scandinavian Ice Sheet, show magni- tudes of relative sea level rise of around 250 m since the last glaciation (Lambeck et al., 2001). In the far field, at thousands of kilometres away from an ice sheet, the hydrostatic effects may dominate. Even there, the addition of water to the ocean basins and the effects of the mass of that extra water may lead to deformation of the Earth’s surface by tens of metres. The amount of solid Earth deformation in both the near field and far field depends, besides on the location and melt rates of ice caps and glaciers, on the thickness of the lithosphere and mantle viscosity, which vary locally. The variations in the solid earth and these glacio-hydrostatic effects determine the global geographic distribution of sea level changes (Lambeck et al., 2014). Upon deglaciation, some regions experience relative sea level rise above the global average, while others, typically in the near field, stay well below it. For example, in response to a collapse of the West Antarctic Ice Sheet, relative sea levels on the heavily populated east coast of the United States could rise significantly more than the global average (Bamber et al., 2009). In contrast, loss of the Greenland Ice Sheet might decelerate relative sea level rise in northwestern Europe. Glacio-isostatic processes are also considered in situations of marine ice-sheet instability. Both the West Antarctic Ice Sheet and portions of the east Antarctic Ice Sheet are grounded well below sea level on a landward sloping bed. According to Weertman’s Law (1974), ice retreat over a landward sloping bed could result in marine ice-sheet instability. Recent modelling studies incorporat- ing glacio-isostatic processes, however, suggest that surface rebound diminishes the bed slope in response to ice mass loss (Gomez et al., 2010). Crucial, however, are differences in the timescales of ice retreat and changes in the dynamic topography (Boulton, 1990). Modelling requires data on lithospheric thickness and mantle viscosity, and calibrations to past ice surface elevation, ice thickness and areal extent, and rates of ice retreat. 570 CHAPTER 16 ICE SHEETS AND CLIMATE

16.3 RECONSTRUCTIONS OF ICE EXTENT AND ICE THICKNESS 16.3.1 ICE-CONTACT DEPOSITIONAL RECORDS Because of the globally distributed geodynamic effects on sea level it is important to know the past geographic distribution and thickness of an ice mass. Marine geological and geophysical studies of continental margins and on land reconstruct the geomorphological footprints of ice sheets during the last major glaciation (Boulton et al., 2001; Anderson et al., 2002; Stokes and Tarasov, 2010). On land, surface exposure ages have been widely used to study the development of ice thickness and extent. Using surface exposure dating, Balco and Rovey (2010) reconstructed changes in the extent of the Laurentide Ice Sheet through the Late Pliocene and Pleistocene and found that glaciation intensified from c.1.3 Ma. Furthermore, a compilation of continent-wide surface exposure ages suggests that more than 4 Ma ago the East Antarctic Ice Sheet was periodically thicker than present during glacials and that it thinned asynchronously through the Late Pliocene and Pleistocene (Yamane et al., 2015). Thus, the distribution of ice mass and its effect on the dynamic topography in the past were widely different from today with implications for globally distributed sea level models that include dynamic topography (Raymo et al., 2011). Prior to the last major ice expansion, however, much geomorphological evidence has been oblit- erated by subsequent erosion. Therefore, offshore stratigraphic records of ice extent are interpreted through combinations of seismostratigraphy and drilling (Cooper et al., 1991; Vorren and Laberg, 1997; O’Brien et al., 2007). For example, trough-mouth fans have been recognized as powerful ice-sheet monitors (Vorren and Laberg, 1997). These fans are found on the continental slopes in both hemispheres seaward of large glacial troughs. Trough-mouth fans are composed of glacigenic debrites recording the maximum extent of ice streams, interbedded with hemipelagic muds deposited during periods with reduced ice sheets. The sand fraction of glacigenic sediments has been used to assess the past positions of areas of enhanced glacial erosion, which are typically found under the margins of ice sheets (Boulton, 1996; Passchier, 2007). For example, the sand provenance of glacio-marine diamicts in the Antarctic AND-2A drillhole yielded evidence that the areas of glacial erosion shifted landward, implying that the East Antarctic Ice Sheet retreated during the Miocene Climatic Optimum (Hauptvogel and Passchier, 2012). Furthermore, sand provenance studies at IODP Site U1358 off the Wilkes Subglacial Basin (Orejola et al., 2014) are in agreement with the surface exposure ages of a thicker and more advanced East Antarctic Ice Sheet during glacials more than 4 million years ago and a thinning in the Late Pliocene to Pleistocene as discussed earlier (Yamane et al., 2015). Geomorphological maps and continental margin records are essential in providing boundary conditions for ice sheetÀsea level model reconstructions. These records of ice thickness and extent, however, are usually of poor chronological resolution. Palaeorecords of orbital scale ice sheet variability are typically collected offshore, where more complete records with fewer hiatus and biomagnetostratigraphic chronologies allow for higher-resolution correlations of changes in ice sheets, climate, and sea level. The interpretation of offshore records, however, is ambiguous without the context of the land-based evidence, because of the interference of oceanographic processes. 16.3 RECONSTRUCTIONS OF ICE EXTENT AND ICE THICKNESS 571

