Encyclopedia of Modern Coral Reefs: Structure, Form and Process

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ACANTHASTER PLANCI Ian Miller Australian Institute of Marine Science, Townsville, QLD, Australia Acanthaster planci (Class Asteroidea; Order Spinulosida; “crown-of-thorns sea star or starfish”) is a large (up to 70 cm), mobile, multi-armed (7–23) sea star covered in sharp, toxic spines. It feeds almost exclusively on hard corals and is found on coral reefs throughout the IndoPacific. No other reef sea stars remotely resemble its appearance, nor possess comparable life-history traits as a predator on corals. Crown-of-thorns are prone to population outbreaks, with aggregations of thousands or more adults per hectare not uncommon (Figure 1). Such populations often advance in fronts through coral habitat, leaving formerly luxuriant coral areas dead in their wake. The sea star has a number of life history traits that predisposes these destructive population outbreaks: absence of any equivalent coral predator (little competition for food); a large stomach (that is pushed out through the mouth to digest coral tissue externally); a high fecundity (a mature female can produce some 50 million eggs); planktonic larvae (that can feed in the water column and disperse over long distances); rapid growth (
10 cm.y1, that is faster than any other coral reef sea star); large size and toxic spiny armature (that provide protection from potential predators); multi-armed morphology and tube feet (allowing them to climb and feed in nearly any position). Repeated population outbreaks have decimated hard corals throughout the Indo-Pacific over the last 50 years.

Outbreaks were first observed in the 1960s. The geographical extent (two oceans) and impact (an ecosystem changed from one dominated by hard corals to one dominated by algae), shocked scientists. A key management issue was whether human activity had somehow precipitated the population outbreaks. Two main hypotheses have been developed that implicate anthropogenic factors. The first is “the predator removal hypothesis” (Endean, 1969), which holds that overfishing (in particular sweetlips (Family Lethrinidae), some wrasses (Family Labridae) and some triggerfish (Family Balistidae)) and collecting of predators of the sea star (notably a large Gastropod mollusc the giant triton Charonia tritonis), allow crown-of-thorns to build up in numbers on a reef. On reaching a critical abundance, their reproduction and larval dispersal leads to successful recruitment of larvae on reefs downstream in prevailing currents. A cascade of outbreaks across tracts of neighboring reefs ensues. The second hypothesis (possibly synergistic with the first) is “the nutrient enrichment hypothesis” (Birkeland, 1982, Lucas, 1982). In this scenario, river runoff from human-modified catchments enhances nutrients in coastal waters, resulting in an increase in phytoplankton upon which the sea star larvae feed. Because crown-of-thorns produce such a vast quantity of eggs even a small increase in survivorship leads to larger settlement of larvae onto a reef, which in turn leads to a primary outbreak. Today, despite repeated outbreaks and years of research the exact events leading to the initiation of an outbreak remain enigmatic. This is because the lifehistory of crown-of-thorns makes it difficult to disentangle the natural processes leading to an outbreak from those forced by human activities. As a result crown-ofthorns remains a major management problem for coral

David Hopley (ed.), Encyclopedia of Modern Coral Reefs, DOI 10.1007/978-90-481-2639-2, # Springer Science+Business Media B.V. 2011

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ACCOMMODATION SPACE

ACCOMMODATION SPACE Tom Spencer University of Cambridge, Cambridge, UK

Definition The space available, in both a vertical and a lateral sense, within which corals can grow, increase framework and sediments accumulate.

Acanthaster Planci, Figure 1 A crown-of-thorns feeding aggregation. Such outbreaks of the sea star are a major recurrent cause of coral mortality on coral reefs throughout the Indo-Pacific (photo: AIMS LTMP).

reefs. Where adults have been collected as a control measure, coral has been saved from predation only over relatively small areas (hectares). In the past, coral cover has generally recovered within 10–15 years of an outbreak. However coral resilience in the face of future outbreaks is uncertain (Done, 1987). This is because the size and frequency of other impacts that can effect coral reefs (such as cyclones, coral bleaching, and ocean acidification) are predicted to increase in coming years due to greenhouse gas emissions. Without full recovery, repeated outbreaks will eventually lead to the degradation of the coral reef community.

Bibliography Birkeland, C., 1982. Terrestrial runoff as a cause of outbreaks of Acanthaster planci (Echinodermata: Asteroidea). Marine Biology, 69, 175–185. Birkeland, C., and Lucas, S. L., 1990. Acanthaster planci: major management problem of coral reefs. Boca Raton, Florida: CRC Press. Done, T. J., 1987. Simulation of the effects of Acanthaster planci on the population structure of massive corals in the genus Porites: evidence of population resilience? Coral Reefs, 6, 75–90. Endean, R., 1969. Report on Investigations Made into Aspects of the Current Acanthaster planci (Crown-of-thorns) Infestations of Certain Reefs of the Great Barrier Reef. Fisheries Branch, Queensland Dept. of Primary Industries, Brisbane. p. 35. Lucas, J. S., 1982. Quantitative studies of feeding and nutrition during larval development of the coral reef asteroid Acanthaster planci (L.). Journal of Experimental Marine Biology and Ecology, 65, 173–194.

Cross-references Coral Reef, Definition Corals: Environmental Controls on Growth

Accommodation space For corals, accommodation space is constrained vertically by the water-air interface and its volume broadly determined by reef widths and slope angles. For sedimentary accumulations on reef platforms, the lower boundary is governed by reef margin position, reef flat elevation and lagoon depth and the upper boundary set by the height of wave run-up during storm events. The rate at which accommodation space can be filled depends upon rates of vertical coral growth, vertical framework accretion and sediment supply, transport and accumulation; these are all controlled by reef productivity and sediment generation processes which may themselves be constrained by the environment processes (e.g., wave exposure locally prevents coral growth from filling accommodation spaces on Hawaii; Grigg, 1998), be periodically interrupted by storms (see Tropical Cyclone/Hurricane) and modulated by sea level change, which ultimately determines the upper margin of the accommodation space (Figure 1). Over the long-term, a subsiding reef basement results in an increase in accommodation space. During glacial periods, emergent reefs were subject to subaerial solution, thus increasing the vertical accommodation space available for reef re-growth on renewed inundation during interglacial periods. There has been debate over the subaerial erosion rates involved, and thus the additional accommodation space generated, ranging from minimal downwearing (e.g., Quinn and Matthews, 1990) to 6–63 cm of surface lowering per 1,000 years (e.g., Gray et al., 1992). During the stable sea-level of the late Holocene, there has been “turn-off ” of both vertical and horizontal growth of some reefs due to the progressive thinning of accommodation space as reefs approached present sea level (which itself may have fallen slightly in some Indo-Pacific locations, thus further reducing accommodation space) and the difficulty of lateral expansion, and maintenance of reef front volume and integrity, over relatively unstable reef front talus deposits in increasing water depths. For example, Smithers et al. (2006) attributed the shut-down of fringing and nearshore reef progradation on the Great Barrier Reef between 5.5–4.8 ka BP and 3.0–2.5 ka BP to the contraction of accommodation space caused by the reefs’ own growth and the complete occupation of favorable reef foundations. In the short term, local accommodation space is an outcome of local reef erosion and the re-configuration of sedimentary accumulations resulting from hurricane and

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Accommodation Space, Figure 1 Different models of fringing reef development show different modes of accommodation space filling. (a): accommodation space is filled by corals showing catch-up or keep-up behavior. (b): reef accretion is lateral, having established at a level with little or no vertical accommodation space. Isochrons are in thousands of radiocarbon years BP (From Kennedy and Woodroffe, 2002).

cyclone impacts (see Tropical Cyclone/Hurricane). It has been argued that rates of sea level rise of
0.5 m by AD 2100 might create new accommodation space and switch reef vertical accretion back on, with carbonate production for the entire Great Barrier Reef rising from the current estimated 50 Mt a1 to 70 Mt a1 (Kinsey and Hopley, 1991).

Bibliography Cowell, P. J., and Thom, B. G., 1994. Morphodynamics of coastal evolution. In Carter, R. W. G., and Woodroffe, C. D., (eds.), Coastal Evolution: late Quaternary shoreline morphodynamics. Cambridge: Cambridge University Press, pp. 33–86. Cowell, P. J., and Kench, P. S., 2002. The morphological response of atoll islands to sea-level rise. Part 1: modifications to the shoreface translation model. Journal of Coastal Research, ICS 2000, 633–644. Gray, S. C., Hein, J. R., Hausmann, R., and Radtke, U., 1992. Geochronology and subsurface stratigraphy of Pukapuka and Rakahanga atolls, Cook Islands: Late Quaternary reef growth and sea level history. Palaeogeography, Palaeoclimatology, Palaeoecology, 91, 377–394. Grigg, R. W., 1998. Holocene coral reef accretion in Hawaii: a function of wave exposure and sea level history. Coral Reefs, 17, 263–272. Kennedy, D. M., and Woodroffe, C. D., 2002. Fringing reef growth and morphology: a review. Earth Science Reviews, 57, 255–277. Kinsey, D. W., and Hopley, D., 1991. The significance of coral reefs as global carbon sinks – response to greenhouse. Palaeogeography, Palaeoclimatology, Palaeoecology, 89, 363–377. Quinn, T. M., and Matthews, R. K., 1990. Post-Miocene diagenetic and eustatic history of Enewetak Atoll: Model and data comparison. Geology, 18, 942–945. Smithers, S. G., Hopley, D., and Parnell, K. E., 2006. Fringing and nearshore coral reefs of the Great Barrier Reef: episodic Holocene development and future prospects. Journal of Coastal Research, 22, 175–187.

ACROPORA Carden C. Wallace Museum of Tropical Queensland, Townsville, QLD, Australia

Synonyms Arborescent corals; Axial branching corals; Midori ishi (Japan); Staghorn corals; Table corals Definition Acropora (Oken, 1815) is the type genus of the hard coral family Acroporidae (class Anthozoa, order Scleractinia

of the phylum Cnidaria). Currently, around 120–140 living species are recognized in this genus, but new species are still being discovered in both living and fossil coral assemblages. The Latin name derives from the growth mode, where branches are formed by a central or axial polyp, which buds off numbers of a second kind, the radial polyps, from around its tip as it extends. New branches are formed by the development of new axial polyps along the branch. This mode of growth, which is similar to the axial mode in flowing plants, allows many variations on a branching theme (Figure 1). It is thought to have been a key character in the evolution of a diverse array of species in Acropora, although other processes are also proposed, such as hybridization and reticulate evolution facilitated by the mass spawning of related species.

Introduction Six coral families (Acroporidae, Faviidae, Mussidae, Poritidae, Fungiidae, and Pocilloporidae) dominate modern world reef composition, in terms of diversity, abundance, geographic range, and contribution to accretion of reef carbonates. Of these, Acroporidae is arguably the most successful, as the two most species-rich genera, Acropora and Montipora, allow it to dominate the species diversity and coral cover of most Indo-Pacific reef locations. Acropora the “staghorn” corals have played a role in the biodiversity, ecology, and structure of coral reefs for almost 60 million years (Schuster, 2003; Wallace and Rosen, 2006). Their mode of skeletal construction, where polyps are supported within an open “synapticular” framework (Figure 2), allow for rapid growth with efficient use of calcium carbonate (Gladfelter, 2008) and provide habitat complexity for other reef biota (Munday, 2002). Strong representation in mass coral spawning and recruitment events, and rapid recolonization after destructive natural events are the characteristics of Acropora (e.g., Babcock et al., 1986; Connell et al., 2004): however, this genus may experience severe localized or widespread loss of diversity from major perturbations such as coral bleaching due to elevated seawater temperature, cold-water events, tsunamis, cyclone damage, and predator population outbreaks, particularly of Acanthaster planci, the crown-ofthorns sea star (Wilkinson, 1998–2008; Berklemans et al., 2004; Marshall and Baird, 2006). Chronic anthropogenic impacts such as nutrient and sediment run-off,

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ACROPORA

Acropora, Figure 1 Examples of colony shapes in Acropora: (a) Arborescent (A. grandis), (b) Arborescent table (A. valenciennesi), (c) Corymbose (A. anthocercis), (d) Digitate (A. gemmifera), (e) Hispidose (A. echinata), and (f) Table (A. clathrata). (Photos: P. Muir.)

overfishing, and coral mining for limestone also have an impact on Acropora (Fabricius, 2005; Fabricius and Wolanski, 2000; Brown, 1997).

Nomenclatural issues Until the late nineteenth century Acropora was known mostly as Madrepora, a broadly applied name, which is

now restricted to a genus of non-zooxanthellate deepwater corals. The name Acropora was stabilized by a decision of the International Code of Nomenclature in the midtwentieth century (Boschma, 1961; China, 1983), which also ruled on a type species, A. muricata (Linneaus, 1758). A dilemma concerning the nature and provenance of this species, described by Linnaeus as being from the

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Acropora, Figure 2 (a) Synapticular formation of Acropora skeleton, shown in scanning electron micrograph of A. abrotanoides. Also seen is the axial corallite (a) and radial corallites (r) (Scale: 500 mm). (b) High power SEM of synapticular formation in Acropora nasuta. Two synapticulae approaching each other will form a node (n), from which another synapticula will develop at right angles (Scale: 100 mm). (Photos: P. Muir and C. Wallace.)

“Asian Sea,” was resolved by designation of a neotype from central Indonesia (Wallace, 1999). The size of the genus was reduced slightly by elevation of a subgenus Isopora to separate genus status because it differed from other Acropora in skeletal and reproductive morphology as well as reproductive mode (Wallace et al., 2007). The remaining species are organized into 20 species groups based on skeletal features (Veron and Wallace, 1984; Wallace, 1999). Genetic studies are revealing numerous dilemmas about species boundaries (Van Oppen et al., 2001, 2002) and at least one named species of Acropora is now known to be an F1 hybrid (Van Oppen et al., 2000; Vollmer and Palumbi, 2002).

The skeleton and polyps All scleractinian corals have skeletons of the crystal aragonite form of calcium carbonate, but formation of a skeleton by the polyps follows different patterns among families, giving structural features by which corals can be identified in both living and fossil form (Wells, 1956; Roniewicz, 1996). In Acroporidae, most components of the skeleton are formed by the development of simple rods or “synapticulae,” which allow for a strong but light and open growth (Figure 2; Nothdurft and Webb, 2007; Rosen, 1986). The potential of this mode for rapid growth in three dimensions is exemplified by the axial growth of Acropora. This provides a light scaffolding to support the living colony and allows organization of the tissues into a gastrovascular system in which flagellated gastrodermal cells promote laminar flow up, down and around the branch to transport water and nutrients. The axial polyp extends through most of the branch and is thus much longer than the radials. A notable aspect of the growth mode of Acropora is that some species have symmetrical growth around a central growing point and maintain a limited “determinate” pattern of growth, while others exhibit unlimited and asymmetrical grow, filling

in available space wherever it comes up. These contrasting modes allow Acropora colonies to efficiently fill available space on the reef (see Figure 3a). The polyp cavities are extended by the coenenchyme, a complex network of tubules containing extensions of the gastric cavity. Much of the skeletal variation used for taxonomic delineation of species comes from the shape of the radial corallites and the microstructure of the skeleton (Wallace, 1999). Another form of skeleton, the epitheca, formed by calcite form of calcium carbonate, is present in very small quantities below the living tissues of the branch and acts as a sealant preventing infection and protecting the live polyps and coenenchyme from fluid loss (Barnes, 1972). The polyps of Acropora have a simple tubular structure and 12 tentacles, one of which extends greatly as a “catch tentacle,” particularly when the polyps are feeding at night (Wallace, 1999). Below the tentacles are the mesenteries, which carry the gonads when they develop and have a muscular internal filament, which can extend outside the polyp for defense, clearing space, and possibly feeding (Roff et al., 2009).

