Glacioeustasy and glacioisostasy

The basic idea of isostasy is that the Earth's crust, with a mean density of approximately 2800 kg m~3, is floating on the underlying plastic mantle with a mean density of about 3300 kg m~3 The amount of crustal depression resulting from ice sheet loading is a function of ice thickness and the ratio between the densities of ice and rock. The density of ice is about a third that of the crust, and therefore the crustal depression beneath an ice sheet is about a third of the ice thickness. Normally, the amount of crustal depression increases from the margin towards the centre of the ice sheet, where the ice sheets in most cases are thicker. The marginal depression, however, continues up to 150-180 km beyond the margin of the ice sheets. Therefore the sea can transgress proglacial areas.

The growth and melting of glaciers and ice sheets have a significant effect on global sea-level, causing large regional and global sea-level changes during glacial-interglacial cycles. A relative rise (transgression) or fall (regression) in sea-level caused by glaciers and ice sheets can occur by glacio-eustasy, glacio-isostasy, hydro-isostasy or geoidal eustasy.

Superimposed on long-term trends in sea-level change caused by tectonic activity, changes in mass distribution and the shape of the Earth, changes in the volume and mass of the hydrosphere, and the effects of variations in the rate of rotation or in the axis of tilting of the Earth, are major sea-level oscillations due to expansion and contraction of the ice sheets. During expansion of terrestrial ice sheets, water is extracted from the oceans. During the last glacial maximum, the terrestrial volume of the ice sheets produced a glacio-eustatic (controlled by the growth and contraction of ice sheets) sea-level lowering of about 120-130 m. In contrast, melting of the Greenland and Antarctic ice sheets would cause sea-level rises of 5.5 and 60 m, respectively (total glacio-eustatic sea-level rise of about 70 m). In some of the tectonically stable areas of the world, the glacio-eustatic sea-level variations have been reconstructed. In Bermuda (a stable mid-oceanic carbonate platform), uranium-series and amino-acid dating of corals and spe-leothems from fossil coral reefs and beach deposits have been used to reconstruct eustatic sea-level changes during the last 250,000 years

- Present sea-level

_ Bermuda " Platform rr

Figure 6.1 The Bermudan record of Late Pleistocene sea-level fluctuations. (Adapted from Harmon et al, 1983)

(Fig. 6.1). According to the Bermuda record, only on two occasions during that period has sea-level been higher than at present; at approximately 200,000 yr bp (+2 m) and during the last Eemian interglacial (+5-6 m).

Due to the limitations of local sea-level records, oxygen isotope records from deep ocean sediments have been used to reconstruct sea-level changes. The oxygen isotope record provides a proxy for ice-sheet volume and glacio-eustatic sea-level change (e.g. Shackleton, 1987). Although the curves of oxygen isotope variations through time reflect ice sheet and ocean volume changes, however, absolute changes in water depth are difficult to calculate.

Glacio-eustatic sea-level changes are most reliably recorded for the last 13,000-15,000 yr bp, when the Laurentide, Fennoscandian and British ice sheets retreated and finally melted. In addition, the Antarctic and Greenland ice sheets decreased in volume. A record based on high-precision U-Th dating of fossil corals on Barbados (Fig. 6.2) shows that sea-level was around 121 ± 5 m lower than at present during the last glacial maximum, and global sea-level

Glacial Sea Level

Figure 6.2 Barbados sea-level curve for the last 17,000 years based on U-Th dating of submerged corals. (Adapted from Fairbanks, 1989)

-o.o -o.i -0.2 -0.3 -0.4 -0.5 -0.6 -0.7 -0.8 -0.9 - 1.0

Figure 6.2 Barbados sea-level curve for the last 17,000 years based on U-Th dating of submerged corals. (Adapted from Fairbanks, 1989)

rose to around —60m at 10,000yrbp (Fairbanks, 1989). The Late-glacial sea-level rise was interrupted by two major meltwater pulses, at 14,000 and 11,000 yr bp (Bard et al, 1990).

Glacio-isostasy is crustal deformation resulting from the build-up and decay of great ice sheets. The crustal deformation varies with the rigidity of the crust. The depression at one place must be compensated elsewhere, and hence marginal displacement of the crust involving upward bulging (forebulge) may be one aspect of this compensation (Peltier, 1987). The distance between the margin of the ice sheet to the fore-bulge depends on the flexural parameter of the crust, or the amplitude of bending of the lithosphere, which is mainly related to lithospheric density, thickness and elasticity.

Glacio-isostatic recovery can be considered as a process that accelerates rapidly and then slows gradually. Glacio-isostatic recovery in response to deglaciation can be subdivided into three phases (Andrews, 1970):

(1) Restrained rebound occurs beneath a thinning ice sheet. This period is not recorded by direct sea-level data because the area is covered by ice.

(2) Postglacial uplift is the rebound phase after deglaciation. Relative sea-level variations can be recorded by means of geomorpholo-gical and sedimentological evidence.

(3) Residual uplift is the rebound that takes place several thousand years after the region is deglaciated. Some regions that were occupied by ice sheets during the last ice age are still rising at present because of the long response time of lithospheric recovery. In many areas isostatic uplift is considered not to be complete. Highland Britain still rises by almost 0.2cmyr_1 (Shennan, 1989), the Gulf of Bothnia by 0.8-0.9 cm yr"1 (Broadbent, 1979), while the eastern Hudson Bay area in Canada experiences an uplift of about 1.1cm yr-1 (Hillaire-Marcel, 1980).

Geophysicists are interested in constraining their models of lithospheric deformation using empirical data on crustal movements, while geomorphologists want to reconstruct dispersal centres of former ice sheets. The complete response of sea-level in an area can be divided into three segments of the sea-level curve: (a) the ice build-up or loading phase which occurs during ice sheet expansion; (b) the equilibrium phase with stable high sea-level when the ice sheet is at maximum thickness; and (c) the deglacial unloading phase characterized by relative sea-level fall.

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