16.3.2 ICE EXTENT FROM DEEP-SEA MUDS ON CONTINENTAL MARGINS Clay mineralogy and geochemical methods are employed in deep water stratigraphic records to assess the onset of glaciation and ice extent. On a continental margin the initiation of full-scale glaciation of a cratonic region is recognized by (1) an increase in sediment supply to the continental slope and rise, (2) a change in the weathering state of fine-grained materials, caused by the mechanical erosion by the ice sheet itself, and (3) the aridifying impact of katabatic winds on proglacial weathering environments (Passchier et al., 2013a). A well-documented example is the initiation of continental-scale Antarctic glaciation at the EoceneÀOligocene boundary. It is characterized by a marked shift to physically weathered illite-chlorite-dominated clay composition of Southern Ocean sediments, accompanied by a shift in the δ18O record (Ehrmann and Mackensen, 1992). Similar records of clay mineral shifts have been published for the Northern Hemisphere (e.g., Andrews, 1993; Piper et al., 1994). Novel methods of sediment provenance analyses applied to the detrital mud fraction of deep-sea sediments offshore ice centres are capable of more precisely tracking ice extent through glaciations. Colville et al. (2011) analysed neodymium isotopes in the silt-sized detrital fraction of drift deposits offshore and showed that the southern tip of Greenland was ice-free during the last interglacial: glacial rock flour was supplied only from terrains farther to the north. A similar study of Site U1361 off East Antarctica showed that an ice margin that had retreated supplied sediment from a position within the Wilkes Subglacial Basin during the interglacials of the warm early-mid Pliocene (Fig. 16.3; Cook et al., 2013).

16.3.3 RECORDS OF IRD In order to unravel the mechanisms of deglaciation on glacialÀinterglacial timescales, and to test models driven by proxy data, time series analyses of higher-resolution palaeoclimate data sets are required. Records of IRD are employed to assess variability in the cryosphere at the same time resolution as ice distal marine palaeoclimate proxies, such as oxygen isotope records. The maximum possible temporal resolution depends on linear sedimentation rates, which vary typically between 1 and 10 cm per 1000 years in high-latitude continental rise to abyssal plain settings. Ideally, IRD counts or abundances should be normalized with respect to background sedimentation rates and dry bulk density (Allen and Warnke, 1991; Krissek, 1992). IRD abundance counts are commonly based on the size fraction between sieve diameters of 0.25À2.00 mm or 0.150À2.00 mm. A 0.150 mm sieve diameter is equivalent to a volumetric grain diameter of 0.125 mm using a laser particle sizer (Passchier, 2011; Hansen et al., 2015). The interpretation of IRD records is complex, and depends strongly on the chronological sam- ple resolution, and the palaeo-oceanographic and palaeoglaciological context. On millennial and longer timescales, at least five variables influence ice-rafting fluxes: (1) ice debris content, (2) ice discharge, (3) iceberg size, (4) sea surface temperature (SST), and (5) surface currents. Glacial erosion rates are a strong influence on iceberg debris content. Glacial erosion rates and ice discharge, however, appear to be dependent variables, because fast flow enhances glacial erosion rates (Herman et al., 2015). Hence high accumulation rates of ice-rafted debris in 572 CHAPTER 16 ICE SHEETS AND CLIMATE