Habitats and ecology Acropora is often interpreted as being a reef-front genus, favoring sites with good circulation, high oxygen content due to the strong movement of water, and access to food from oceanic waters. While a diverse “Acropora zone” from the reef top to about 12 m depth, is indeed a characteristic of most oceanic Indo-Pacific reefs, this genus also occurs significantly in specialized habitats such as sandy lagoon floors, deep reef slopes and deepwater Halimeda banks, and in relatively turbid fringing reef locations. The persistence of an abundance of colonies and diversity of species through time on any reef habitat relies on a complex interaction of water quality and physical and biological parameters: the activities of other reef organisms also impact on survivorship at each life stage

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Acropora, Figure 3 (a) Numerous forms of Acropora maximize usage of three-dimensional space in shallow reef shoals: ten species occupy this frame. (b) Steep walls support mainly small plates which maximize purchase and exposure of the polyps to light. (Photos: P. Muir.)

and compete with corals for resources (Done, 1999; Done et al., 2007). Human-influenced deterioration of conditions suitable for survival and/or settlement of Acropora, or favoring survival and population increase of other benthic organisms such as the coral genus Porites, soft corals, and algae, are thought to be involved in the gradual deterioration of Acropora communities in many parts of the world and to threaten the survival of the rarer or more narrowly distributed species of this genus (McClanahan et al., 2008). Many rare species may be threatened by diminishing habitats (Carpenter et al. 2008), although it has been suggested that some rare species with small global population sizes are actually unidirectional hybrids, and that this contributes to increased genetic variability and adaptive potential, making them less vulnerable to extinction (Richards et al., 2008). In the Caribbean, the two major Acropora species A. cervicornis and A. palmata have undergone massive population loss over recent decades in many locations as a result of exposure to repetitive hurricanes, coral diseases, changes in water quality, and outbreaks of other reef organisms (Bythell et al., 1993; Williams et al., 1999) and such rapid declines appear to be unprecedented in the past 4,000 years (Aronson and Precht, 1997).

Sexual and asexual reproduction Species of Acropora reproduce sexually by developing gametes (eggs and sperm) along the mesenteries (radial dividing structures) within the polyps. In Acropora polyps have both sexes (hermaphrodite), with certain mesenteries bearing eggs and others sperm (Wallace, 1985) (Figure 4). Fertilization takes place externally, after the gametes are released into the water column, often during mass spawning events involving many species. The fertilized eggs develop into ciliated larvae known as planulae,

Acropora, Figure 4 Egg sperm bundles of Acropora tenuis leaving polyps and ascending into the water column, during a mass spawning event on the Great Barrier Reef, Australia. (Photo: Z. Florian.)

which spend some days in the water before being ready to settle on reef surface to begin a new colony. Planulae remain viable for days to weeks and may be transported long distances and settle away from the home reef. This contrasts with sexual reproduction in the sister genus Isopora, where sperm is released but the eggs stay within the polyp, where they are fertilized by sperm from other colonies and develop into larvae which are released ready to settle on the reef. The two contrasting modes of reproduction may have different consequences for the genera after loss of the adult corals in a population: for Acropora, the possibility of recruitment of larvae from healthy reefs is greater. This has been seen after mass bleaching of Acropora, for example, in the Maldives (Wallace and Zahir, 2007) and Socotra (western Yemen), where

ACROPORA

Acropora recruits were visible several years before the appearance of Isopora, following the 1998 bleaching event (L. DeVantier, personal communications).

Genetics and phylogeny Acropora has a large and complex genome and this has been studied in detail for certain species and species groups as well as in the context of genus-level phylogenies. Molecular (genetic) studies show corals to have two main evolutionary lines, known currently as the “Robust” and “Complex” clades. Acropora and other members of the family Acroporidae fall within the Complex clade (Romano and Cairns, 2000; le Goff-Vitry et al., 2004; Chen et al., 2002). Evolution of the mitochondrial genome of all Anthozoa is typically slower than that of other animals, making it difficult, for example, to use cytochrome b to study population genetics in these animals, but it is faster in Acropora than in confamilial genera (Van Oppen et al., 1999). The tempo of evolutionary change is faster in the nuclear and slower in the mitochondrial genomes for Acropora (and other corals studied), making them more similar to plants than other animal groups in this respect (Hellberg, 2006; Chen et al., 2009). Genetic studies and laboratory cross-fertilization experiments on Acropora from within species groups (especially the A. aspera, A. cervicornis, and A. humilis groups) have indicated that hybridization and introgression may play a significant role in maintaining variety within populations and associations of Acropora species, perhaps contributing to resilience in the face of adverse conditions (Van Oppen et al., 2000; Wolstenholme et al., 2003). These and many other genetic findings for Acropora and other corals are contributing to a major revision of the characteristics and relationships within the order Scleractinia. Because corals have a hard skeleton and this remains after death, there is a superb fossil record and long-known paleontological information is currently being integrated with the molecular results to develop a new overview of relationships and evolution. Biogeography and evolution The greatest living diversity of Acropora is seen in Indonesia, where 91 species have been recorded (Wallace, 2001) and similar numbers are present in the Philippines and Papua New Guinea. The Indonesian diversity is greatest within the region known as “Wallacea,” that is, the region of islands between the Asian and Australian continental shelves (Wallace, 2001). The Indonesian Acropora diversity includes species with very extensive Indo-Pacific ranges, others restricted to the central IndoPacific, and yet others which have either predominantly Pacific Ocean or Indian Ocean distribution, with some overlaps in Indonesia. (Wallace, 2001; Wallace et al., 2001). In the Caribbean, only three living species occur. Through post-Cretaceous time, Acropora has been present in all the major reef-bearing parts of the world, including

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the fossil deposits of the Middle East and Europe, fossil and modern reefs of the Caribbean and the Indo-Pacific, where the majority of the living species occur. It is known in the fossil record from the Paleocene of Somalia (approx. 60 million years ago) and was possibly present before the end of the Cretaceous (Baron-Szarbo, 2006). Nine of the twenty species groups are represented in the mid-Eocene fossil record of England and France and this is taken to indicate that the diversification of this genus began well ahead of its later Indo-Pacific diversification (Wallace, 2008). Several species of Acropora also await description from the Oligocene to early Miocene of Europe and Miocene-Pliocene of the Caribbean. The first case of Acropora being abundant and dominant on a coral reef is seen in the Oligocene of Greece (Schuster, 2003).

Summary Acropora, the staghorn coral genus, has persisted for some 60 million years and now remains in living form in the Caribbean and the Indo-Pacific, the two main reef-bearing regions of the world. It is regarded as extremely successful because it contains the greatest number of species of any coral genus, and its species typically occur in great abundance. While it plays a major role in many aspects of reef ecology, Acropora has been shown to be extremely vulnerable to major disturbances on reefs, and thus there is concern about its persistence into the future, in the face of changes due to bleaching, disease, and other factors resulting from global climate change. The fossil record tells us that this genus has persisted and diversified through time, and genetic research is indicating that rare species may have a resilience to local extinction because of the potential for hybridization with other species. Recent decadal changes in both the Caribbean and the Pacific, however, show that Acropora can undergo local extinction in certain circumstances. It is clear that the future of this coral genus is intimately linked with the future of the world’s coral reefs. Acknowledgments Dr. P.R. Muir of Museum of Tropical Queensland for preparing figures and reviewing text. Bibliography Aronson, R. B., and Precht, W. F., 1997. Stasis, biological disturbance, and community structure of a Holocene coral reef. Paleobiology, 23, 326–346. Babcock, R. C., Bull, G. D., Harrison, P. L., Heyward, A. J., Oliver, J. K., Wallace, C. C., and Willis, B. L., 1986. Synchronous spawnings of 105 scleractinian coral species on the Great Barrier Reef. Marine Biology, 90, 379–394. Barnes, D. J., 1972. The structure and formation of growth-ridges in scleractinian coral skeletons. Proceedings of the Royal Society of London B, 182, 331–350. Baron-Szarbo, R. C., 2006. Corals of the K/T- boundary: scleractinian corals of the suborders Astrocoeniina, Faviina, Rhipiogyrina and Amphiastraeina. Journal of Systematic Palaeontology, 4, 1–108.

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Berklemans, R., De’ath, G., Kininmonth, S., and Skirving, W. J., 2004. A comparison of the 1998 and 2002 coral bleaching events on the Great Barrier Reef: spatial correlation, patterns, and predictions. Coral Reefs, 23, 74–83. Boschma, H., 1961. Acropora Oken, 1815 (Anthozoa, Madreporaria): proposed validation under the plenary powers. Bulletin of Zoological Nomenclature, 20, 319–330. Brown, B. E., 1997. Coral bleaching: causes and consequences Coral Reefs 16, S129–S138. Bythell, J. C., Gladfelter, E., and Bythell, M., 1993. Chronic and catastrophic natural mortality of three common Caribbean corals. Coral Reefs, 12, 143–152. Chen, C. A., Wallace, C. C., and Wolstenholme, J., 2002. Analysis of mitochondrial 12S RNA gene supports the two-clade hypothesis of evolutionary history of scleractinian corals. Molecular Phylogenetics and Evolution, 23, 137–149. Chen, I.-P., Tang, C.-Y., Chiou C.-Y., Hsu, J.-H., Wei, N. V., Wallace, C. C., Muir, P., Wu, H., and Chen, C. A., 2009. Comparative analyses of coding and noncoding DNA regions indicate that Acropora (Anthozoa: Scleractina) possesses a similar evolutionary tempo of nuclear vs. mitochondrial genomes as in plants. Marine Biotechnology, 11, 141–152. China, W. E., 1983. Opinion 674: Acropora Oken, 1815 (Anthozoa, Madreporaria): validated under the plenary powers. Bulletin of Zoological Nomenclature, 18, 334–335. Connell, J. J., Hughes, T. P., Wallace, C. C., Tanner, J. E., Harms, K. E., and Kerr, A. M., 2004. A long-term study of competition and diversity of corals. Ecological Monographs, 74, 179–210. Done, T. J., 1999. Coral community adaptability to environmental change at the scales of regions, reefs and reef zones. American Zoologist, 39, 66–79. Done, T., Turak, E., Wakefield, M., DeVantier, L., McDonald, A., and Fisk, D., 2007. Decadal changes in turbid-water coral communities at Pandora Reef: loss of resilience or too soon to tell? Coral Reefs, 26, 789–815. Fabricius, K. E., 2005. Effects of terrestrial runoff on the ecology of corals and coral reefs: review and synthesis. Marine Pollution Bulletin, 50, 125–146. Fabricius, K., and Wolanski, E., 2000. Rapid smothering of coral reef organisms by muddy marine snow. Estuarine, Coastal and Shelf Science, 50, 115–120. Gladfelter, E. 2008. Coral skeletons: from calcium carbonate to intricate architecture. 11th International Coral Reef Symposium, Abstracts, p. 15. Hellberg, M. E., 2006. No variation and low substitution rates in coral mtDNA despite high nuclear variation. BMC Evolutionary Biology, 6, 24. le Goff-Vitry, M. C., Rogers, A. D., and Baglow, D., 2004. A deepsea slant on the molecular phylogeny of the Scleractinia. Molecular phylogenetics and evolution. Molecular Phylogenetics and Evolution, 30, 167–177. Linneaus, 1758. Systema Naturae (edition 10) 1, 1–824 Laurentii Salvii, Holmiae. Marshall, P. A., and Baird, A. H., 2006. Bleaching of corals on the Great Barrier Reef: differential susceptibilities among taxa. Coral Reefs, 19, 155–163. McClanahan, T. R., Buddemeir, R. W., Hoeegh-Guildberg, O., and Sammarco, P., 2008. Projecting the current trajectory of coral reefs. In Polunin, N. V. C., (ed.), Aquatic Ecosystems. Cambridge: Cambridge University Press, pp. 242–260. Munday, P. L., 2002. Does habitat availability determine geographicalscale abundance of coral-dwelling fishes? Coral Reefs, 21, 105–116. Nothdurft, L. D., and Webb, G. E., 2007. Microstructure of common reef-building coral genera Acropora, Pocillopora, Goniastrea

and Porites: constraints on spatial resolution in geochemical sampling. Facies, 53, 1–26. Oken, L., 1815. Steinkorallen. Lehrbuch Naturgesch, 3, 59–74. Richards, Z. T., van Oppen, M. J. H., Wallace, C. C., Willis, B. L., and Miller, D. J., 2008. Some rare Indo-Pacific coral species are probable hybrids. PLoS ONE, 3(9), e3240. doi:10.1371/ journal.pone.0003240. Roff, G., Dove, S. G., and Dunn, S. R., 2009. Mesenterial filaments make a clean sweep of substrated for coral growth. Coral Reefs, 28, 70. Romano, S. L., and Cairns, S. D., 2000. Molecular phylogenetic hypotheses for the evolution of scleractinian corals. Bulletin of Marine Science, 67, 1043–1068. Roniewicz, E., 1996. The key role of skeletal microsctructure in recognizing high-rank scleractinian taxa in the stratographic record. Palaeontological Society Papers, 1, 187–206. Rosen, B. R., 1986. Modular growth and form of corals: a matter of metamers? Philosophical Transactions of the Royal Society of London B, 313, 115–142. Schuster, F., 2003. Oligocene and Miocene examples of Acroporadominated palaeoenvironments: Mesohellenic Basin (NW Greece) and northern Gulf of Suez (Egypt). In Proceedings 9th International Coral Reef Symposium, Bali, Indonesia, Vol. 1, pp. 199–203. Van Oppen, M. J. H., Willis, B. L., and Miller, D. 1999. Atypically low rate of cytochrome b evolution in the scleractinian coral genus Acropora. Proceedings of the Royal Society of London B, 266, 179–183. Van Oppen, M. J. H., Willis, B. L., van Vugt, H., and Miller, D., 2000. Examination of species boundaries in the Acropora cervicornis group (Scleractinia, Cnidaria) using nuclear DNA sequence analyses. Molecular Ecology, 9, 1363–1373. Van Oppen, M., Mc Donald, B., Willis, B., and Miller, D., 2001. The evolutionary history of the coral genus Acropora (Scleractinia, Cnidaria) based on a mitochondrial and a nuclear marker: reticulation, incomplete lineage sorting, or morphological convergence? Molecular Biology and Evolution, 18, 1315–1329. Van Oppen, M. J. H., Willis, B. L., van Rheede, T., and Miller, D., 2002. Spawning times, reproductive compatibilities and genetic structuring in the Acropora aspera group: evidence for natural hybridization and semi-permiable boundaries in corals. Molecular Ecology, 11, 1363–1376. Veron, J. E. N., and Wallace, C. C., 1984. Scleractinia of Eastern Australia. Part V. Family Acroporidae. Townsville: Australian Institute of Marine Science. Vollmer, S. V., and Palumbi, S. R., 2002. Hybridization and the evolution of reef coral diversity. Science, 296, 2023–2025. Wallace, C. C., 1999. Staghorn Corals of the World: A Revision of the Coral Genus Acropora (Scleractinia; Astrocoeniina; Acroporidae) Worldwide, with Emphasis on Morphology, Phylogeny and Biogeography. Melbourne: CSIRO. Wallace, C. C., 2001. Wallace’s line and marine organisms: the distribution of staghorn corals (Acropora) in Indonesia. In Metcalf, I. (ed.), Faunal and Floral Migrations and Evolution in SE Asia–Australasia. Rotterdam: Balkema, pp. 168–178. Wallace, C. C., 2008. New species and records from the Eocene of England and France for the reef-building coral genus Acropora (Scleractinia; Astrocoeniina; Acroporidae). Journal of Paleontology, 82, 313–328. Wallace, C. C., and Rosen, B. R. R., 2006. Diverse staghorn corals (Acropora) in high-latitude Eocene assemblages: implications for the evolution of modern diversity patterns of reef corals. Proceedings of the Royal Society B, 273, 975–982. Wallace, C. C., and Zahir, H., 2007. The “Xarifa” expedition and the atolls of the Maldives, 50 years on. Coral Reefs, 26, 3–5. Wallace, C. C., Richards, Z., and Suharsono, 2001. Regional distribution patterns of Acropora and their use in the conservation of

