Records, processes, and effects of sea-level change

One of the most societally-relevant objectives within the earth sciences is to understand the history and impact of global sea-level (eustatic) fluctuations at different timescales. Sea-level fluctuations result from changes in the volume of water in the ocean or the volume of ocean basins as a consequence of the complex spatial and temporal interplay of a spectrum of intertwined forcing processes. Reconstructions of the magnitude and timing of sea-level movements can therefore provide clues to the complex interactions between the cryosphere, the oceans, the atmosphere, and the lithosphere that drive global climate change (Fulthrope et al., 2008).

Over the past 100 m.y., sea-level change reflects global climate evolution from a time characterized by ephemeral Antarctic ice sheets (100 Ma to 33 Ma), to a time when large ice sheets occurred primarily in Antarctica (33 Ma to 2.5 Ma), and finally to a world with large Antarctic and Northern Hemisphere ice sheets (2.5 Ma to present) (Miller et al., 2005a) (Fig. 2.2). Over the last 2.5 m.y., the structure of the sea-level curve shows greater high-frequency fluctuations than does the insolation index. The lack of an overall correlation indicates that, in addition to summer insolation at 65˚N, other parameters have influenced global climate. By the time ice sheets became long-term features of the Northern Hemisphere landscape, reaching the continental shelves and becoming unstable, the subtle interplay between Milankovitch cycles, ice-sheet dynamics, and shifts in ocean circulation had begun to drive the climate system (Lambeck, 2004).

Eustatic changes on the 10⁴ to 10⁶ year scales are controlled primarily by the cyclic growth and decay of continental ice sheets. Over the past ~800 kyr, major rapid eustatic changes occurred at intervals of ~100 kyr with maximum amplitudes of 120–140 m (Fig.

2.3), and involved changes in ice volume of 50–60 million km³; lesser cycles of a few tens of thousands of years and shorter duration are superimposed on these. Sea-level fluctuations during the last glacial cycle, from about 130 kyr to present, have responded to the dominant oscillations in insolation, with periodicities of ~40 kyr and ~20 kyr, but the relative amplitudes and phase relationships show no consistent match.

The evolution of high-latitude global ice volumes, as inferred from observations of far-field sea-level change, serves as a fundamental constraint on climate models for the last 3 m.y. (Lambeck et al., 2002a). Although the correlation between ice and ocean volumes is incontrovertible, the causal link is commonly obscured. Local effects change the position of the sea surface relative to the land and impose additional signals that

overprint the record of global sea level. Apart from tectonic movements, they include a variety of processes resulting from ice-sheet unloading and the redistribution of water masses between ice sheets and the global ocean: glacio-hydro-isostatic adjustments, the changing gravitational potential of the ice sheets, and equatorial ocean-siphoning effects. The resultant relative sea-level change is therefore a function not only of the change in ice volume but also of the planet‘s rheology. A number of attempts have been

Figure 2.2 Global sea level (light blue) for the interval 7 Ma to 100 Ma derived by backstripping data. Revised backstripped sea-level curve (plotted in brown) is based on 7 new wells and new age and paleoenvironmental data for the 5 holes used to derive the light blue curve(Kominz et al., 2008). Global sea level (purple) for the interval 0 Ma to 7 Ma derived from a benthic foraminifer δ¹⁸O synthesis from 0 to 100 Ma (red) (Miller et al., 2005a). The Miller et al.

(2005a) backstripped sea-level curve was smoothed with a 21-point Gaussian convolution filter (black).

The pink box at ̀11 Ma is a sea-level estimate derived from the Marion Plateau (John et al., 2004). Light green boxes indicate times of spreading rate increases on various ocean ridges (Cande and Kent, 1992).

The dark green box indicates the opening of the Norwegian-Greenland Sea and concomitant extrusion of basalts (modified from Browning et al., 2008).