ice-proximal settings could be indicative of high ice discharge. Likewise, SSTs and ocean currents are both constrained by the position of oceanic fronts (Bond et al., 1997; Death et al., 2006): SSTs control iceberg survival and IRD sedimentation rates, and the strength and directions of ocean currents control iceberg tracks. Traditionally, records of IRD, in particular Heinrich events, have been interpreted as resulting from largely glaciologically controlled bingeÀpurge behaviour of an ice sheet (MacAyeal, 1993). Heinrich events are ice-rafting episodes dating from the last glacial (70À10 ka) in the North Atlantic that are spaced about every 5À10 ka. In the model, the binge phase is characterized by growth of the Laurentide Ice Sheet over a frozen bed. As the ice sheet thickens, the geothermal heat flux to the bed increases leading to basal melt. During the purge phase ice is discharged through ice streams over a lubricated bed. Sediment core analyses from ice-proximal sites (Box 16.1), however, suggest that meltwater discharge accompanied the growth phase of the ice sheet and that the ice sheet was not frozen to the bed (Rashid et al., 2012). More recently, based on satellite observations, iceberg calving under a warming scenario is modelled in two major ways: (1) increased snow accumulation, forcing more ice across the grounding line (Winkelmann et al., 2012); and (2) ice-shelf disintegration, with additional diminishing of the buttressing effect on outlet glaciers (Hulbe et al., 2004). Modelling suggests that, upon warming, ice discharge forced by increased snow accumulation for any basin is larger than that of ice- shelf disintegration (Ivins, 2009; Winkelmann et al., 2012). Furthermore, palaeorecords suggest that both mechanisms operate within the same basin across a warming trend (Hansen et al., 2015). The implication is that marine ice sheets exhibit threshold behaviour and change their mode of calving and ice loss depending on the climate state. IRD peak abundances show diachronous distributions between proximal sites near the source of land ice and distal sites in the open ocean (Box 16.1), and the timing of peak IRD sedimentation rates cannot likely be assumed constant across large ice-rafting distances (Dowdeswell et al., 1999; Rashid et al., 2012). IRD peak abundances in distal settings are strongly modulated by SST and ocean currents, and may be unrelated to ice discharge. For example, both millennial-scale and Heinrich event peak IRD abundances in the central North Atlantic coincide with cold SST conditions (Bond et al., 1997; Chapman and Shackleton, 1998). Coupling of IRD with proximal continental margin records is therefore essential, because these records provide unique evidence of the extent and behaviour of grounded ice sheets (Box 16.1). In proximal high-latitude settings, adjacent to large outlet glaciers, IRD maxima record increased ice discharge related to the effects of increased erosion and ice debris content during episodes of fast flow (Dowdeswell et al., 1999; Death et al., 2006; Passchier, 2011; Rashid et al., 2012; Weber et al., 2014; Herman et al., 2015). On glacialÀinterglacial timescales, proximal records show that IRD maxima are recorded at the onsets of interglacials (Passchier, 2011; Weber et al., 2014), in agreement with other proximal continental margin records of dynamic behaviour and fast ice discharge at the onset of deglacial times (O´ Cofaigh et al., 2005; Nygard et al., 2007; Passchier et al., 2010). Upon glacial termination, ice streams typically widen and discharge more ice (Stokes and Tarasov, 2010).Theincreaseinicedischargelateinaglacia- tion may also be tied to the basal temperature evolution of an ice sheet, with warm-based ice and basal sliding developing as an ice sheet thickens, followed by a rapid termination (Marshall and Clark, 2002). BOX 16.1 HEINRICH EVENTS IN SEDIMENTARY RECORDS AND MODELS