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coral reefs in Indonesia. Indonesian Journal of Marine and Coastal Resources, 4, 1–19. Wallace, C. C, Chen, C. A. C., Fukami, H., and Muir, P. R., 2007. Recognition of separate genera within Acropora based on new morphological, reproductive and genetic evidence from A. togianensis, and elevation of the subgenus Isopora Studer, 1878 to genus (Scleractinia: Astrocoeniidae; Acroporidae). Coral Reefs, 26, 231–239. Wells, J. W., 1956. Scleractinia. In Moore, R. C. (ed.), Treatise on Invertebrate Paleontology, Part F (Coelenterata), Lawrence: The University of Kansas Press, pp. F328–F444. Wilkinson, C. R. (ed.), (1998, 2000, 2002, 2004, 2008) Status of the Coral Reefs of the World. Townsville: Australian Institute of Marine Science. Williams, E. H., Jr., Bartels, P. J., and Bunkley-Williams, L., 1999. Predicted disappearance of coral-reef ramparts: a direct result of major ecological disturbances. Global Change Biology, 5, 839–845. Wolstenholme, J. K., Wallace, C. C., and Chen, C., 2003. Species boundaries within the Acropora humilis species group (Cnidaria; Scleractinia): a morphological and molecular interpretation of evolution. Coral Reefs, 22, 155–166.

Cross-references Carbonate Budgets and Reef Framework Accumulation Corals: Biology, Skeletal Deposition, and Reef-Building General Evolution of Carbonate Reefs Porites Scleractinia, Evolution and Taxonomy

ADAPTATION David Obura CORDIO East Africa, Mombasa, Kenya

Definition Adaptation is the process of change in the structure or function of an organism or parts of an organism that makes it better suited to the environment in which it lives. Adaptations (or traits that are adaptive) that are heritable, i.e., coded in genes or that have consequences on the reproductive success of genes, contribute to natural selection. Acclimatization refers to adjustment to local conditions that occurs within the lifetime of an individual, in response to external environmental conditions, for example, through behavioral changes, or increased tolerance of stressful conditions. Acclimation is similar, though is applied more narrowly to artificial conditions and experimentation. Both can lead to true adaptation if and when the relevant traits are passed on to the next generations. Introduction Adaptation is a core concept of evolutionary biology, its significance recognized by Charles Darwin (Darwin, Charles (1809–1882)) as a central tenet of his theory of evolution by natural selection (Darwin, 1856). Simply stated, differences in individuals of a species, or among species, may confer differential survival or performance, and thus influence which individuals survive and

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reproduce, and thereby the passing of favorable traits on to offspring. Where traits shift or adjust to suit local conditions, beneficial ones can be viewed as “adaptations,” and may confer evolutionary success. Adaptation is a whole-organism phenomenon. Change in a trait that confers benefits in one area but imposes costs or dysfunction in another may not lead to adaptation if the costs outweigh the benefits. Thus tradeoffs between traits are an essential part of the process of adaptation, placing constraints on what changes are beneficial. This is recognized in life history theory, which relates how organisms divide limited energy and resources to different functions and processes, how the balance between these may change with external conditions, and how these changes result in differential success among life history strategies (Stearns, 1992). Adaptation is such a wide-ranging process that generalizations about it are often not true in all conditions. For example, severe environmental conditions may pose strong selective pressures leading to rapid adaptation to environmental stress in that part of the population that survives (Hoffman and Parsons, 1991). In this case, adaptation occurs by extermination of “unfit” genes under rapid environmental change. On the other hand, benign conditions enable beneficial traits to accumulate in a population over successive generations. In this case, adaptation occurs by competitive success of beneficial genes over less-fit genes. Not all traits that occur in an organism are certain to be adaptive; they may simply not have a negative impact on the individual’s or species’ survival. Thus, it is necessary to determine specifically if a trait is adaptive through careful observation or experimentation, rather than by simply observing its presence (Gould and Lewontin, 1994). The evolution and taxonomy (Coral Cay Classification and Evolution) of extant reef corals provides key insights into the unusual evolutionary pressures faced by corals and hence of their capacity for adaptation. Coral reefs are typically considered to occur in relatively benign and stable environmental conditions (Corals: Environmental Controls on Growth), with high density and diversity of organisms. This creates conditions for high levels of niche diversification and diversification of interactions, and for these to become stable over time. Thus adaption, or coadaptation (among mutually interacting species), is common on reefs and can be distinguished in many forms of interactions, such as: 1. Primary production: Different functional groups of primary producers, characterized by whether they form hard crusts that resist herbivory and cement reefs, fast-growing low-biomass algal filaments and turfs with high recovery rates from removal, and large-bodied fleshy algal fronds that resist herbivory through low palatability and compete with other sessile organisms to monopolize space (Algae, Coralline; Algae-Macro; Algae, Turf ); 2. Predator–prey dynamics: A vast array of different prey and predators, and of defense and predation mechanisms. Adaptations of fish consumers are clear in their jaw structures toward their prey (Figure 1): scraping

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herbivores with hard fused teeth that can graze algal crusts, fish with long tubular mouths for sucking polyp tissue out of a coral skeleton, or top predators with needle-sharp pointed teeth for grabbing fast swimming prey and holding it until it can be torn and swallowed. Defense mechanisms may be even more numerous including chemical defenses of unpalatable algae and invertebrates that produce toxins, hard-shelled defenses such as in snails and crustaceans, or mimicry by palatable species of unpalatable or poisonous species often in completely different taxonomic groups; 3. Mutualisms: With limited space on a coral reef, numerous organisms adapt to live together cooperatively, in stark contrast to predator–prey or competitive interactions, in which one wins over the other. Many species live in or on one another, such as many types of crustaceans and anemones living on anemones and hard and soft corals, or fish and shrimp that share skills to build, maintain, and defend a burrow, or of microscopic single-celled algae such as zooxanthellae living in the tissue of a host such as a coral, to the mutual benefit of both.

Coral–zooxanthellae symbiosis as an illustration of adaptation The symbiosis between corals (the host, a sessile macroinvertebrate) and its zooxanthellae (the endosymbiont, an autotrophic single-celled algae) is the example par excellence of a type of partnership that has recurred among different partners in shallow tropical seas for hundreds of millions of years. The host is attached (sessile) on a shallow bottom, providing a sheltered nutrient-rich microhabitat in the photic zone for the endosymbiont. The endosymbiont fixes energy of sunlight into carbohydrates, which are passed to the host for consumption, and may also enhance the intracellular chemical environment for cellular processes of the host, such as calcification (Corals: Biology, Skeletal Deposition, and Reef-Building). The adaptations enabled by this symbiosis can be illustrated at physiological, organismal, and ecological–geological scales. Because the symbiosis combines two organisms – a photosynthetic algae and a consumer animal – two forms of energy capture and nutrition are possible. Under low-nutrient, high light conditions, photosynthesis is maximized, tight nutrient-cycling between the symbiotic partners enables growth in both, and autotrophy is the dominant mode of energy capture for the holobiont. By contrast, in highly turbid environments with limited light, heterotrophy by the coral predominates. Under these conditions feeding by the coral on plankton, detritus, and dissolved organic matter (Corals: Environmental Controls on Growth) may compensate for limited autotrophy and enable corals to thrive. The reliance of different coral species on autotrophy vs. heterotrophy, and their ability to shift

Adaptation, Figure 1 Top – butterflyfish (Chaetodontidae) have long tubular mouths for sucking polyp tissue out of a coral skeleton, and a flattened shape for manouevering in narrow spaces between coral branches. Middle – hawfish (Cirrhitidae) live on and within coral colonies. Bottom – parrotfish (Scaridae) have hard fused teeth that can graze algal crusts and excavate the rock substrate.

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without intermediate growth forms being known, these may appear to be separate species (Veron, 2000). Further adaptations of growth form can relate to, for example, sediment shedding in high-sediment conditions. The ability of coral holobionts to calcify extremely efficiently and thereby produce durable skeletons has enabled successive generations to colonize the skeletons left by previous generations and thereby raise up above the substrate resulting in reef construction over time (Coral Reef, Definition). The specific growth forms and other adaptations of the dominant corals in these communities affect the morphology and development of Reef Structure. Over geological history different but functionally equivalent symbioses have recurred: rugose corals in the Palaeozoic; rudist bivalves in the Cretaceous; scleractinian corals in the Cenozoic. The fossil reefs they have left behind show characteristic adaptations to the periods in which they lived.

Adaptation, Figure 2 Growth form adaptations of Pocillopora damicornis, which forms thin delicate branches in calm and deeper conditions (top) and thick robust branches in shallow rough conditions (bottom).

between the two, differs, reflecting adaptations to different environments and ecological niches. At the organismal level, a classic example of the adaptability of corals is in the diverse growth forms possible in some species. For example, growth form may vary depending on hydrodynamics (affecting how robust the skeleton must be) or light and sediment regimes (affecting shape, orientation, and self-shading (Corals: Environmental Controls on Growth). Pocillopora damicornis is a fast-growing, opportunistic, branching coral species that illustrates this well (Figure 2): individuals growing in calm and light-limited conditions may have very fine branches (no selection for robust growth but strong selection to minimize self-shading) while individuals of the same species in rough, well-illuminated environments may have very robust branches (resistant to breakage, no need to minimize shading effects). Individual colonies from the extremes of these distributions may appear so different that

Current investigations, controversies, and gaps in current knowledge Coral bleaching The life history of symbiotic corals and the as-yet incompletely understood phenomenon of coral bleaching (Temperature Change: Bleaching) provide an interesting case study of adaptation in action. As currently understood, the coral host provides a safe habitat and nutrients to the endosymbiotic algae. They in turn contribute to a range of the coral host’s physiological processes by transferring energy from sunlight in the form of fixed carbon, and chemically facilitating a variety of intracellular processes, such as calcification (Muscatine, 1990). The symbiosis is obligate because coral species that are symbiotic do not successfully compete in nature when asymbiotic. However, the symbiosis can be disrupted temporarily as happens under stress (e.g., heat, cold, hyposalinity) when the symbionts part company and the coral “bleaches,” or turns white (Figure 3). This occurs by a reduction in the photosynthetic capacity of the holobiont (by reduction in cholorophyll concentration in individual symbionts, and/or by reduction in symbiont densities), primarily to counter the damaging effects of overproduction of free oxygen radicals by the symbionts. As a stress response, bleaching must have some capacity for acclimatization and adaptation (Coles and Brown, 2003; Obura, 2009). The Adaptive Bleaching Hypothesis first expressed this idea as a mechanism that allows coral symbionts to adapt to changing environmental conditions (Buddemeier and Fautin, 1993), a position countered by other on the grounds of insufficient evidence (HoeghGuldberg, 2005). Further advances in the field may come through seeing bleaching as an extreme state of a range of symbiotic responses to changing environmental conditions (Obura, 2009). These include, from least to most severe: fluctuating symbiont densities, such as occur under

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the switch to different symbionts may be a temporary response to severe stress, with reversion to normal symbiont populations following a return to normal conditions. There are tradeoffs between high growth and reproduction versus slow growth and stress resistance that affect the bleaching response of corals, illustrating the adaptive dimensions of bleaching. Corals characterized by rapid growth and reproduction and thin coral tissues tend to bleach and die at lower levels of stress (e.g., Acropora, Pocillopora), while those characterized by slow growth and thick coral tissues tend to bleach and survive at higher levels of environmental stress (e.g., Porites) (Loya et al., 2001). Greater tissue thickness and larger polyp size may be adaptive through their shading of zooxanthellae, ameliorating stress from high light and temperatures.

