Sea level changes and application of Oxygen isotopes

Sea level changes In the geological time scale, global sea level changes are tied primarily to long-term (10 7 –10 8 -year scale) 10 7 –10 8 years sea level changes are slow and relative (< 10 m Myr) Tectonism and short-term (10 3 –10 6 -year scale) A global sea level rise of 120 m followed the Last Glacial Maximum, with rates that exceeded 10 times the modern rate of rise (> 40 mm / yr versus ~ 3 mm/ yr). The last glacial period was the most recent glacial period within the current ice age occurring during the last years of the Pleistocene, from approximately 20,000 to 26,000 years ago. Reasons for sea level changes are 1, Temperature variations in the climate 2, Tectonism 3, Sedimentation 4, Ocean floor production. 5, Epirogeny

Sea level changes in the geological records Figure : Sea level estimates derived from New Jersey (Miller et al., 2005a; Kominz et al., 2008) that provide a testable record of global sea level for the past 100 million years. The fgure also shows the oxygen isotopic synthesis of cramer et al. (2009). NhiS = Northern hemisphere ice sheets. Ma = millions of years ago

Annual average mean sea level Tide gauge measurements and satellite altimetry suggest that sea level has risen worldwide approximately inches (12-22 cm) during the last century (IPCC, 2007). Along the coasts of the north Indian Ocean, tide-gauge data show an average rise of 12.9 cm per century (Unnikrishnan. et al)

The global sea levels have generally remained stable from about 7000 yrs before to the present; the rise in sea level has been marginal (Fleming et al., 1998). Horton et al. (2005) for the Malay–Thai Peninsula revealed an upward trend of the sea level from 22 m at 9670–9250 cal yr B.P. to a mid-Holocene high- stand of 5 m at 4850–4450 cal yr B.P. Spatial variability in the sea level history and in the absence of Sea level curve for the A&N region we are unable to account for any such changes (Rajendran et al., 2008).

Rise in mean sea level: Global and in Indian Ocean By: M. Ravichandran (personal communication)

Oxygen Isotopes Oxygen is one of the most significant keys to deciphering past climates. Oxygen comes in heavy and light varieties, or isotopes, which are useful for paleoclimate research. Like all elements, oxygen is made up of a nucleus of protons and neutrons, surrounded by a cloud of electrons. All oxygen atoms have 8 protons, but the nucleus might contain 8, 9, or 10 neutrons. Light oxygen-16, with 8 protons and 8 neutrons, is the most common isotope found in nature. Heavy oxygen-18, with 8 protons and 10 neutrons. δ 18 O is a measure of the ratio of stable isotopes 18 O: 16 O. 18 O: 16 O data from foraminifera and ice core are used as a proxy for temperature. The definition is, in "per mil" (, parts per thousand) Standard known isotopic composition : VSMOW Source :

Role of Oxygen Isotopes : Ice sheets preferentially store the lighter isotope 16 O, and their growth results in higher δ 18 O values in the ocean. Growth of the latitude Ice Sheet lowered sea level 120 m by 20,000 years ago and increased global seawater δ 18 O values by ~ 1.2 (Fairbanks, 1989). Conversely, melting the present-day global inventory of ice (~ 25 x 106 km 3 and 64 m of sea level equivalent; Lythe et al., 2001) would raise global δ 18 O values by ~ 0.9 (Shackleton and Kennett, 1975).

Evaporation and condensation are the two processes that most influence the ratio of heavy oxygen to light oxygen in the oceans. Water molecules containing light oxygen evaporate slightly more readily than water molecules containing a heavy oxygen atom. At the same time, water vapor molecules containing the heavy variety of oxygen condense more readily. As air cools by rising into the atmosphere or moving toward the poles, moisture begins to condense and fall as precipitation. At first, the rain contains a higher ratio of water made of heavy oxygen, since those molecules condense more easily than water vapor containing light oxygen. Source :

The remaining moisture in the air becomes depleted of heavy oxygen as the air continues to move poleward into colder regions. As the moisture reaches the upper latitudes, the falling rain or snow is made up of more and more water molecules containing light oxygen. Water vapor gradually loses 18 O as it travels from the equator to the poles. Because water molecules with heavy 18 O isotopes in them condense more easily than normal water molecules, air becomes progressively depleted in 18 O as it travels to high latitudes and becomes colder and drier. In turn, the snow that forms most glacial ice is also depleted in 18 O. As glacial ice melts, it returns 16 O-rich fresh water to the ocean. Therefore, oxygen isotopes preserved in ocean sediments provide evidence for past ice ages and records of salinity. (Illustration by Robert Simmon, NASA GSFC)

Fossilized oxygen isotopes The shells of tiny plants and animals and corals are typically made of calcium carbonate (CaCO3), which is the same as limestone, or chalk, or silicon dioxide (SiO2), similar to the compound common in quartz sand. As the shells form, they tend to incorporate more heavy oxygen than light oxygen, regardless of the oxygen ratio in the water. The biological and chemical processes that cause the shells to incorporate greater proportions of heavy oxygen become even more pronounced as the temperature drops, so that shells formed in cold waters have an even larger proportion of heavy oxygen than shells formed in warmer waters, where the difference is less notable. Therefore, low δ18O values in foraminifera indicate warm periods and high values cold periods. Source :

R CORRECTION The most common method for determining ages of marine sediments is radiocarbon dating of fossil calcareous tests of surface dwelling foraminifers and is frequently employed in high- resolution paleoclimate studies of the Late Quaternary period. The 14C ages of marine fossils are on the average 400 years older than contemporary terrestrial wood, since the reservoir (seawater) from which these foraminifers derive carbon has lower 14C/12C ratios compared to the atmosphere, due to mixing with deeper 14C depleted water. Considerable spatial variability is seen in the apparent 14C ages of marine calcareous shells due to variations in the regional ocean circulation patterns. For a given region (s)at any given time (t), the difference between the regional marine 14C age Rs(t) and the global model marine 14C age Rg(t), is given by R(s) (Stuiver and Braziunas 1993). The ΔR for a specific location (s) can be calculated from known age shells, collected prior to atmospheric bomb testing using the formulae: Rs(t) – Rg(t) = ΔR(s) The ΔR standard error is calculated by the formulae: ΔRσ = (σRs(t) 2 + σRg(t) 2 ) By: Dutta et al.