Top: Simplified facies distribution for Heinrich layers found in sediment cores from ice-proximal and distal sites (near mid-ocean ridges). Early calving and deposition of ice-rafted debris (IRD) are associated with meltwater discharge (Facies A and B), followed by calving and deposition of IRD (Facies C). Bottom: Heinrich events as they are represented in models that use data from cores retrieved from the distal sites only. LIS, Laurentide Ice Sheet; EIS, Eurasian Ice Sheet. From Rashid, H., Saint-Ange, F., Barber, D.C., Smith, M.E., Devalia, N., 2012. Fine scale sediment structure and geochemical signature between eastern and western North Atlantic during Heinrich events 1 and 2. Quat. Sci. Rev. 46, 136À150. Data sets from mid-ocean ridge flanks above the calcite compensation depth are favored because their foraminiferal carbonate content allows for the correlation of IRD content with stable isotope analysis. The IRD accumulation rates at these distal sites, however, are strongly influenced by iceberg survival controlled by sea surface temperature (SST). Analyses of ice-proximal records of ice-rafting and meltwater discharge are necessary, because the SST effect interferes with the ice discharge signal of the IRD at the distal sites. 574 CHAPTER 16 ICE SHEETS AND CLIMATE

16.4 ORBITAL FORCING OF GLACIATION AND CLIMATE FEEDBACKS 16.4.1 ONSET OF MAJOR NORTHERN HEMISPHERE GLACIATION Gravel-sized IRD as old as B46 Ma is found in stratigraphic records of the central Arctic Lomonosov Ridge at IODP Site 302 (Fig. 16.4; St. John, 2008). Following on from the previous discussion: whereas the IRD indicates that icebergs were calving into the Arctic Ocean and survived long-distance transport as early as the mid-Eocene, the IRD is not evidence of continental-scale glaciation. Due to the presence of perennial sea ice, the Arctic basin has been a difficult target for drilling palaeoclimate archives that record pre-Pleistocene glaciations. The record of Pleistocene glaciations is much better known, but questions remain regarding the onset and intensification of major Pleistocene Northern Hemisphere glaciation between 2.8 and 0.8 Ma. For studies of the Pliocene and Pleistocene Northern Hemisphere ice sheets, IRD records have typically been collected from mid-ocean ridge flank settings, where the seafloor emerges above the calcite compensation depth (CCD). Below the CCD, ocean water is undersaturated with respect to calcium carbonate due to the cold temperatures and high hydrostatic pressure at depth. Carbonate skeletal material secreted by organisms in the water column dissolves as it settles below the CCD. Although records from mid-ocean ridge flanks above the CCD allow comparison between stable isotope records in foraminiferal calcite and ice-rafted debris, these sites are typically positioned far from glaciated continental margins and preserve a signal of iceberg survival. In the past, Maslin et al. (1992) suggested that an increase in IRD at 2.78 Ma at Site 882 in the north Pacific (Fig. 16.4) marked the onset of full-scale Northern Hemisphere glaciation, despite a predominantly northern and local origin for the IRD (McKelvey et al., 1992). In the Arctic Ocean, deposition of IRD rich in detrital carbonate derived from North America is common since 2.8 Ma, but peak quan- tities of the IRD occur in interglacials, including the Holocene (Spielhagen et al., 1997). In both cases the IRD was derived from local glaciers. Therefore, although cooling occurred from 2.8 to 2.7 Ma onward, IRD records do not provide convincing evidence of continental-scale glaciation in North America and these early records are now commonly regarded as a signal of increased iceberg survival in response to cooling (Bailey et al., 2013). The implication is that continental-scale glaciation commenced later and that more ice-proximal continental margin evidence is needed to estab- lish changes in ice extent.