Adaptation, Figure 3 Bleaching of corals has been portrayed as an adaptation for the coral–zooxanthellae symbiosis to resist stressful conditions, whereby zooxanthellae and/or chlorophyll is lost such that the normal color of the coral (left panel) fades such that the white skeleton becomes visible through the transparent coral tissue (right panel). All hard coral genera may bleach; shown here are the genera Pocillopora (top), Galaxea (middle), and Lobophyllia (bottom).

normal seasonal changes in the environment; shuffling between different clades of symbionts already in the coral, in response to more extreme environmental fluctuations; taking on new symbiont clades from the water column, after bleaching caused by severe stress. In the latter case,

Long-term change Global climate change (Climate Change and Coral Reefs) is proceeding in multiple environmental parameters critical to coral growth and survival. Water temperature and acidification of ocean waters are two of the fundamental ones (Hoegh-Guldberg et al., 2008), in addition to many others such as changes in sea level, storm tracks, wave regimes, precipitation, and terrestrial runoff that will affect reef growth. The ability of corals and zooxanthellae to adapt to the changes in these basic environmental parameters will fundamentally affect their ability to continue to grow and build reef structures. Increases in the frequency and severity of coral bleaching events globally are an indicator of water temperatures exceeding the temperature envelopes to which corals are historically adapted. The adaptive potential of bleaching and other regulatory processes is currently unknown (Hughes et al., 2003), however the adaptive basis for coral bleaching explained above, and scenarios for the degree of adaptation needed for corals to survive future change (Donner, 2009), provide tools for understanding this process as it unfolds. The adaptive potential of corals to seawater acidification is less known than that for temperature, and as a basic chemical parameter controlling calcification it may be that there is very little ability for corals and other calcifying marine organisms to adapt to more acidic conditions (Ocean Acidification, Effects on Calcification). In this time of global change, the adaptive capacity of corals will be a critical feature in determining how reef ecosystems respond to change. Bibliography Brown, B. E., 1997. Adaptations of reef corals to physical environmental stress. Advances in Marine Biology, 31, 220–299. Buddemeier, R., and Fautin, D.,1993. Coral bleaching as an adaptive mechanism. Bioscience, 43, 320–326. Coles, S., and Brown, B. E., 2003. Coral bleaching – capacity for acclimatization and adaptation. Advances in Marine Biology, 46, 183–224. Darwin, C. R., 1856. On the origin of species by means of natural selection, or the preservation of favoured races in the struggle for life. London: Murray.

AERIAL PHOTOGRAPHY OF CORAL REEFS

Donner, S. D., 2009. Coping with commitment: projected thermal stress on Coral Reefs under different future scenarios. PLoS one, 4(6), e5712, doi:10.1371/journal.pone.0005712. Gould, S., and Lewontin, L., 1994. The spandrels of San Marco and the Panglossian paradigm – a critique of the adaptationist programme. In Sober, E. (ed.), Unifying Concepts in Ecology. Cambridge: Massachussetts, MIT Press. Hoegh-Guldberg, O., (ed.), 2005. Understanding the stress response of corals and symbiodinium in a rapidly changing environment (workshop proceedings). May 10–June 3 2005. Unidad Académica Puerto Morelos, Instituto de Ciencias del Mary Limnología, UNAM Mexico. Hoegh-Guldberg, O., Mumby, P. J., Hooten, A. J., Steneck, R. S., Greenfield, P., Gomez, E., Harvell, C. D., Sale, P. F., Edwards, A. J., Caldeira, K., Knowlton, N., Eakin, C. M., Iglesias-Prieto, R., Muthiga, N., Bradbury, R. H., Dubi, A., and Hatziolos, M. E., 2008. Coral reefs under rapid climate change and ocean acidification. Science, 318, 1737–1742. Hoffman, A., and Parsons, P., 1991. Evolutionary Genetics and Environmental Stress. Oxford: Oxford University Press, 284 pp. Hughes, T. P., Baird, A. H., Bellwood, D. R., Card, M., Connolly, S. R., Folke, C., Grosberg, R., Hoegh-Guldberg, O., Jackson, J. B., Kleypas, J., Lough, J. M., Marshall, P., Nyström, M., Palumbi, S. R., Pandolfi, J. M., Rosen, B., and Roughgarden, J., 2003. Climate change, human impacts, and the resilience of coral reefs. Science, 301, 929–933. Loya, Y., Sakai, K.,Yamazato, K., Nakano, Y., Samabali, H., and van Woesik, R., 2001. Coral bleaching: the winners and the losers. Ecology Letters, 4, 122–131. Muscatine, L., 1990. The role of symbiotic algae in carbon and energy flux in reef corals. Coral Reefs, 25, 1–29. Obura, D. O., 2009. Corals bleach to resist stress. Marine Pollution Bulletin, 58, 206–212, DOI 10.1016/j.marpolbul.2008.10.002. Stearns, S., 1992. The Evolution of Life Histories. Oxford: Oxford University Press. Veron, J., 2000. Corals of the world. Townsville: Australian Institute of Marine Science, 489 pp.

Cross-references Algae, Coralline Algae-Macro Algae, Turf Climate Change and Coral Reefs Coral Reef, Definition Corals: Biology, Skeletal Deposition, and Reef-Building Corals: Environmental Controls on Growth Darwin, Charles (1809–1882) Ocean Acidification, Effects on Calcification Porites Reef Structure Temperature Change: Bleaching

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earliest to take advantage of remote sensing techniques (Hopley, 1978). Both aircraft and balloons (e.g., Rützler, 1978) formed the initial platforms, usually for vertically mounted cameras using black and white film. On the Great Barrier Reef (GBR), the earliest vertical aerial photography was in 1925, when the Royal Australian Air Force photographed the Low Isles at a scale of 1:2,400 in 1928 for the Yonge Expedition (see Great Barrier Reef Committee). Simultaneously, Umbgrove (1928, 1929) was photographing reefs in Indonesia to aid the extensive work he was carrying out there. Aerial photography was used in many reef studies for the next 50 years, though systematic approaches were rare (Steers, 1945; Teichert and Fairbridge, 1948, 1950). Color photography was used in some areas, especially those related to tourism projects. Between 1964 and 1972, the whole of the GBR was photographed at scales between 1:50,000 and 1:80,000 (later to be used in combination with satellite imagery in providing the detail for the first zoning maps of the GBR Marine Park, Hopley et al., 1989). In the 1970s, experiments were made with emulsions outside the visible range. For example, it was found that the near infrared part of the spectrum (0.7–0.86 µm), though giving poor water penetration, uses its whole tonal range over only a meter or so of water depth, i.e., was ideal for mapping exposed reef flats at low tide. Moreover, the zooxanthellae within the coral tissue, like terrestrial vegetation, provides a very strong reflective signal (Hopley and van Steveninck, 1977) (Figure 1). Initial experiments were made using both color and near infrared film from flying heights as low as 1,000 ft (Linfoot and Thamrongnawasawat, 1993; Thamrongnawasawat and Catt, 1994; Thamrongnawasawat and Hopley, 1995), providing pixels of 18 m. In a comparison of techniques for mapping reefal habitats, Mumby and Green (2000) quote

Aerial Photography of Coral Reefs, Figure 3 Color near the infrared photograph of Pandora Reef near Townsville, GBR on a medium low tide. The full tonal range of the IR reflectance is used on the features, which are exposed or have a water cover of
1 m including living coral margins, shingle ridges, and reef flat pools with depths up to 1 m.

81% accuracy for CASI compared with the levels of 57% for 1:10,000 color aerial photography and 200 m off San Salvador Seamount, Bahamas showing the deepest crustose coralline.

Algae, Coralline, Figure 9 In situ images of Target Pathogen in the Pacific.

to depths as great as 274 m (Figure 10, Littler and Littler, 1994). The great abundances of corallines in the poorly known deep-sea realm underscore their widespread contributions to productivity, the marine food web, sedimentology, and reef biogenesis in clear tropical seas. The diversity of coralline algal forms is astonishing, ranging from small filamentous strands to some of the larger and most beautiful head-forming organisms on coral reefs

(Figures 1–3). The predominant members of this functional indicator group (Littler and Littler, 2007) tend to be slow-growing competitively subordinate taxa abundant in most reef systems. As a indicator group they are functionally resilient and able to expedite the recovery/restoration of a particular coral-reef system relatively quickly, given that some thin forms of crustose coralline algae accelerate colonization and chemically attract and facilitate the survival of coral larvae (Harrington et al., 2004); whereas, the other two fleshy-algal functional indicator groups (i.e., turfs, macroalgae) tend to overgrow and inhibit coral settlement and survival. Because most crustose coralline algae continually slough-off upper surface layers (Figure 2), they play a key role in physically preventing the settlement and colonization of many fleshy fouling organisms on coral reefs (Littler and Littler, 1997). Crustose corallines, because of their slow growth, tolerate a wide-range of nutrient levels and generally are conspicuous, but not as predominant as corals, under low concentrations of nutrients and high levels of herbivory (Littler and Littler, 2007). Accordingly, they do well in the presence of both low and elevated nutrients [i.e., most are not inhibited by nutrient stress and many are maintained competitor-free by surface cell-layer shedding (Johnson and Mann, 1986), even at lower levels of grazing (Littler and Littler, 1997)]. Therefore, crustose coralline algae do not require elevated nutrients as might be inferred

ALGAE, CORALLINE

(Littler and Littler, 2007); instead, the degree to which they rise to dominance is largely controlled indirectly by the factors influencing the abundances of other functional groups, primarily the corals and fleshy macroalgae. The key point is that crustose corallines dominate mainly by default (i.e., under conditions of minimal competition), where corals are inhibited (e.g., by elevated nutrients or by strong wave action), and where fleshy algae are removed by intense herbivory. The wave-pounded intertidal algal ridges are built predominantly by Porolithon (Hydrolithon) onkodes, P. gardineri, P. craspedium, and Lithophyllum kotschyanum in the Indo-Pacific and P. pachydermum, and L congestum in the Atlantic; all are coralline species that appear uniquely tolerant of aerial exposure. The transition from frondose- to turf- to coralline-algal communities has been reported (Steneck, 1989) to closely correlate with increasing herbivory gradients on coral reefs. In addition to their protective reef-building nature, coralline algae provide a number of other goods and services. Since the eighteenth century, unattached corallines (maërl) have been harvested as acid-soil pH conditioners. In Britain and France, hundreds of thousands of tons of Phymatolithon calcareum and Lithothamnion corallioides continue to be dredged annually. Enormous maërl beds, several km2 in area, mainly composed of species belonging to the genera Lithothamnion and Lithophyllum, are present off the coast of Brazil and have begun to be commercially harvested. Maërl is also used as a mineral food additive for cows, hogs, and other livestock, as well as in the filtration and neutralization of acidic drinking water. Corallines are used in modern medical science in the preparation of dental bone implants (Shors, 1999). The cellular carbonate skeleton provides an ideal matrix for the adherence and regeneration of bone and tooth structures. Coralline algal fossils have proven to be extremely beneficial in deriving paleoecological and paleoclimatic information, and also have been employed as stratigraphic markers of particular significance in petroleum geology. As a spectacularly colorful component of live rock for the flourishing marine aquarium trade, coralline algae are highly desired for their architectural and attractive aesthetic qualities. However, the most important contribution of coralline algae worldwide may well prove to be in ameliorating the greenhouse carbon dioxide buildup associated with global climate change. It is the balance between calcification and respiration – which produce carbon dioxide – and the consumption of CO2 by photosynthesis that will determine whether corallines act as a “sink” (absorbing CO2) or as a source of CO2. Experiments that studied how various calcifying systems take up and give off carbon dioxide have shown that the rise in CO2 produced by calcification is mitigated by its removal through increased photosynthesis (Ohde, 1995; Iglesias-Rodriguez et al., 2008), with a net effect that is unlikely to either contribute greatly or significantly reduce the rise in atmospheric CO2. However, rising levels of CO2 and concomitant acidification of seawater inhibit all

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reef builders, including coralline algae (Kleypas et al., 1999). By binding vast accumulations of CaCO3 during calcification and photosynthesis, corallines may play a role in slowing future acidification of marine habitats such as coral reefs.

Bibliography Agassiz, A., 1888. Three Cruises of the United States Coast and Geodetic Survey Steamer “Blake.” Boston: Houghton, Mifflin Co, Vol. 2, 220 p. Aguirre, J., Riding, R., and Braga, J. C., 2000. Diversity of coralline red algae: origination and extinction patterns from the Early Cretaceous to the Pleistocene. Paleobiology, 26(4), 651–667. Bailey, J. C., 1999. Phylogenetic positions of Lithophyllum incrustans and Titanoderma pustulatum (Corallinaceae, Rhodophyta) based on 18S rRNA gene sequence analyses, with a revised classification of the Lithophylloideae. Phycologia, 38, 208–216. Bory de Saint-Vincent, J. B., 1832. Notice sur les polypiers de la Grèce. Expédition Scientifique de Morée (Section des Sciences Physiques), 3(1), 204–209, pl. 54. Darwin, C. R., 1842. The Structure and Distribution of Coral Reefs. Being the First Part of the Geology of the Voyage of the Beagle, Under the Command of Capt. FitzRoy, R.N. During the Years 1832 to 1836. London: Smith Elder and Co, pp. i–xii, 1–214, pls I–II. Dawson, E. Y., 1961. The rim of the reef. Natural History, 70, 8–17. Esteban, M., 1996. An overview of Miocene reefs from Mediterranean areas: general trends and facies models. In Franseen, E., Esteban, M., Ward, W. C., and Rouchy, J. M. (eds.), Models for Carbonate Stratigraphy from Miocene Reef Complexes of the Mediterranean Regions. Society of Economic Paleontologists and Miner, Concepts in Sedimentology and Paleontology Series, Vol. 5, pp. 3–53. Halfar, J., and Mutti, M., 2005. Global dominance of coralline redalgal facies: a response to Miocene oceanographic events. Geology, 33(6), 481–484. Halfar, J., Godinez-Orta, L., Mutti, M., Valdez-Holguin, J., and Borges, J., 2004. Nutrient and temperature controls on modern carbonate production: an example from the Gulf of California, Mexico. Geology, 32(3), 213–216. Harrington, L., Fabricius, K., De’Ath, G., and Negri, A., 2004. Recognition and selection of settlement substrata determine post-settlement survival in corals. Ecology, 85, 3428–3437. Iglesias-Rodriguez, M. D., Halloran, P. R., Rosalind E. M., Rickaby, R. E. M., Hall, I. R., Elena Colmenero-Hidalgo, E., Gittins, J. R., Green, D. R. H., Tyrrell, T., Gibbs, S. J., von Dassow, P., Rehm, E., Armbrust, E. V., and Boessenkool, K. P., 2008. Phytoplankton Calcification in a High-CO2 World. Science, 320(5874), 336–340. Johnson, C. R., and Mann, K. H., 1986. The crustose coralline alga, Phymatolithon Foslie, inhibits the overgrowth of seaweeds without relying on herbivores. Journal of Experimental Marine Biology and Ecology, 96(2), 127–146. Kleypas, J. A., Buddemeier, R. W., Archer, D., Gattuso, J. P., Langdon, C., and Opdyke, B. N., 1999. Geochemical consequences of increased atmospheric CO2 on coral reefs. Science, 284(5411), 118–120. Littler, M. M., 1973. The population and community structure of Hawaiian fringing-reef crustose Corallinaceae (Rhodophyta, Cryptonemiales). Journal of Experimental Marine Biology and Ecology, 11, 103–120. Littler, M. M., and Kauker, B., 1984. Heterotrichy and survival strategies in the red alga Corallina officinalis L. Botanica Marina, 27, 37–44.