Figure 2.3 A Pleistocene stack of 57 globally distributed ¹⁸O records (from Lisiecki and Raymo, 2005).

made to model both global hydro-isostatic adjustments and equatorial ocean siphoning (e.g., Lambeck et al., 2002a; Milne et al., 2002; Peltier, 2002) to simulate the lithospheric response to specific deglaciation histories and predict the general shape of local sea-level curves; however, aspects of these models remain controversial and significant deviations between model predictions and coral-based sea-level records have been noted in several regions (e.g., Bassett et al., 2005). New records are therefore needed to constrain and validate rheologic models of the mantle that predict the Earth‘s response to surface loads.

The discrepancies between ice-sheet models and far-field sea-level records must be tested to improve our ability to model past ice-sheet behavior and dynamics and enhance predictions of future changes. Central to this aim is the requirement to obtain accurate records of sea-level change on a regional scale, at various latitudes, in different tectonic settings, and at variable distances from former glaciated regions. Observations from sites distal to glaciated regions (i.e., the ‗far-field‘) are less affected by isostatic deformation and therefore better suited to constrain glacial eustasy. In contrast, sea-level data from sites proximal to former ice sheets (i.e., the ‗near field‘) provide information on local ice sheet dynamics.

Quantitative studies of sea-level change and ice-sheet fluctuations coupled to global climate variability have not been undertaken so far. The range and rate of temporal and spatial variability of various ice sheets (e.g., ice extent, ice volume and contribution to global sea level, thermal conditions) and the sensitivity of various parts of the cryosphere to changes in Earth‘s climate (e.g., atmospheric greenhouse gas concentrations, sea-surface and land temperatures, orbital cycles) are poorly constrained, and thus the causes of the climate fluctuations that repeatedly built up and destroyed ice sheets remain unclear. In particular, the timing and magnitude of past rapid ice-sheet collapses need to be documented and the driving mechanisms clarified.

Ice-sheet responses for different climate backgrounds remain uncertain due to the difficulties of establishing reliable geophysical ice models that take into account the basal conditions of the ice sheets. The behavior of individual ice sheets needs to be investigated, including the thermal characteristics of Antarctic ice sheets during past

warmer-than-present climates and the response of the Antarctic ice sheets to orbital forcing.

Past eustatic variations can be estimated from shoreline markers, reefs and atolls, oxygen isotopes (¹⁸O), and the flooding history of continental margins and cratons. Oxygen-isotope values provide a proxy for glacioeustasy but the relationship between ¹⁸O and sea-level fluctuations is more complex than sometimes assumed due to uncertainties regarding the effects of temperature, diagenesis, and evaporation-precipitation processes on calcite ¹⁸O values. Tropical coral reefs are unique recorders of sea level and environmental changes and can provide unparalleled records of the timing and magnitude of Quaternary sea-level change by dating the ‗fossil sunshine‘ (i.e., shallow dwelling corals). They are therefore of pivotal importance to resolving the rates of millennial-scale eustatic changes, clarifying the mechanisms that drive glacial-interglacial cycles, and constraining geophysical models.

Because the magnitude of Pleistocene sea-level change was in the 120-130 m range, the relevant reef and sediment archives are mostly stored on modern fore-reef slopes and can therefore be investigated only by drilling. Direct, accurate, and high-resolution observations of coral reef-based Pleistocene and Holocene sea levels exist only for the last glacial cycle (from about 130 kyr ago to present) and are limited mainly to two snapshots: the period following the Last Glacial Maximum (LGM) and the last interglacial period, approximately 125 kyr ago. Existing coral-reef records are too limited to constrain accurately the Pleistocene sea-level fluctuations and therefore new drilling is necessary. Furthermore, most of the existing records concern uplifted and presently emerged parts of reefs and reef terraces in active subduction zones where vertical tectonic movements may be large and often discontinuous, implying that apparent sea-level records may be biased by variations in the rates of uplift. Hence, there is a clear need to obtain sea-level records in tectonically stable regions or in areas where vertical movements are slow and/or regular. The reconstruction of the evolution of high-latitude global ice volumes will primarily rely on far-field sea-level change during interglacial and glacial periods. Glacial terminations, which are critical to understanding ice-sheet dynamics and to determining the timing and volume of meltwater released during deglaciation events, are discussed in section 5.2.3.