16.4.2 INTENSIFICATION OF NORTHERN HEMISPHERE GLACIATION B1MA IRD records need to be reconciled with evidence from continents and continental margins. These records have shown evidence of some glacial activity by smaller marine-based ice sheets between 2.8 and 1.5 Ma with glacial intensification to continental ice sheet proportions between 1.5 and 1.1 Ma. In northern Eurasia, the early Pleistocene ice sheets at the higher latitudes, such as the Barents Sea Ice Sheet, were small marine-based ice sheets. ODP Site 986 (Fig. 16.4) records major glacial expansion with ice extending to the shelf break on the SvalbardÀBarents Sea margin between 1.5 and 1.3 Ma (Butt et al., 2000). At lower latitude, the oldest recorded glacial advance by the Scandinavian Ice Sheet onto the Norwegian Shelf is dated at 1.1 Ma at the Troll 8903 borehole (Sejrup et al., 2000). Trough-mouth fan sedimentation with glaciers crossing the shelf to the shelf break was established by the Mid-Pleistocene, 16.4 ORBITAL FORCING OF GLACIATION AND CLIMATE FEEDBACKS 575

FIGURE 16.4 Large ice sheets of the Northern Hemisphere during the Last Glacial Maximum (LGM). LGM reconstruction from Ehlers and Gibbard (2004a,b). Map created using GeoMap App (Ryan et al., 2009). Grey dots represent deep drillholes of the DSDP, ODP, and IODP programs and numbers refer to site numbers mentioned in the text. 576 CHAPTER 16 ICE SHEETS AND CLIMATE

B0.9 Ma (Vorren and Laberg, 1997).By0.7Magroundedicehadalsoreachedthenorthern Eurasian continental margin and began to supply sediment eroded from the shelf to the central Arctic basins (Spielhagen et al., 1997). In North America, chemically weathered till-like deposits with abundant sedimentary clasts are interpreted as marking the onset of glaciation more than 2 million years ago, whereas unweathered tills were largely deposited in the Brunhes Chron, after 0.78 Ma (Roy et al., 2004). Furthermore, based on surface exposure data, Balco and Rovey (2010), found evidence for an isolated erosional event at 2.4 Ma, but reconstruct major advances of the Laurentide Ice Sheet only from about 1.3 Ma. At high latitude, a persistent increase in sedimentation rates from 1.2 Ma onward is observed at IODP Site U1417 on the southern Alaska margin (Gulick et al., 2015). Based on palaeomagnetic measurements, Barendregt and Irving (1998) argued that separate Cordilleran, Keewatin, and Labrador Ice Sheets existed prior to the BrunhesÀMatuyama reversal (0.78 Ma), whereas a continent-wide Laurentide Ice Sheet was present after this time. Evidence of full-scale continental glaciation of eastern North America with ice reaching the coast is represented by a shift in sediment provenance between B1 and 0.7 Ma at Ocean Drilling Program Site 645 in Baffin Bay (Fig. 16.4), between Canada and Greenland (Thiebault´ et al., 1989; Andrews, 1993). A change in clay mineralogy to illite-chlorite dominated assemblages illustrates the change in sediment supply caused by the initiation of physical erosion of crystalline basement by a thick continental-scale Laurentide Ice Sheet (Clark and Pollard, 1998). Off eastern Canada, in the Grand Banks area, no glacial activity was recorded prior to Marine Isotope Stage (MIS) 24 at B1 Ma, followed by local upland glaciation (Piper et al., 1994). Around MIS 17 (B0.7 Ma) an increase in illite and plagioclase is observed, followed by an unusual high rate of sedimentation reflecting the development of a full continental-scale ice sheet (Piper et al., 1994). Coincidentally, McHugh et al. (2002) report increased debris flow activity from MIS 18 (B0.8 Ma) on the New Jersey margin at ODP Site 1073. These records demonstrate that glaciation by thick continental-scale Eurasian and Laurentide Ice Sheets, which extended to the shelf break at mid-latitudes, commenced between B1 and 0.7 Ma.