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Littler, M. M., and Littler, D. S., 1994. Plant life of the deep ocean realm. Biologie in Unserer Zeit, 24(6), 330–335 [In German]. Littler, M. M., and Littler, D. S., 1995. CLOD spreading in the seasurface microlayer: response. Science, 270, 897. Littler, M. M., and Littler, D. S., 1997. Disease-induced mass mortality of crustose coralline algae on coral reefs provides rationale for the conservation of herbivorous fish stocks. In Proceedings of the Eighth International Coral Reef Symposium. Panama, pp. 719–724. Littler, M. M., and Littler, D. S., 1998. An undescribed fungal pathogen of reef-forming crustose coralline algae discovered in American Samoa. Coral Reefs, 17(2), 144. Littler, M. M., and Littler, D. S., 1999. Castles built by a chiton from the Great Astrolabe Reef, Fiji. Coral Reefs, 18(2), 146. Littler, M. M., and Littler, D. S., 2007. Assessment of coral reefs using herbivory/nutrient assays and indicator groups of benthic primary producers: a critical synthesis, proposed protocols, and critique of management strategies. Aquatic Conservation: Marine and Freshwater Ecosystems, 17, 195–215. Littler, M. M., Littler, D. S., and Brooks, B. L., 2007. Target phenomena on south Pacific reefs: strip harvesting by prudent pathogens? Reef Encounter, 34, 23–24. Littler, M. M., Littler, D. S., and Hanisak, M. D. 1991. Deep-water rhodolith distribution, productivity and growth history at sites of formation and subsequent degradation. Journal of Experimental Marine Biology and Ecology, 91, 1–20. Littler, M. M., Littler, D. S., and Taylor, P. R., 1995. Selective herbivore increases biomass of its prey: a chiton-coralline reefbuilding association. Ecology, 76(5), 1661–1681. Ohde, S., 1995. Calcium carbonate production and carbon dioxide flux on a coral reef, Okinawa. In Sakai, H., and Nozaki, Y. (eds.), Biogeochemical Processes and Ocean Flux in the Western Pacific. Tokyo: Terra Scientific Publishing Company (TERRAPUB), pp. 93–98. Philippi, R. A., 1837. Beweis, dass die Nulliporen Pflanzen sind. Archiv Für Naturgeschicthe, 3, 387–393, pl. 9, figs 2–6. Shors, E. C., 1999. Coralline bone graft substitutes. Orthopedic Clinics of North America, 30, 599–613. Silva, P. C., and Johansen, H. W., 1986. A reappraisal of the order Corallinales (Rhodophyceae). European Journal of Phycology, 21, 245–254. Sloane, H., 1707. A Voyage to the Islands, Madera, Barbados, Nieves, S. Christophers and Jamaica. London: Privately published, Vol. 1, 364 pp. Steneck, R. S., 1983. Quantifying herbivory on coral reefs: just scratching the surface and still biting off more than we can chew. In Reaka, M. L. (ed.), The Ecology of Deep and Shallow Coral Reefs. Symposia Series for Undersea Research, Vol. 1, pp. 1103–1112. Steneck, R. S., 1985. Adaptations of crustose coralline algae to herbivory: patterns in space and time. In Toomy, D., and Nitecki, M. (eds.), Paleoalgology. Berlin: Springer-Verlag, pp. 352–366. Steneck, R. S., 1989. Herbivory on coral reefs: a synthesis. In Proceedings of the Sixth International Coral Reef Symposium. Australia, Townsville, Vol. 1, pp. 37–49. Woelkerling, W. J., 1988. The Coralline Red Algae: An Analysis of the Genera and Subfamilies of Nongeniculate Corallinaceae. London: British Museum (Natural History).

Cross-references Algal Rims Atoll Islands (Motu) Barrier Reef (Ribbon Reef ) Calcite Coral Reef, Definition

Fossil Coralline Algae Fringing Reefs General Evolution of Carbonate Reefs Ocean Acidification, Effects on Calcification Pacific Coral Reefs: An Introduction Reef Front Wave Energy Reef Structure Rhodoliths Spurs and Grooves

ALGAE-MACRO Mark M. Littler, Diane S. Littler Smithsonian Institution, Washington, DC, USA

Synonyms Macroalgae; Macroscopic algae; Multicellular photosynthetic cryptogams; Seaweeds Definition Macroalgae are multicellular marine plants that are easily observed by the unaided eye and whose “thallus” (plant body) is characterized by holdfasts for attachment, and by “laminae,” reproductive “sori,” gas bladders, and/or stipes. They lack the various structures that characterize higher plants, such as true leaves, roots, and encased reproductive organs. Introduction Macroalgae (“seaweeds”) belong to either one of three groups of eukaryotic algae: the Rhodophyta (red algae), Chlorophyta (green algae), and Phaeophyceae (brown algae) or to the prokaryotic colony-forming Cyanobacteria/ Cyanophyta (blue-green algae). These four groups do not have a common multicellular ancestor (i.e., collectively they are polyphyletic); although, their chloroplasts – common to all – appear to have had a single blue-green algal (Cyanobacteria) origin. The presence of chloroplasts and subsequent capacity for photosynthesis give marine macroalgae an ecological role as primary producer that is similar to other marine plants, notably “seagrasses.” However, seagrasses are not seaweeds; rather, they are rooted, flower, and seed-bearing “higher” plants (Angiosperms). The macroalgal thallus (i.e., plant body) consists of blades (leaf-like lamina), reproductive sori (spore clusters), gas bladders [flotation organs (on blades in rockweeds, between lamina and stipes in kelps)], stipes [stem-like structures (may be absent)], and holdfasts [with or without haptera (finger-like extensions anchoring to substrates)]. The stipe and blade combined are known as the frond. Macroalgae grow attached to stable substrata in seawater (or brackish water) under light levels sufficient for photosynthesis. Seaweeds are most commonly found in shallow waters on rocky shores; however, the green algal group Bryopsidales includes rhizoidal forms adapted

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to proliferating in sedimentary environments. At the shallowest level are algae that inhabit the high-intertidal spray zone, whereas, the deepest living forms are attached to the seabed under nearly 295 m of water (Littler and Littler, 1994; see Algae, Coralline, Figure 10). The deepest macroalgae are calcified crustose coralline species.

Human utilization of macroalgae Macroalgae have a variety of uses. They are used extensively as food by coastal cultures, particularly in Southeast Asia. Seaweeds are harvested or cultivated for the extraction of alginate, agar, and carrageenan – gelatinous substances collectively known as hydrocolloids or phycocolloids. Colloids have great commercial importance, especially in the production of food additives. The gelling, water-retention, emulsifying, and other physical properties of colloids are critical to the food industry. Agar is used in foods such as candies, canned meats, desserts, bottled drinks, and gelatin molds. Carrageenan is used in the manufacture of salad dressings, condiments, dietary foods, as preservatives in canned meat and fish, milk products, and bakery goods. Alginates are utilized for many of the same uses as carrageenan, but are also used in the production of paper sizings, glues, colorings, gels, explosive stabilizers, fabric prints, hydro-spraying, and drill lubricants. Macroalgae have long been used as fertilizers and soil conditioners. Seaweeds are currently being investigated as sources of biodiesel and biomethane. Algal extracts are also widely used in toothpastes, cosmetics, and paints. In the biomedical and pharmaceutical industries, alginates are used in wound dressings and production of dental molds. In microbiological/diagnostic research, agar

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is extensively used as the culture substrate of choice. Seaweeds are also a source of iodine, an element necessary for thyroid function. The vast array of natural products that algae produce represents a potential “gold mine” of medicinal compounds that are still yielding promising results.

Ecological significance of macroalgae Top-down control of macroalgae by abundant populations of large mobile herbivores is particularly well studied for coral reefs, beginning over four decades ago with the caging study of Stephenson and Searles (1960). As examples, Sammarco et al. (1974), Ogden and Lobel (1978), Sammarco (1983), Carpenter (1986), Lewis (1986), Morrisson (1988), and numerous other workers (see review by McCook et al. (2001)) have demonstrated that lowering herbivory without increased nutrient inputs (usually assumed) mostly results in rapid increases in fleshy algae. However, when coral reefs are exposed to an increase in nutrients (bottom-up), fleshy macroalgae (Figure 1) may be favored over the slower growing but highly desirable corals (Lapointe et al., 1997). On healthy oligotrophic coral reefs, even very low nutrient increases may exceed critical levels that can shift relative dominances by stimulating macroalgal production, while inhibiting corals (Littler and Littler, 1984). Interestingly, large biomasses/standing stocks of slow-growing perennial macroalgae (e.g., rockweeds) can develop given sufficient time, even under low inorganic nutrient concentrations (McCook, 1999). Also, Sargassum spp. can coexist with corals in oligotrophic waters by utilizing particulate organic sources of nutrients (Schaffelke, 1999). Therefore, in this context, large macroalgal biomasses do not

Algae-Macro, Figure 1 Images of frondose macroalgae overgrowing corals.

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necessarily require or indicate detrimentally abundant dissolved nutrients. Fleshy macroalgae can outcompete corals (Birkeland, 1977; Bellwood et al., 2006), many of which are inhibited under elevated nutrient levels (reviewed in Marubini and Davies, 1996). Fast-growing macroalgae are opportunists that benefit from disturbances that release space resources from established longer-lived organisms. They can also take over space from living corals (Birkeland, 1977) when provided with sufficient nutrients. As a result, frondose macroalgae (those that form carpets of horizontal thalli) are generally recognized as harmful to coral reefs due to the linkage between excessive blooms and coastal eutrophication (ECOHAB, 1997). The competitive dominance of fast-growing macroalgae is inferred from their overshadowing canopy heights, as well as from inverse correlations in abundances between algae and other benthic producers (Lewis, 1986), particularly under elevated nutrient concentrations (e.g., Littler et al., 1993; Lapointe et al., 1997). Macroalgae, such as Halimeda spp. (Figure 2), also can gain competitive advantage by serving as carriers of coral diseases (Nugues et al., 2004). The fleshy macroalgal form-group has proven to be particularly attractive to herbivores (see Hay, 1981; Littler et al., 1983a, b) and only becomes abundant where grazing is lowered or swamped by excessive algal growth [chemically defended forms such as Cyanobacteria (Figure 3; Paul et al., 2007) are exceptions]. Overcompensation by high levels of herbivory may explain some of the reported cases (e.g., Smith et al., 2001) of specific corals surviving high-nutrient coral-reef environments.

Major macroalgal groups Rhodophyta (red algae) Rhodophyta are generally some shade of red, the result of having large quantities of the red pigment phycoerythrin

in their photosynthetic cells. This red pigment in combination with various other pigments is responsible for the vast array of colors ranging from translucent pale pink, lavender, purple, maroon, burgundy to iridescent blue (Figure 4), but generally with some hint of red. The pigment phycoerythrin is water-soluble; therefore, red algae immersed in hot water will stain the liquid red or pink and the thalli will eventually turn green. Other red-algal characters are eukaryotic cells lacking motile gametes (without flagella and centrioles), floridean starch as the food reserve, and chloroplasts containing unstacked thylakoids without an external endoplasmic reticulum. Pit connections and pit plugs are unique and distinctive features of red algae that form during the process of cytokinesis following mitosis. Most red algae are also multicellular, macroscopic, marine, and have sexual reproduction. They display alternation of life-history phases including a gametophyte phase and two sporophyte phases. The red algae are almost exclusively marine and comprise the most diversified and the largest group of tropical reef plants, with estimates of up to 10,000 species. Their diversity of forms is astonishing, ranging from small filamentous turfs to some of the larger and most beautifully delicate organisms on coral reefs (Figure 4). Calcareous red algae can dominate some reefs and often surpass corals in reef-building importance [e.g., Porolithon (Hydrolithon) craspedium, Figure 5]. Most often, corals (Cnidaria) supply the bulk building blocks, whereas, coralline algae do much of the cementing together of debris. The crustose coralline algae [forms that deposit a type of calcium carbonate (calcite limestone) that is harder and denser than the aragonite of corals] also build the “algal ridge” (see Algae, Coralline, Figure 5). The raised algal ridge, by absorbing tremendous wave energy, not only protects land masses that would otherwise erode, but also shelters the more delicate corals and other reef organisms.

Algae-Macro, Figure 2 Halimeda opuntia competing with coral.

Algae-Macro, Figure 3 Lyngbya polychroa, a chemically defended blue-green alga overgrowing Millepora.

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Algae-Macro, Figure 4 Array of different forms and colors of red algae.

Phaeophyceae (brown algae) The most frequent color within Phaeophyceae is some shade of brown, from pale beige to yellow-brown to almost black. This color is the result of large quantities of the brown pigment fucoxanthin. Fucoxanthin is alcohol soluble and the liquid will turn brown after immersion of a specimen, with the thallus eventually becoming green. They have cellulose walls, with alginic acid and fucoidin also important components. Brown algae are unique among macroalgae in developing into multicellular forms with differentiated tissues, and they reproduce by means of motile flagellated spores. Most brown algae have a life history that consists of an alternation between morphologically similar haploid and diploid plants. Scytosiphon lomentaria alternates between four distinct morphological generations, which is considered to be a “bet-hedging” survival strategy (Littler and Littler, 1983). The Phaeophyceae comprise about 2,000 species and are almost exclusively marine algae and, as kelps (Laminariales), attain their greatest abundance, size, and diversity in cold temperate to polar waters. They occur from the high intertidal (Fucales) to 115 m deep (Sargassum hystrix, Littler and Littler, 1994). Tropical brown algae include microscopic filaments, sheets, coarsely branched, and crust forms. Nearly all brown algae have fine (microscopic) hairs emanating from their surfaces that may serve to increase surface area for nutrient uptake. Brown algae are also well represented and very important on coral-reef ecosystems, with certain species of rockweeds able to dominate the fleshy algal biomass in

back-reef areas. for example, Sargassum and Turbinaria, growing just behind the reef crest, can form small-scale forests up to several meters high that provide biomass, habitat, and shelter for numerous fishes and invertebrates. Interestingly, we have observed vast drifting rafts of floating Sargassum in Fiji that were reminiscent of the Atlantic Ocean’s Sargasso Sea accumulations.

Chlorophyta (green algae) The most representative color of the Chlorophyta is some shade of green, the result of having large quantities of predominantly green chlorophyll pigments. Chlorophyll is also present in all of the other algal divisions; however, members of those groups have additional pigments that often overshadow and mask the green chlorophyll color. The green algae, in additional to chlorophylls a and b, also contain various subordinate carotenoid and xanthophyll pigments. The green algae, with upwards of 7,000 species, are the ancestral relatives of vascular plants (grasses, trees, seagrasses, etc.), which also contain these same basic pigments. Green seaweeds range from microscopic threadlike filaments to thin sheets, while others are spongy, gelatinous papery, leathery, or brittle in texture, ranging to 1.5 m in length (Figure 6). The green algae store their energy reserves as starch; therefore, a drop of potassium iodide (tincture of iodine) on a branch or blade will stain blue-black (iodine is taken up by starch granules). All produce flagellated spores and gametes giving them the advantage of motility.

