Coral reef as an archive of sea level variation

Massive tropical corals like Porites are considered among others a paleo-sea level indicator due to their dense massive structure that is resistant to strong waves, mechanical breakage as well as erosion compared to branching species (Corrège, 2006). They grow between a few millimeters to 3 cm per year, forming annual growth bands which can be used for chronology (Grottoli 2001).

In general, they grow from very close to sea level to ~25 m below sea level (Carpenter et al., 2008; Pratchett et al., 2013). Because of a large range of growth the normal massive Porites alone do not necessarily provide precise constraints on the position of local sea level. For this reason in Fossil Porites micro atolls are considered as a useful marker of the sea level (Woodroffe, C. D, 2005). Porites micro atoll put a distinct constraint on the position of a past sea level because they grow only a few centimeters below the sea level, their vertical growth is limited by the longer period of exposure at low tide (Flora and Ely, 2003; Smithers and Woodroffe, 2001; Woodroffe, C. D, 2005). In case of low availability of fossil micro atoll, a combination of normal fossil Porites and micro atoll in the reconstruction can help to reduce the depth uncertainty in the sea level reconstruction. Under normal conditions growth direction of the corals in the tropical oceans follows the water level (Davies and Marshal., 1980). When the sea level increases, the dominant direction is vertical. The still-stand position or falling of the sea level imposes restrictions on the vertical growth of a reef resulting into reef exposure (Davies and Marshal., 1980; Eisenhauer et al., 1999). Most of the corals tend to with stand short period of exposure, while longer periods are generally fatal and cause mortality.

Geological evidences from South Pacific and Indian Ocean islands (far-field areas) predicted that the sea level was ~1-3m above the present at ~6.5 ka (Mid Holocene) during the sea level high stand (Woodroffe and Horton, 2005). This was caused by the immense volumes of water from

the melting of ice sheets during the last interglacial resulting into raising sea level in regions which are far from the major glaciation centers (far-field locations). For the Late Holocene these areas experienced the relative sea level (rsl) drop to the present level (Banerjee, 2000;

Deschamps et al., 2012; Eisenhauer et al., 1993; Woodroffe and Horton, 2005).

The concept of rising and falling of the sea level in the far-field areas is governed by the theory concerning the former ice sheets dynamics and Glacial Isostatic adjustments (Lambeck et al., 2002; Milne et al., 2009; Mitrovica and Milne, 2002) that induce water migrations through the process. During the Last Glacial Maximum (~20,000yrs ago) the weight of continental ice sheets was exerting the downward pressure causing deformation of the crust that forced sub-lithospheric flow of mantle away from the centers of load that has caused the low geoid in these areas. This in turn caused formation of the fore-bulge (flexural bulge) around the ice load (Fig.

3a). The ice sheets exerted a gravitational pull of the ocean water causing a slight sea level increase in the around these areas (Mitrovica and Milne, 2002). During deglaciation the lack of gravitational pull which was previously exerted by ice mass on the ocean water resulted in a sea level drop nearby the formerly glaciated area and in a sea level rise higher than the eustatic value at the opposite end (Mitrovica and Milne, 2002; Woodward, 1888). Therefore, the ocean averaged sea level change exactly corresponds to the eustatic change (Suess and Waagen, 1888), but the local sea level change may be significantly different, or even opposite in sign depending on gravity and the distance from the former ice masses. Because the load of ice decreases as the ice melt, the formerly glaciated areas undergo isostatic rebound (rise of the land mass) which is caused by sub-lithospheric flow back of the mantle towards the unloaded former glaciated regions inducing the collapsing of the fore-bulges in order reach new isostatic equilibrium (Fig.

3b).

Figure 3a: Illustrates the downward pressure exerted into the Earth lithosphere forcing sub-lithospheric flow of mantle away from the centers of load to form the fore-bulges on the periphery of the ice cover. Figure 3b: Isostatic rebound of the Earth crust as a result of melting of the continental ice sheets causing for-bulges to sink as a result of mantle flowing away from the bulges.(Source: http://xenon.colorado.edu/spotlight/index.php?product=spotlight&

station=CHUR).

This resulted into the so called “Ocean siphoning effect” migration of water from the far-field equatorial oceans towards the subsiding peripheral fore bulges (near-field areas) to fill the space vacated by these subsiding bulges (Fig. 4) causing sea level regression in the far-field areas (Mitrovica and Milne, 2002). Note, an increase in gravity of the rebound areas also add an effect (gravitational pull) on the migration of water towards the near-field areas.

Figure 4: Isostatic rebound (crustal rebound) of the Earth lithosphere to maintain isostatic equilibrium (Source:

Mitrovica and Milne, 2002). This causes the collapse of the fore-bulge that induces migration of water to the near field areas to fill the space vacated by collapsing fore-bulges. As a consequence, the sea level fall in the far-field equatorial regions that resulted into exposure of the coral platforms along the coastal areas of the far-field regions.

These are the characteristic platforms that are found in the Pacific and Indian Ocean.

As a consequence of sea level regression, the coral reefs developed extended emerged fossil reef platforms (Fig. 5) which are currently 1-3 m above the modern sea level (Eisenhauer et al., 1999;

Eisenhauer et al., 1993; Grossman et al., 1998; Montaggioni and Pirazzoli, 1984; Pirazzoli et al., 1988; Rashid et al., 2014; Woodroffe and Horton, 2005). These platforms are characteristic of Chagos Islands, Cocos (Keeling), Maldives, Laccadives in Indian Ocean and most of islands in the Southern Pacific Ocean (Eisenhauer et al., 1999; Montaggioni and Pirazzoli, 1984; Pirazzoli et al., 1988). Barbados is described by Pirazzoli (1996) as being located in an intermediate field site which corresponds to the peripheral bulge around a former ice margin. This tends to subside in late and post-glacial times, to compensate the uplift in nearby formerly glaciated areas.