16.4.3 ORBITAL FORCING OF ICE DYNAMICS High-resolution palaeoclimate records show a strong forcing by orbital parameters (Box 16.2), but shifts in the frequencies of orbital forcing are still not completely understood. One of the most enigmatic phe- nomena is the shift from 41-kiloyear obliquity cycles to 100-kiloyear eccentricity cycles in palaeoclimate records across the Mid-Pleistocene Transition (MPT), 1.1À0.8 Ma (Fig. 16.1). Obliquity cycles in snow accumulation are explained by the strong effect of the tilt of the Earth’s axis on the seasonality of insolation and moisture supply to ice sheets from lower latitudes (Vimeux et al., 1999), whereas B20 ka precession cycles are a strong driver of summer insolation at higher latitudes (Hansen et al., 2015). Both deep-sea stable isotope records and high-latitude marine palaeoclimate records, however, show eccentricity cycles, even prior to the onset of major Northern Hemisphere glaciation (Naish et al., 2001; Lisiecki and Raymo, 2005; Pa¨like et al., 2006; Passchier, 2011; Liebrand et al., 2011; Passchier et al., 2013b). The strong eccentricity response of palaeoclimate records is surprising given the relatively weak effect of changes in the shape of the orbit on insolation at high latitudes. Furthermore, for the past 5 million years (Lisiecki, 2010) low eccentricity of the orbit with small changes to its shape coin- cides with a large 100-ka response in stable isotope geochemistry of benthic foraminifera. Low eccentricity forcing favors warm interglacials, such as MIS 11, whereas peak eccentricity forcing generates colder and less pronounced interglacials, such as MIS 7 (Fig. 16.1). 16.4 ORBITAL FORCING OF GLACIATION AND CLIMATE FEEDBACKS 577

BOX 16.2 ORBITAL CYCLICITY

The Earth’s orbit around the Sun has an elliptical shape and the Sun is off-centre within the orbit. The shape of the orbit changes every B100,000 years from more to less elliptical, which influences the distance between the Earth and the Sun. Above is a ‘warm’ summer configuration for the Northern Hemisphere. The influence of eccentricity alone on the amount of incoming solar radiation (insolation) received, however, is weak. In contrast, variations in the tilt angle of the Earth’s axis have a large influence on changes in the distribution of insolation across the globe. Without the tilt of the Earth’s axis seasons would not exist. The fact that the Earth’s axis is tilted determines that the Northern and Southern hemispheres receive maximum insolation in different seasons as the Earth completes its annual rotation around the Sun. The tilt angle increases and decreases from about 22À24½ degrees on B40,000-year obliquity cycles, affecting the seasonal contrast of surface temperatures and temperature gradients across the globe. The precession cycle is caused by a ‘wobble’ in the orientation of the Earth’s axis. As the Earth rotates around the Sun the orientation of its axis shifts slowly each year relative to its position within the orbit (close or far away from the Sun). One complete cycle takes around 20,000 years. Precession cycles affect the insolation received at high latitudes during the summer melt season.