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other sediments due to continual sloughing and following disintegration. In many tropical locales, the sparkling white sand beaches are mostly bleached and eroded calciumcarbonate (aragonite) skeletons of Halimeda. “Halimeda hash” (i.e., the coarse oatmeal-like accumulations of Halimeda segments, Figure 7) has been used in powergenerating plants and other fossil-fueled industries as smoke-stack scrubbers/neutralizers to precipitate sulfurous acid and other precursors to acid rain.

Algae-Macro, Figure 5 The reef-building calcareous macrophyte Porolithon (Hydrolithon) craspedium.

Green algae are always present on tropical coral reefs and lagoon floors, often intermixed among seagrass shoots. These are the siphonaceaous (giant-celled) forms of Bryopsidales, such as Halimeda, Avrainvillea, Udotea, and Caulerpa that employ a unique cytoplasmic streaming/blade abandonment mechanism to eliminate epiphytes (Littler and Littler, 1999). Most Bryopsidales have a rhizophytic “rooted” growth form and readily take-up pore water nutrients by cytoplasmic streaming (Williams, 1984). The deepest occurring fleshy upright alga (Rhipiliopsis profunda) is a member of this group and was found by submersible attached to bedrock at a depth of 210 m (Littler and Littler, 1994). Some genera of filamentous or sheet-like green algae are extremely tolerant of stressful conditions and can be indicators of fresh-water seeps, recently disturbed areas (as early colonizers of newly exposed substrates), habitats of low herbivory (high herbivory eliminates palatable greens), and especially areas with an overabundance of nutrients (e.g., bird roosting islands, polluted areas). Calcified green algae are major contributors to the production of marine sediments. Some genera, such as Udotea and Penicillus, produce enormous amounts of fine silt and

Cyanobacteria (blue-green “algae”) This ancient, highly controversial, and difficult group is a “prokaryote,” not a true plant. The Cyanobacteria were the first group to evolve photosynthesis, the process that powers the biological world. On tropical reefs, they comprise masses of microscopic organisms that are strung together into large filamentous clumps or colonies (Figure 8). In life, most of these large aggregations have distinctive colors, shapes, or growth forms that provide distinctive recognition features. However, these are lost in preserved specimens, and thus went unappreciated by earlier museum/herbarium-bound taxonomists. Most commonly, the color of blue-green algae is some peculiar shade of pink to purple to black – a combination of red from the pigment phycoerythrin, blue from phycocyanin, and green from chlorophyll. Colonies may form filamentous tufts, sheets, or globular spheres (Figure 8). Some filamentous colonies show the ability to differentiate into several specialized cell types: vegetative cells (the normal, photosynthetic cells that are formed under favorable growing conditions), akinetes (the stress-resistant long-lived spores that form when environmental conditions become harsh), and thickwalled heterocysts, which contain the enzyme nitrogenase for nitrogen fixation (see below; Herrero and Flores, 2008). Heterocysts also form under specific environmental conditions (anoxia, hypoxia) or where nitrogen is limiting. Many Cyanobacteria also form motile reproductive filaments called hormogonia that glide free from the parent colony at special weaker cells (necridia) and disperse to form new colonies. Like the other groups of seaweeds, excessive standing biomass of Cyanobacteria is usually considered detrimental to the health of coral-reef systems and people. They produce chemical compounds that can be toxic to fish, plankton, and invertebrates. For example, swimmers’ itch, a skin irritation that beach goers commonly experience, is caused by blooms of the blue-green alga Lyngbya majuscula (Figure 9). Black-band disease of corals (Figure 10), found throughout all tropical oceans, is caused by blue-green algae and associated microorganisms (Ruetzler et al., 1983). Certain Cyanobacteria produce neurotoxins, hepatotoxins, cytotoxins, and endotoxins that can be dangerous to animals and humans (Paul et al., 2007). Several cases of human- and many cases of livestock-poisoning have been documented. The nitrogen fixing capacity of some blue-green algae is extremely important, although often overlooked.

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Algae-Macro, Figure 6 Array of different green algal forms.

Algae-Macro, Figure 7 Halimeda “hash”; i.e., dead calcareous segments.

Heterocyst-forming species bind nitrogen gas into ammonia (NH3), nitrite (NO2), or nitrate (NO3) that can be absorbed by all plants. This role is crucial for tropical reef systems and especially nutrient-depauperate atoll reefs, which are extremely low in “fixed” nitrogen. Some of

these organisms contribute significantly to global ecology and the oxygen cycle. For example, the marine cyanobacterium Prochlorococcus (0.5–0.8-µm diameter spheres) accounts for >50% of the total photosynthetic production of the open ocean and 20% of the planet’s atmospheric oxygen (Partensky et al., 1999). Cyanobacteria are the only group of organisms that are able to reduce nitrogen and carbon in aerobic conditions, a feature that may be responsible for their evolutionary and ecological success. Blue-green algae are abundant worldwide and ubiquitous on coral reefs, where they often occur under extreme environmental conditions. The universally present black band in the splash zones that make rocks or boat ramps slippery is a layer of microscopic blue-green algae. Such blue-greens can withstand exposure to severe drying, extreme salinity, rain water, bright sun, and high heat and still flourish. Cyanobacteria are the oldest known life forms on earth. Stromatolites containing fossilized oxygenproducing Cyanobacteria date to 1.5 billion years ago (Zhang and Golubic, 1987). The ability of Cyanobacteria to perform oxygenic photosynthesis is thought to have converted the early reducing atmosphere of Earth into an oxidizing one. Chloroplasts – the organelles responsible for photosynthesis in all higher plants and eukaryotic algae – evolved from Cyanobacteria via endosymbiosis.

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Algae-Macro, Figure 8 Array of different blue-green algal forms.

Algae-Macro, Figure 9 Raft of the noxious blue-green alga, Lyngbya majuscula. Algae-Macro, Figure 10 The black-band disease, Phormidium corallyticum, attacking a brain coral.

This evolutionary step dramatically changed the composition of life forms on Earth, triggering an explosion of biodiversity and leading to the near-extinction of oxygenintolerant organisms.

Summary Marine macroalgae are among the oldest and most significant inhabitants of coral reefs. Because of the rapid degradation of tropical reefs worldwide, it is imperative that the role and diversity of macroalgae be studied in a timely,

efficient, and scientifically verifiable manner. It is of paramount importance to characterize the world’s coral-reef environments and to understand the responses of these foundation species. The fleshy macroalgal forms are the food of herbivores, and only become abundant when their rate of production exceeds the capacity of herbivores to consume them. On healthy oligotrophic coral reefs, even very low nutrient increases may shift relative dominance from corals to macroalgae by both stimulating macroalgal

ALGAE-MACRO

production and inhibiting corals (Marubini and Davies, 1996). As a result, frondose macroalgae as a group are generally recognized as harmful to the longevity of coral reefs due to the linkage between excessive blooms and coastal eutrophication. Reef plant complexity has evolved along very different evolutionary lines. The range of sizes, shapes, life histories, pigments, and biochemical and physiological pathways is remarkable. The biodiversity of coral-reef plant life is unequalled. Macroalgae from four evolutionary lines dominate and, in conjunction with coelenterate corals, are the major primary producers and builders of coral-reef habitats and carbonate architecture. Previously, marine plants have been understudied on coral reefs. However, their rapid growth and short generation time make them ideal subjects for experimental studies. Some are commercially valuable and/or preferred table fare for many humans. Marine plants are essential to the world’s biogeochemical cycles and serve as potentially important sources of pharmaceuticals. The critical role that seaweeds play in reef ecosystems overlaps with other fields of marine sciences, such as fisheries resources, marine chemistry, ecology, geology, and coral-reef conservation.

Bibliography Bellwood, D. R., Hughes, T. P., and Hoey, A. S., 2006. Sleeping functional group drives coral reef recovery. Current Biology, 16, 2434–2439. Birkeland, C., 1977. The importance of rate of biomass accumulation in early successional stages of benthic communities to the survival of coral recruits. Proceedings of the Third International Coral Reef Symposium, 1, 15–21. Carpenter, R. C., 1986. Partitioning herbivory and its effects on coral reef algal communities. Ecological Monographs, 56, 345–363. ECOHAB., 1997. The ecology and oceanography of harmful algal blooms – a national research agenda. In Anderson, D. M. (ed.), Proceedings of the National Workshop, Massachusetts: WHOI, pp. 1–66. Hay, M. E., 1981. Herbivory, algal distribution, and the maintenance of between-habitat diversity on a tropical fringing reef. The American Naturalist, 118, 520–540. Herrero, A., and Flores, E. (eds), 2008. The Cyanobacteria: Molecular Biology, Genomics and Evolution (1st ed.). Norfolk, UK: Caister Academic Press. Lapointe, B. E., Littler, M. M., and Littler, D. S., 1997. Macroalgal overgrowth of fringing coral reefs at Discovery Bay, Jamaica: bottom-up versus top-down control. Proceedings of the Eighth International Coral Reef Symposium, 1, 927–932. Lewis, S. M., 1986. The role of herbivorous fishes in the organization of a Caribbean reef community. Ecological Monographs, 56, 183–200. Littler, M. M., and Littler, D. S., 1983. Heteromorphic life history strategies in the brown alga Scytosiphon lomentaria (Lyngb.). Journal of Phycology, 19(4), 425–431. Littler, M. M., Littler, D. S., and Taylor, P. R., 1983a. Evolutionary strategies in a tropical barrier reef system: functional-form groups of marine macroalgae. Journal of Phycology, 19, 229–237.

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Littler, M. M., Taylor, P. R., and Littler, D. S., 1983b. Algal resistance to herbivory on a Caribbean barrier reef. Coral Reefs, 2, 111–118. Littler, M. M., and Littler, D. S., 1984. Models of tropical reef biogenesis: the contribution of algae. In Round, F. E., and Chapman, D. J. (eds.), Progress in Phycological Research, Vol. 3, Bristol: Biopress, pp. 323–364. Littler, M. M., Littler, D. S., and Lapointe, B. E., 1993. Modification of tropical reef community structure due to cultural eutrophication: the southwest coast of Martinique. Proceedings of the Seventh International Coral Reef Symposium, 1, 335–143. Littler, M. M., and Littler, D. S., 1994. Plant life of the deep ocean realm. Biologie in Unserer Zeit, 24(6), 330–335 (In German). Littler, M. M., and Littler, D. S., 1999. Blade abandonment/proliferation: a novel mechanism for rapid epiphyte control in marine macrophytes. Ecology, 80(5), 1736–1746. Marubini, F., and Davies, P. S., 1996. Nitrate increases zooxanthellae population density and reduces skeletogenesis in corals. Marine Biology, 127, 319–328. McCook, L. J., 1999. Macroalgae, nutrients and phase shifts on coral reefs: scientific issues and management consequences for the Great Barrier Reef. Coral Reefs, 18, 357–367. McCook, L. J., Jompa, J., and Diaz-Pulido, G., 2001. Competition between corals and algae on coral reefs: a review of evidence and mechanisms. Coral Reefs, 19, 400–417. Morrisson, D., 1988. Comparing fish and urchin grazing in shallow and deeper coral reef algal communities. Ecology, 69, 1367–1382. Nugues, M. M., Smith, G. W., van Hooidonk, R. J., Seabra, M. I., and Bak, R. P. M., 2004. Algal contact as a trigger for coral disease. Ecology Letters, 7, 919–923. Ogden, J. C., and Lobel, P. S., 1978. The role of herbivorous fishes and urchins in coral reef communities. Environmental Biology of Fishes, 3, 49–63. Partensky, F., Hess, W. R., and Vaulot, D., 1999. Prochlorococcus, a marine photosynthetic prokaryote of global significance. Microbiology and Molecular Biology Reviews, 63, 106–127. Paul, V. P., Arthur, K. E., Ritson-Williams, R., Ross, C., and Sharp, K., 2007. Chemical defenses: from compounds to communities. Biological Bulletin, 213, 226–251. Ruetzler, K., Santavy, D. L., and Antonius, A., 1983. The black band disease of Atlantic reef corals, III: distribution, ecology, and development. Marine Ecology, 4, 329–358. Sammarco, P. W., 1983. Effects of fish grazing and damselfish territoriality on coral reef algae. I. Algal community structure. Marine Ecology Progress Series, 13, 1–14. Sammarco, P. W., Levinton, J. S., and Ogden, J. C., 1974. Grazing and control of coral reef community structure by Diadema antillarum Phillipi (Echinodermata: Echinoidea): a preliminary study. Journal of Marine Research, 32, 47–53. Schaffelke, B., 1999. Particulate nutrients as a novel nutrient source for tropical Sargassum species. Journal of Phycology, 35, 1150–1157. Smith, J. E., Smith, C. M., and Hunter, C. L., 2001. An experimental analysis of the effects of herbivory and nutrient enrichment on benthic community dynamics on a Hawaiian reef. Coral Reefs, 19, 332–342. Stephenson, W., and Searles, R. B., 1960. Experimental studies on the ecology of intertidal environments at Heron Island. I. Exclusion of fish from beach rock. Australian Journal of Marine and Freshwater Research, 2, 241–267. Williams, S. L., 1984. Uptake of sediment ammonium and translocation in a marine green macroalga Caulerpa cupressoides. Limnology and Oceanography, 29(2), 374–379. Zhang, Y., and Golubic, S., 1987. Endolithic microfossils (Cyanophyta) from early Proterozoic stromatolites, Hebei, China. Acta Micropaleont Sinica, 4, 1–12.

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ALGAE, TURF

Cross-references Algae, Blue-Green Boring Algae, Coralline Algae, Turf Algal Rims Fossil Coralline Algae Halimeda Halimeda Bioherms Nutrient Pollution/Eutrophication Rhodoliths Seagrasses Stromatolites

ALGAE, TURF Mark M. Littler, Diane S. Littler Smithsonian Institution, Washington, DC, USA

Synonyms Algae turfs; Algal mats; Low-growing algae; Prostrate algae Definition Algae turfs are sparse to thick mats of diminutive and juvenile algae less than 2 cm high. Turf communities (Figure 1) are composed of juvenile macroalgae and faster-growing filamentous species

accompanied by the ubiquitous blue-greens, diatoms, and detrital sediments. These juvenile and microalgal species assemblages have a high diversity (>100 species in some western Atlantic turfs), although only 30–50 species commonly co-occur at any one time. Turf algal assemblages, when viewed as a functional indicator group (Littler and Littler, 2007), remain relatively stable year round. They are often able to recover rapidly after being partially consumed by herbivores. Algal turfs characteristically trap ambient sediments and smother corals and other competitors for space by gradual encroachment. Domination by turf algae suggests not only desirably low nutrient levels (bottom-up) but also an inadequate herbivory (top-down) component required for healthy coral-dominated reefs (Littler et al., 2006). Algal turfs have been shown to form extensive horizontal mats under reduced nutrient-loading rates (Fong et al., 1987) or infrequent nutrient inputs (Fujita et al., 1988). Their relatively small size and rapid regeneration from basal remnants (perennation) results in only moderate losses to herbivory at low grazing pressures. Turf algal community structure can be affected by the behavior of territorial damselfish. Because of their preferential grazing and protection (chasing off of larger herbivores), damselfish cultivate more diverse alga turfs that have higher biomass within their territories. Turf algae have opportunistic (“weedy”) life-history characteristics, including high surface area to volume ratios and the ability to maintain substantial nutrient uptake and growth rates

Algae, Turf, Figure 1 Examples of naturally occurring algal turf communities. Upper left: Asparagopsis taxiformis Sporophyte. Upper right: Gelidiopsis intricata. Lower left: Wurdemannia miniata. Lower right: Oscillatoria rosea.