Figure 5a: Exposed fossil reef platform in Moorea Island that was formed as a consequence of migration of water away from these areas as a result of ocean siphoning effect. Figure 5b: Fossil Porites in growth position (in situ).

Figure 5c: Fossil Porites micro atoll in growth position. These images were taken during our field study.

To reconstruct sea level variations for these platforms which are currently exposed above the present sea level the elevation of an exposed fossil coral relative to present mean sea level and its age estimate (either by ¹⁴C or U/Th dating) is used. For islands of volcanic origin (e.g. Society Islands), the island’s specific subsidence rate needs to be applied in order to constrain the actual height of a sample above the sea level (Rashid et al., 2014).

1.5.1. The principle applied for coral U/Th age dating

Corals can be accurately dated using U-series disequilibrium method which is based on the radioactive decay of radionuclides within the naturally occurring decay chains. Dating of fossil corals is done by measuring concentrations of radioisotopes incorporated within the skeletal matrix at the time of accretion (skeletogenesis) and the concentration of the daughter nuclei that have grown with time since accretion (Bourdon et al., 2003). For U/Th dating of corals we consider ²³⁸U as the original parent taken from the sea water and ingrowth of the ²³⁰Th as the daughter nuclei (decayed from the parent nuclide) within the coral skeleton. In this case it is assumed that during skeletogenesis ²³⁰Th is not incorporated into the skeleton due to its insolubility in the sea water, and the system remains closed after deposition (U and Th are not remobilized). In nature uranium mainly exists in two oxidation states (U⁴⁺ and U⁶⁺). The U⁶⁺ is soluble and appears to be dominant in the form as uranyl ionand in various uranyl carbonate forms (Edwards et al., 2003). The U⁴⁺ is insoluble and thus far less mobile. In contrast to uranium Th is insoluble in natural waters. Because it is particle reactive, once transported into the sea water it is adsorbed to particles or minerals and settles mostly at the bottom waters in sediments.

During skeletogenesis (formation of skeletal CaCO3) corals incorporate uranium from the seawater in the CaCO3 crystal matrix and almost free of ²³⁰Th. Fractionation between the different U isotopes (²³⁴U and ²³⁸U) does not occur during coral growth therefore distribution of

234U/²³⁸U (δ²³⁴U) in the ocean is homogenous at an average of 149.6±3 ‰ (Delanghe et al., 2002).

As time passes ²³⁴U in the sample, with a half-life of 245,250±490 years, decays to ²³⁰Th (Cheng et al., 2000). The decay of excess ²³⁴U and the growth of ²³⁰Th can be expressed as a function of age (Kaufman and Broeker, 1965) under the following criteria:

1. The sample should have primary aragonitic skeletal structure.

2. The carbonate should remain a closed system with respect to uranium and its decay products.

3. Uranium ²³⁸U concentration should reflect the modern analogues from the same region.

4. The initial δ²³⁴U values should lie within the range of modern corals and sea water between 141‰-157‰.

5. The ²³²Th concentration should be <2 ppb

6. Abundance of calcite from X-ray diffraction must be below detection limits (<1% calcite) Unfortunately the assumption that initial incorporation of uranium from the sea water is free from ²³⁰Th during skeletogenesis is not fulfilled (Edwards et al., 2003) therefore, correction of initial non radiogenic ²³⁰Th (detrital) is necessary. Since it is not possible to distinguish between the radiogenic and detrital ²³⁰Th, therefore, ²³²Th is used as an indicator of detrital contamination because it has the same chemical properties as ²³⁰Th. Note that, the formula applied for age calculations and the detrital corrections are shown in the methodology (section 2.3. 2). Only U-series ages that fulfil all these requirements are considered to be reliable and can be used for paleo-climate reconstructions.

1.5.2. Corals as archives for sea surface temperature (SST) reconstruction

Massive scleractinian corals offer a reliable tool for estimating environmental and climatic parameters to the time period beyond instrumental records. This is because they incorporate chemical signatures (elements) from the ambient sea water during calcification process (Felis and Pätzold, 2004). The behaviors of many of these elements (isotopes to trace metals) are believed to be more or less controlled by external environmental conditions of ambient sea water, thus providing us with an ability to use them as a reliable tracer (Corrège, 2006). Several proxies such as Mg/Ca, Mg/Li, B/Ca have been proposed for SST estimates but their robustness have not yet completely proved (Gagan et al., 2000; Oomori et al., 1983; Min et al., 1995;

Mitsuguchi et al., 1996; Quinn and Sampson, 2002; Shen and Dunbar, 1995; Watanabe et al., 2001). Currently δ¹⁸O isotopes, Sr/Ca and U/Ca are commonly used for paleo-temperature studies where Sr/Ca is considered to be the most promising proxy in reconstructing the past climate history. However, U/Ca proxy is still not very well established (Min et al., 1995). Proxy evidence of paleo-climate reconstruction helps a better understanding the climate system that allows analysis of the current climate into a broader context which improves projections of future climate. The use of multi-proxy approach improves the strength of paleo-climate reconstruction because it combines the information from different proxies that uses advantage of the strengths of one proxy and minimizes the limitations of some other proxies.