Lisiecki (2010) suggested that the strong 100-ka cycles in climate records from times of weak 100- kiloyear forcing could be a result of weak precession. Because precession is a strong driver of summer insolation and melt at high latitudes, weak precession potentially favors long-term climate variables, such as ice sheets, or the carbon cycle, to generate large responses. Modelling experiments for the Late Pleistocene Northern Hemisphere ice sheets show that a reduction in eccentricity diminishes insolation maxima (Box 16.1), allowing ice sheets to grow (Abe-Ouchi et al., 2013). As ice sheets grow, the ice- elevation and ice-albedo feedbacks (Fig. 16.2) further diminish the effect of insolation forcing. Within a prescribed set of climate boundary conditions, ice caps coalesce into large continental ice sheets that are not able to retreat to full interglacial extent at precession and obliquity timescales (Bintanja and Van de Wal, 2008). Thick, wet-based ice sheets are more susceptible to ice-sheet instability at peak insolation and display rapid terminations roughly at 100-kiloyear frequency as eccentricity increases (Clark and Pollard, 1998; Marshall and Clark, 2002; Abe-Ouchi et al., 2013). These inferences from modelling are consistent with the geological record for the past 5 Ma. Reduced glacial extent is correlated to low 100-ka eccentricity power at 4.2À3.6 Ma in the deep-sea proxy (Lisiecki, 2010). This period predates major Northern Hemisphere glaciation. A high-resolution IRD record gives evidence of precession forcing of marine ice-sheet instability in East Antarctic subglacial basins at 4.2À3.6 Ma (Hansen et al., 2015). Furthermore, widespread ice 578 CHAPTER 16 ICE SHEETS AND CLIMATE

retreat is documented by low IRD deposition at Antarctic continental margin Site 1165 (Passchier, 2011), condensed sections and episodes of ice retreat landward of current interglacial positions at IODP Sites U1359 and U1361 (Tauxe et al., 2012; Cook et al., 2013), and pelagic sedimentation at drillsite AND-1B on Antarctica’s continental shelf (Naish et al., 2009). In contrast, the 100-ka eccentricity power in the LR04 benthic oxygen isotope stack (Lisiecki, 2010) increased stepwise, first at 3.6À3.3 Ma leading up to a prominent excursion in the oxygen isotope records known as the M2 ‘superglacial’ (Fig. 16.1), and secondly at 1.1À0.8 Ma during the MPT. These periods are associated with the onset of ice-sheet advance in Antarctica and the Northern Hemisphere, respec- tively (O’Brien et al., 2007; Passchier, 2011; Balco and Rovey, 2010). In the past 5 million years, it is within these cold and expanded ice sheet states of the climate system that the 100-kiloyear eccentricity response of the benthic isotope stack is stronger than expected. The mechanisms behind the changes from a warm to a cold climate state and vice versa, however, remain poorly understood and may involve albedo and carbon cycle feedbacks among others, as discussed in Section 16.4.4.

16.4.4 GREENHOUSE RADIATIVE FORCING AND ALBEDO FEEDBACKS Deep-sea proxy records suggest that the onset of East Antarctic glaciation B34 Ma ago coincided with a drop in atmospheric pCO2 (Pagani et al., 2011). A reduction in greenhouse gases in the atmosphere may have contributed to high-latitude cooling and allowed the growth of a continental ice sheet in Antarctica. The role of greenhouse gases in the glacial onset, however, is still debated, because of a stepwise pattern of ice growth and cooling (Coxall et al., 2005). Likewise, the role of greenhouse radiative forcing in the dynamical changes of the MPT B1 Ma is not fully resolved. The changes in pCO2 across the MPT are relatively small, whereas the glacial effects are large, suggesting a role for albedo feedbacks (Ho¨nisch et al., 2009; Van de Wal and Bintanja, 2009). Ice growth increases surface albedo as white surfaces reflect incoming solar radiation. Analyses of differences between Pliocene and Late Pleistocene palaeorecords confirm that the albedo feedbacks associated with glaciation are more important than previously thought and that they significantly amplify the Earth system sensitivity to the radiative forcing of greenhouse gases (Mart´ınez-Bot´ıetal.,2015). It implies that, at least on long timescales, deglaciation is a positive feedback that amplifies the warming effects of greenhouse gas radiative forcing, and that it is not counteracted by larger negative feedbacks. As concentrations of atmospheric greenhouse gases increase, surface temperatures increase that allow ice to melt. As ice melts, surfaces that were once reflecting solar radiation are replaced with bare land or ocean that begin to absorb solar radiation and heat up. This is a positive feedback loop as the deglaciation amplifies the warming initiated through the increase in atmospheric greenhouse gases.