ALGAL RIMS

under low-nutrient conditions (Rosenberg and Ramus, 1984). Turfs also contain populations of nitrogen-fixing Cyanobacteria (Adey, 1998) that can enrich the other low-growing members within the dense turf community in oligotrophic waters.

Bibliography Adey, W. H., 1998. Coral reefs: algal structured and mediated ecosystems in shallow, turbulent, alkaline waters. Journal of Phycology, 34, 393–406. Fong, P., Rudnicki, R., and Zedler, J. B., 1987. Algal community response to nitrogen and phosphorus loading in experimental mesocosms: management recommendations for Southern California lagoons. Report of the California State Water Control Board, pp. 88. Fujita, R. M., Wheeler, P. A., and Edwards, R. L., 1988. Metabolic regulation of ammonium uptake by Ulva rigida (Chlorophyta): a compartmental analysis of the rate-limiting step for uptake. Journal of Phycology, 24, 560–566. Littler, M. M., Littler, D. S., and Brooks, B. L., 2006. Harmful algae on tropical coral reefs: bottom-up eutrophication and top-down herbivory. Harmful Algae, 5(5), 565–585. Littler, M. M., and Littler, D. S., 2007. Assessment of coral reefs using herbivory/nutrient assays and indicator groups of benthic primary producers: a critical synthesis, proposed protocols, and critique of management strategies. Aquatic Conservation: Marine and Freshwater Ecosystems, 17, 195–215. Rosenberg, G., and Ramus, J., 1984. Uptake of inorganic nitrogen and seaweed surface area: volume ratios. Aquatic Botany, 19, 65–72.

Cross-references Algae, Turf Nutrient Pollution/Eutrophication

ALGAL RIMS Jacques Laborel Université de la Méditerranée, Marseille Cedex 9, France

Definition and morphology Algal rims are marine biogenic formations of various size and shape, generally edificated by Coralline algae associated with other organisms, developing upon the windward edge of coral reefs or rocky coasts in tropical and subtropical seas. Thin reef-like structures (often referred to as biostromes) may develop on the outer edge of reef-flats or rocky windward coasts submitted to strong surf, both in tideless or tidal areas. They were first described from the Pacific (Tracey et al., 1948), and were subsequently found in the North Atlantic (Agassiz, 1895), Brazil (Kempf and Laborel, 1968) and the Caribbean area (Gessner, 1970; Adey and Burke, 1976) Related formations are also known from the Mediterranean (Blanc and Molinier, 1955). They are mainly built by massive or encrusting coralline algae (mostly Hydrolithon) Hydrocorals (Millepora spp.), Vermetid Gastropods and some corals. Specific

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composition varies with surf, slope and nature of substrate (Focke and Gebelein, 1978). Plant and animal populations mingle seaward with those of the reef’s outer slope; laterally, the rim may pass to the spur-and-groove structures or to rim-like formations developing directly on rocky shores. On coral reefs, algal rims often develop by fusion of algal heads. Extreme surf conditions may lead to the development of boilers, terraced pinnacles or blowholes. Similar morphological structures may be obtained by erosive processes. Bermudian “erosive boilers” generated from the erosion of an emerged stack of soft rock were described (Ginsburg and Schroeder, 1973).

Relation with sea level – Some biological components of algal rims (Dendropoma vermetids, some Lithophyllum and Hydrolithon) have a very narrow repartition around MSL; their presence in cores or on elevated shorelines is a precise indicator of past sea levels, with metric or decimetric approximation, and widely used around the world’s oceans (Adey, 1986; Pirazzoli et al., 1988; Laborel et al., 1994). Stony corals having a range of – 5 to10 m tend to be less accurate indicators. Bibliography Adey, W. H., and Burke, R. B., 1976. Holocene bioherms (algal ridges and bank barrier reefs) of the eastern Caribbean. Bulletin of the Geological Society of America, 87, 95–109. Adey, W. H., 1986. Coralline algae as indicators of sea-level. In Van de Plassche, O. (ed.), Sea-level Research: a Manual for the Collection and Evaluation of Data. Amsterdam: Free University of Amsterdam, pp. 229–279. Agassiz, A., 1895. A visit to the Bermudas in 1894. Bulletin of the Museum of comparative Zoology, Harvard. Coll. 26, 209–281. Blanc, J. J., and Molinier, R., 1955. Les formations organogènes construites superficielles en Méditerranée occidentale. Bulletin de l’ Instititut océanographique de Monaco, 1067, 1–26. Focke, J., and Gebelein, C., 1978. Marine lithification of reef rock and rhodolites at a fore-reef slope locality off Bermuda. Geologie en Mijnbouw, 57, 163–171. Gessner, F., 1970. Lithothamnium terrassen in Karibischen Meer. Internationale Revue der Gesamten Hydrobiologie, 55, 757–762. Ginsburg, R. N., and Schroeder, J. H., 1973. Growth and submarine fossilisation of algal cup reefs, Bermuda. Sedimentology, 20, 574–614. Ginsburg, R. N., and Schroeder, J. H., 1973. In Biology and Geology of Coral reefs. Jones and Endean (eds.). Academic press, Biology 1, 9:271–324. Kempf, M., and Laborel, J., 1968. Formations de Vermets et d’Algues calcaires des côtes du Brésil. Recueil des travaux de la Station Marine d’Endoume, 43, 9–23. Laborel, J., and Laborel - Deguen, F., 1994. Biological indicators of relative sea-level variation and of co-seismic displacements in the Mediterranean area. Journal of Coastal research, 10(2), 395–415. Laborel, J., Morhange, C., Lafond, R., Le Campion, J., Laborel – Deguen, F., and Sartoretto, S., 1994. Biological evidence of sea-level rise during the last 4500 years on the rocky coasts of continental southwestern France and Corsica. Marine Geology, 120, 203–223.

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ANTECEDENT PLATFORMS

Pirazzoli, P. A., Montaggioni, L. F., Salvat, B., and Faure, G., 1988. Late Holocene sea level indicators from twelve atolls in the central and eastern Tuamotus (Pacific Ocean). Coral reefs, 7(2), 57–68. Tracey, J. I., Ladd, J. S., and Hoffmeister, J. E., 1948. Reefs of Bikini, Marshall Islands. Bulletin of the Geological Society of America, 59, 861–878.

Cross-references Algae, Coralline Forereef/Reef Front Reef Front Wave Energy Spurs and Grooves

ANTECEDENT PLATFORMS Peter J. Davies University of Sydney, Sydney, NSW, Australia

Definition Antecedent Platforms are the surfaces, usually older reefs which have been exposed to the atmosphere, and which form the foundation for later reef growth. Introduction In the study of coral reefs, the term “antecedent platform” has a special place for three reasons. First, in various guises, it stood as a challenge to Darwin’s championing of subsidence as the principle factor in coral reef growth; second,

it stood against the need for glacial eustacy as a factor in reef growth; and third, as a special case (limestone platforms), and at a later time, it co-defined a new paradigm for the evolution of the foundations of modern reefs. The thinking on antecedent platforms moved therefore from the broad to the specific, from “the alternative” to “center-stage,” albeit coupled with other processes. This contribution traces these views accenting more the later process-related views but placed within the context of the earlier history.

In the beginning!! In the middle and early nineteenth century, ideas on reef evolution and growth were dominated by Darwin’s theory that subsidence played a critical role in the evolution of coral reefs from fringing to barrier to atolls (1842) (Figure 1). However, not all agreed on either the occurrence or primacy of subsidence in reef growth. A succession of contributors proposed that reefs accumulated on pre-existing reef platforms (Rein, 1870, 1881), volcanic foundations in the deep sea (Murray, 1880, 1887, 1889), or surfaces of submarine planation (Wharton, 1890, 1897; Gardiner, 1898, 1903, 1904; Agassiz, 1898, 1899). Indeed, Andrews (1900, 1902), for example, proposed that the Great Barrier Reef itself rested on a surface of submarine planation. All attempted to counter the need for subsidence in the evolution of coral reefs, and this continued even after the results of the Funafuti borehole were published, which showed that subsidence at any rate was a crucial factor operating in the evolution of at least that reef (Cullis, 1904), and

Antecedent Platforms, Figure 1 (a) The Darwinian evolution of coral reefs and (b) Daly’s glacial control theory of reef evolution. (Page 10 – Purdy’d 1974 paper.)

ANTECEDENT PLATFORMS

subsequent drilling throughout the Pacific has confirmed this (Schlanger, 1963). However, post-Funafuti, the importance of the antecedent platform was as a foil to another idea in understanding the growth of coral reefs, that is, glacial eustacy. In a series of controversial papers, Daly (1910, 1915, 1917, 1919, 1934) proposed that global glaciations had the effect of both lowering sea level and reducing sea surface temperatures, the effects of both substantially impacting the growth of coral reefs. His glacialcontrol theory, applied initially to atolls, used the glacial lowering of sea level to produce foundations, which were then planated by wave abrasion as the effective process in turning fringing reefs into barrier reefs and then into atolls (Figure 1b). These ideas were extended by Daly himself and by Vaughan (1914, 1919, 1923) to Florida and the Bahamas as well as to the Great Barrier Reef. While some of Daly’s evidence has been called into question (e.g., not all atolls have the same depth; his proposed processes and rates of processes are unrealistic, and his proposed timescales are wrong), there can be little doubt today that glacio-eustatic sea-level fluctuations have profoundly affected how we think about the nature of the platforms on which today’s coral reefs have grown. The next step was taken by Kuenen (1933, 1947) coupling glacially produced sea-level fluctuations and subsidence to sub-aereal and marine intertidal erosion to produce partially or completely planated antecedent platforms on which further reef development could occur. And then, in the 1940s, very important papers but in poorly distributed publications, Japanese scientists (Yabe, 1942; Asano, 1942; Tayama, 1952) coupled glacio-eustatic sea-level fall to the nature of the reef foundations for the first time. Thus began the germ of a new set of ideas. The first two papers were by Hoffmeister and Ladd (1944, 1945), which admitted to the reality of sea-level fall but placed the major emphasis on a suitable substrate for subsequent reef growth, that is, any bench or bank situated at a proper depth in the coral seas is a potential reef foundation. In the 1945 paper, however, they reported the results of experiments simulating the effects of rainfall on an exposed limestone surface (a slab of Solenhoffen limestone) as a way to explain the origin of raised atolls. They inferred from their crude experiments that the saucer shape of raised atolls may be related to solution. This was left to MacNeil (1954) who accepted Kuenen’s coupling of glacial lowering of sea level and subsidence but invoked sub-aereal erosion to produce the diagnostic annular rim of atolls, which was therefore inherited from a period of earlier sub-aereal erosion (Figure 2). The saucer-shaped basin was therefore the logical consequence of sub-aereal solution acting on exposed Pleistocene atoll foundations. MacNeil (op.cit) emphasizes the importance of limestone-solution processes involving surface and subsurface solution and re-deposition and surface case hardening in producing relief features inherited by subsequent reef growth during an ensuing sea-level rise. The accent was clearly placed on inheritance from an exposed eroded surface. Thus, atoll lagoons form on eroded lows, and atoll

41

rims occupy the surrounding highs. For whatever reasons, MacNeil’s ideas were not received with ultra-enthusiasm by the scientific community.

Antecedence post-1974 – the purdy revolution In 1974, Purdy (1974) published a seminal paper, inferring cause and effect in linking reef shape to karst-induced antecedent morphology. Drowned atolls reflect drowned karst topography; reef passes originate as drainage breaches in the solution rim; faroes are the karst product of breaching; peripheral islands are exposures of the fossil drainage divide, and spurs and grooves are expressions of lapies. Thus, the karst-induced differences in relief are perpetuated and indeed accelerated by growth, but reef growth per se has little to do with the basic configuration. Paradoxically, a paper proving that the Holocene reef off the north Jamaica coast was a mere mantling of low sealevel Pleistocene terraces (Goreau and Land, 1974) was published alongside Purdy’s karst revolutionary paper. In marked contrast to MacNeil’s ideas, the 1974 paper sparked a widespread rethink of the role of antecedence and for a number of reasons: (1) Purdy’s ideas applied to barrier reefs as well as atolls and therefore to a much wider scientific community, (2) he discussed processes that were important to the oil industry in developing porosity, once again opening the subject to a science community outside of the academic, (3) he coupled eloquent laboratory experiments to well-illustrated examples of mega-karst morphology from some of the world’s notable karst areas, and (4) he used well-illustrated seismic data coupled to drillhole data across a major barrier reef complex (effectively for the first time), which became well known through the work of Ginsburg and coworkers (Ginsburg and James, 1976; James and Ginsburg, 1976) and Rutzler and Macintyre (1982). Furthermore, when data was lacking he resorted to compellingly acceptable eloquent deductions. Like Hoffmeister and Ladd (1944), Purdy (1974) subjected limestone blocks to acid rain and produced features analogous to both karst and modern reef forms (Figure 2). He emphasized solution rims, enclosed depressions at various scales, conical karst, tower karst, and karst marginal plains as important natural and experimental features, which could have analogs in both atolls and barrier reef systems. He coupled glacially effected sea-level oscillations, subsidence, and sub-aereal erosion to explain their formation (Figure 3). Atoll morphology derives from the development of solution rims on emerged limestone masses; where rainfall is high, breaches in the karst rim give rise to subsequent atoll-passes, while in the interior, a conical karst may develop as antecedent foundations to lagoonal patch reefs; also, collapse dolines may form as a consequence of extensive subterranean dissolution, the forerunner of “blue holes” according to Purdy (1974). Alternately, where rainfall is low, conical depressions (solution dolines) form within the solution rim. The morphology of atolls and barrier reefs is solution determined rather than growth predicated (Purdy and Winterer, 2001) (Figure 4).