16.5 ABRUPT CLIMATE CHANGE AND RATES OF SEA LEVEL RISE Levels of greenhouse gases in the atmosphere and surface temperatures are currently rising rapidly. For short timescales into the future, however, the effects of storage of heat and carbon in the ocean, the effects of ocean warming on marine ice sheets, the effects of meltwater fluxes on ocean circulation and the expected rates of sea level rise remain poorly constrained. Much of what is known about abrupt climate change and the maximum rates of ice decay and sea level rise is derived from modelling constrained by numerous palaeoclimate records of the last deglaciation, B19 ka to present 16.5 ABRUPT CLIMATE CHANGE AND RATES OF SEA LEVEL RISE 579

FIGURE 16.5 Varves in a section of IODP Hole 347-64D Core 3 from the Baltic Sea. Data from Andren,´ T., Jørgensen, B.B., Cotterill, C., Green, S., Andren,´ E., Ash, J., et al., 2015. Site M0064. In: Andren,´ T., Jørgensen, B.B., Cotterill, C., Green, S., the Expedition 347 Scientists (Eds.), Proc. IODP, 347: College Station, TX (Integrated Ocean Drilling Program). http://dx.doi.org/10.2204/iodp.proc.347.108.2015 (Andren´ et al., 2015). The lightÀdark annual couplets accumulated in a glacial lake. The light layers are silt-rich summer meltwater deposits, and the dark layers are clay-rich winter deposits from gravity settling.

(Clark et al., 2012). Ice cores are palaeoclimate archives used to reconstruct Late Pleistocene atmospheric conditions. However, glacially influenced sediment archives, in particular varves (Fig. 16.5), are emerging as powerful monitors of how ice sheets responded to atmospheric forcing on annualÀmillennial timescales (Rittenour et al., 2000; Ridge et al., 2012). Glacial varves consist of sand/ silt and clay couplets, with sand/silt laminae representing discharge during the melt season, and clay laminae representing particulate settling in winter. Varve couplet thickness can be correlated to known rates of ice retreat (Ridge et al., 2012). Feedbacks between ice sheet melt, atmospheric greenhouse gas concentrations, and ocean circulation are likely responsible for episodes of abrupt climate change, such as the Older Dryas, BøllingÀAllerød, and Younger Dryas events B17À11 ka. Furthermore, the abrupt climate changes were asynchronous between the Northern and Southern Hemispheres. Following the onset of deglaciation, the Older Dryas cold event began as a meltwater pulse associated with Heinrich Event 1 released from the Northern Hemisphere ice sheets that disrupted the Atlantic Meridional Overturning Circulation. Freshening of the North Atlantic by the meltwater release decreased North Atlantic Deep Water (NADW) formation, with possible widespread implications for other compo- nents of the climate system. Cooling in the north may have pushed atmospheric frontal systems south, increasing atmospheric greenhouse gas concentrations as a result of enhanced wind-driven upwelling in a warmer Southern Ocean (Anderson et al., 2009). At B15 ka, a rapid increase in radiative forcing from greenhouse gases coincided with marine ice-sheet instability in Antarctica and an increase in the rate of sea level rise to a maximum of B4 m per century (Fairbanks, 1989; Clark et al., 2012; Weber et al., 2014), compared to the current average rate of B0.3 m per century. Despite advances in data collection of ice sheets and climate the exact mechanisms and climate feedbacks that control the behaviour of large terrestrial and marine-based ice sheets, and on what timescales, remain uncertain. The spatial coverage of high-latitude palaeoclimate archives is sparse (Figs. 16.3 and 16.4) and the available records are incomplete. Likewise, the effects ice sheets have on ocean circulation and the global carbon cycle through albedo, oceanic carbon storage, and weathering feedbacks under conditions of higher levels of atmospheric CO2 are poorly known. To advance these questions, future developments are focused on integrating ice sheet models within Earth System Models and drilling of high-resolution palaeoclimate archives of ice-sheet and sea level change. 580 CHAPTER 16 ICE SHEETS AND CLIMATE

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