ANTECEDENT PLATFORMS

subsequent drilling throughout the Pacific has confirmed this (Schlanger, 1963). However, post-Funafuti, the importance of the antecedent platform was as a foil to another idea in understanding the growth of coral reefs, that is, glacial eustacy. In a series of controversial papers, Daly (1910, 1915, 1917, 1919, 1934) proposed that global glaciations had the effect of both lowering sea level and reducing sea surface temperatures, the effects of both substantially impacting the growth of coral reefs. His glacialcontrol theory, applied initially to atolls, used the glacial lowering of sea level to produce foundations, which were then planated by wave abrasion as the effective process in turning fringing reefs into barrier reefs and then into atolls (Figure 1b). These ideas were extended by Daly himself and by Vaughan (1914, 1919, 1923) to Florida and the Bahamas as well as to the Great Barrier Reef. While some of Daly’s evidence has been called into question (e.g., not all atolls have the same depth; his proposed processes and rates of processes are unrealistic, and his proposed timescales are wrong), there can be little doubt today that glacio-eustatic sea-level fluctuations have profoundly affected how we think about the nature of the platforms on which today’s coral reefs have grown. The next step was taken by Kuenen (1933, 1947) coupling glacially produced sea-level fluctuations and subsidence to sub-aereal and marine intertidal erosion to produce partially or completely planated antecedent platforms on which further reef development could occur. And then, in the 1940s, very important papers but in poorly distributed publications, Japanese scientists (Yabe, 1942; Asano, 1942; Tayama, 1952) coupled glacio-eustatic sea-level fall to the nature of the reef foundations for the first time. Thus began the germ of a new set of ideas. The first two papers were by Hoffmeister and Ladd (1944, 1945), which admitted to the reality of sea-level fall but placed the major emphasis on a suitable substrate for subsequent reef growth, that is, any bench or bank situated at a proper depth in the coral seas is a potential reef foundation. In the 1945 paper, however, they reported the results of experiments simulating the effects of rainfall on an exposed limestone surface (a slab of Solenhoffen limestone) as a way to explain the origin of raised atolls. They inferred from their crude experiments that the saucer shape of raised atolls may be related to solution. This was left to MacNeil (1954) who accepted Kuenen’s coupling of glacial lowering of sea level and subsidence but invoked sub-aereal erosion to produce the diagnostic annular rim of atolls, which was therefore inherited from a period of earlier sub-aereal erosion (Figure 2). The saucer-shaped basin was therefore the logical consequence of sub-aereal solution acting on exposed Pleistocene atoll foundations. MacNeil (op.cit) emphasizes the importance of limestone-solution processes involving surface and subsurface solution and re-deposition and surface case hardening in producing relief features inherited by subsequent reef growth during an ensuing sea-level rise. The accent was clearly placed on inheritance from an exposed eroded surface. Thus, atoll lagoons form on eroded lows, and atoll

41

rims occupy the surrounding highs. For whatever reasons, MacNeil’s ideas were not received with ultra-enthusiasm by the scientific community.

Antecedence post-1974 – the purdy revolution In 1974, Purdy (1974) published a seminal paper, inferring cause and effect in linking reef shape to karst-induced antecedent morphology. Drowned atolls reflect drowned karst topography; reef passes originate as drainage breaches in the solution rim; faroes are the karst product of breaching; peripheral islands are exposures of the fossil drainage divide, and spurs and grooves are expressions of lapies. Thus, the karst-induced differences in relief are perpetuated and indeed accelerated by growth, but reef growth per se has little to do with the basic configuration. Paradoxically, a paper proving that the Holocene reef off the north Jamaica coast was a mere mantling of low sealevel Pleistocene terraces (Goreau and Land, 1974) was published alongside Purdy’s karst revolutionary paper. In marked contrast to MacNeil’s ideas, the 1974 paper sparked a widespread rethink of the role of antecedence and for a number of reasons: (1) Purdy’s ideas applied to barrier reefs as well as atolls and therefore to a much wider scientific community, (2) he discussed processes that were important to the oil industry in developing porosity, once again opening the subject to a science community outside of the academic, (3) he coupled eloquent laboratory experiments to well-illustrated examples of mega-karst morphology from some of the world’s notable karst areas, and (4) he used well-illustrated seismic data coupled to drillhole data across a major barrier reef complex (effectively for the first time), which became well known through the work of Ginsburg and coworkers (Ginsburg and James, 1976; James and Ginsburg, 1976) and Rutzler and Macintyre (1982). Furthermore, when data was lacking he resorted to compellingly acceptable eloquent deductions. Like Hoffmeister and Ladd (1944), Purdy (1974) subjected limestone blocks to acid rain and produced features analogous to both karst and modern reef forms (Figure 2). He emphasized solution rims, enclosed depressions at various scales, conical karst, tower karst, and karst marginal plains as important natural and experimental features, which could have analogs in both atolls and barrier reef systems. He coupled glacially effected sea-level oscillations, subsidence, and sub-aereal erosion to explain their formation (Figure 3). Atoll morphology derives from the development of solution rims on emerged limestone masses; where rainfall is high, breaches in the karst rim give rise to subsequent atoll-passes, while in the interior, a conical karst may develop as antecedent foundations to lagoonal patch reefs; also, collapse dolines may form as a consequence of extensive subterranean dissolution, the forerunner of “blue holes” according to Purdy (1974). Alternately, where rainfall is low, conical depressions (solution dolines) form within the solution rim. The morphology of atolls and barrier reefs is solution determined rather than growth predicated (Purdy and Winterer, 2001) (Figure 4).

ANTECEDENT PLATFORMS

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Antecedent Platforms, Figure 3 The effects of rainfall on blocks simulated by dripping acid onto limestone surfaces (a) Rainfall (acid) is sufficient to produce a sinle acid menicus over limestone surface; this produces a peripheral rim; (b) Rainfall (acid) insufficient to cover top surface so menisus breaks up into a series of smaller menisci; the result is a rim bounded blockmin which the central depression has been residual solution prominences; and (c) Rainfall (acid) is more than sufficient and runs down the blocks: result is that the block just gets smaller.

solution is directly related to the age of the limestones, that is, depressions in the 120 k reef are only 1 m deep, while in the >480 k reef, they are 12 m deep. Clearly, any karstic solution is cumulative, and Purdy (1974) recognized this, repeating again in 2001 (Purdy and Winterer) that in terms of gross morphology, the dissolution process is cumulative. Most now agree on this. However, an important question is still the extent to which the previous reef (usually the 125,000 years old reef ) has affected the growth pattern of the modern reef, and this will depend on the amount and degree of erosion of the previous reef. Purdy himself was unsure about this, noting the work by Land et al. (1967) again on Bermuda and concluding that a net reduction of the exposed surface by up to about 4.5 m was likely. However, as a general conclusion, this is at odds with the facts, that is, beneath Belize and the Great Barrier Reef, the depth to the 125 k antecedent surface can be anything from 10 to 25 m below present sea level. Assuming the 125 k reef grew to sea level, then the present depth indicates 10–25 m of erosion plus subsidence in the intervening period, clearly at odds with the Bermuda conclusions. It is however in agreement with conclusions in Purdy and Winterer (2001) for gross surface erosion rates on Pacific atolls. In addition, if one uses the vertical erosion rates for coral and coralline algae published by Trudgill (1976) for Aldabra, then vertical erosion of the coral flat and algal flats in the Great Barrier Reef and exposure for around 95,000 years out of the last 125,000 years, then the surface of the 129 k reef is eroded only about 10 m. Coupled with estimated subsidence in the same time frame (5 m or so), the preHolocene surface would be at 10 m prior to the growth of the Holocene reef. This is in fact the depth that it is at on a number of reefs in the Great Barrier Reef (Davies and Hopley, 1983). While this solution/erosion is enough

to produce small-scale features, some of which may be inherited by the modern reef, it is clearly insufficient to have produced the large-scale features on the scale of barrier reef systems or reef tracts. Such features, Purdy says are the result of repetitive cumulative karstification. In the Great Barrier Reef, the idea of a karst marginal plain is no longer in favor (Hopley et al., 2007). Since the seminal 1974 paper, Purdy has become more sure of his ground although even in 1974, he showed the direction in which his thoughts were moving – “Thus the major premise of the subsidence theory (Darwin’s) has been confirmed. It would be dangerous, however, to assume that this subsidence necessarily proves the genetic succession of reef types advocated by Darwin” (Purdy, 1974, p. 10). Further, quoting Vaughan’s warning in 1919 (p. 325) – “although the theoretical possibility of the conversion of a fringing reef into a barrier reef and a barrier reef into an atoll may not be denied, no instance of such a conversion has yet been discovered,” Purdy makes it clear that the evolution of reef types proposed by Darwin is open to question. In 2006, he states that “there are no examples of the subsidence-predicted transition of fringing reefs to barrier reefs to atolls. Moreover, the common occurrence of fringing reefs within barrier reefs negates subsidence as a causal factor in their presumed progressive evolutionary development” (Purdy and Winterer, 2006, p. 143). Instead, Purdy advocates a solution morphology template accentuated by reef construction particularly for barrier reefs.

Growth antecedence This was first proposed by Bloom (1974) and states quite simply that original reef facies exerts a fundamental control

44

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Atolls

Barrier reefs

Rain water, pH generally 40-m depths in clear oceanic waters (Yentsch et al., 2002). Reef development tends to cease at around 4–8% mean surface irradiance (Yentsch et al., 2002; Cooper et al., 2007). Coral cover and species richness decline with increasing depth at high turbidity, with the most phototrophic species being the first to disappear. Importantly, macroalgae are also missing at deeper depths despite the high availability of nutrients, as they are quickly light limited (Hillebrand, 2005). On these dark reef slopes, filter-feeder communities may be prolific. In turbid but shallow water, both light and nutrient levels are high and coral cover and coral growth rates may be very high – but only if herbivores are effective in keeping macroalgae in bay and if hydrodynamics prevent the accumulation of sediments. If herbivores are missing, the slowgrowing corals tend to be outcompeted by fast-growing macroalgae. Fishes play an essential role in the ecology of coral reefs, with low abundances of herbivorous fish potentially leading to a proliferation of macroalgae and declining coral cover (Hawkins and Roberts, 2004). Importantly, the abundance of herbivorous fish on coral reefs may potentially decline not only from overfishing, but also from nutrient pollution. Wolanski et al. (2004) reported a negative relationship between water clarity and the abundance of herbivores on the Great Barrier Reef, and a similar result was found along a water-quality gradient

in the Caribbean (Mallela et al., 2007). On the Great Barrier Reef, herbivorous fish are not usually targeted by fishers suggesting minimal fishing pressure, yet macroalgae cover extensive areas on many inshore reefs (De’ath and Fabricius, 2010), suggesting that control by herbivores is ineffective. The mechanisms for the apparent negative relationship between abundances of herbivorous fishes and poor water quality is still poorly understood; they may include an avoidance of turbidity (Wolanski et al., 2004) or a reduced food palatability, if algae trap large amounts of sediments (Mallela et al., 2007). If such negative relationship was more widespread than presently realized, then water quality would not only directly but also indirectly promote macroalgae, by releasing both nutrient limitation and grazing pressure. A systematic investigation of the potential relationships between water quality and the abundance herbivorous fishes is urgently needed.

Effects of altered sediment properties Coastal nutrient pollution is often linked to altered sediment properties or sedimentation regimes. Coarse sediment grains usually settle within a few kilometers of the discharge source, often contributing to forming river deltas. However, fine particles including detritus, clay, and silt particles may remain suspended for prolonged periods of time. These fine particles carry a large proportion of the total river load of nutrients and pesticides, which eventually settle onto the seafloor and onto the benthic organisms, sometimes tens to hundreds of kilometers away from the source. Previous studies have mostly focused on assessing the effects of increased rates of sedimentation on coral reefs. Sedimentation rate is a strong environmental driver for coral reefs. It is known to reduce coral recruitment rates and coral biodiversity, and many sensitive species are missing or underrepresented in areas of high sedimentation (Rogers, 1990; Fabricius, 2005). Photophysiological stress in corals increases linearly as the product of the duration and amount of sediment exposure, i.e., two units of sediment deposited on the coral for one time unit, exerts a similar decline in photosynthetic yields as half the amount deposited for twice the time (Philipp and Fabricius, 2003). High sedimentation rates (up to 100 mg dry weight cm2) can kill exposed coral tissue within a few days. Exposure to a few days of sedimentation can therefore cause long-term damage to coral populations, by removing whole cohorts of small and sensitive corals. Previous studies have suggested a sedimentation threshold of 10 mg cm2 day1, as reefs may be severely damaged at higher rates (Rogers, 1990). Eutrophication may, however, not always increase the rate of sedimentation, but may alter sediment properties. Experiments have shown that when coral tissue is exposed to sedimentation, the degree of damage increases with bacterial activity (Hodgson, 1990) and with decreasing grain sediment sizes (Weber et al., 2006). Nutrient-rich sediments and sediments with high organic contents cause

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greater stress and damage to corals than do sediments that are poor in organic matter (Weber et al., 2006). Fine and organically enriched sediments may support different bacterial communities and metabolic processes, further altering the microenvironment on the reefs biofilm surfaces. Fine sediments may also bind with marine snow aggregates, and this additional organic material is highly damaging to coral recruits (Fabricius et al., 2003). At very high levels of pollution, the decomposition of organically enriched sediments can lead to hypoxia in poorly flushed and highly stratified locations, further damaging benthos. Not only the amount of sedimentation and its geochemical and physical properties, but also its biological properties and organic contents are ecologically relevant and determine the extent of sedimentation damage in coral reefs. Corals are most sensitive to sedimentation during their recruitment stage, and reduced coral recruitment is one of the most deleterious effects of eutrophication on coral reefs. Coral settlement is sparse on sediment-covered surfaces, and the tolerance of coral recruits to sediment is at least one order of magnitude lower than that of adult corals (Fabricius, 2005). Newly settled corals may be killed by exposure to 12 mg cm2 day1 for 2,000 individual coral reefs between the land and the open ocean. Retention times here are still being debated, but some estimate them to be up to 300 days for dissolved materials (Luick et al., 2007). Particulate materials are likely to be retained for even longer periods of time, as they are repeatedly deposited and resuspended from the shallow sea floor. So, although nutrient enrichment is less severe on the Great Barrier Reef than in many other more densely populated regions, symptoms such as macroalgal dominance and low coral diversity on numerous inshore reefs have been attributed to enhanced terrestrial runoff (van Woesik et al., 1999; Fabricius and De’ath, 2004; Fabricius et al., 2005; De’ath and Fabricius, 2010). Lastly, biological processes are believed to modify the resistance and resilience of coral reefs, but many of these processes are as yet poorly understood. For example, abundant herbivorous fish strongly control macroalgal abundances, thereby promoting resilience (Littler and Littler, 2007). It is also still unresolved to what extent resistance and resilience are codetermined by biodiversity; biogeographic regions with low biodiversity have fewer species to replace the loss of sensitive species and may be more likely to undergo structural and functional changes in their communities (Bellwood et al., 2004). For example in the Caribbean, the loss of the dominant coral species Acropora palmata and Acropora cervicornis and the one remaining important algal grazer, Diadema antillarum, has led to a widespread collapse of reef

ecosystems (Lessios et al., 1984). It is also unknown whether or not the resistance and resilience of reefs vary along latitudinal gradients, as reefs in higher latitudes naturally have lower calcification rates, higher macroalgal biomass, and lower coral biodiversity than do low latitude reefs. Lastly, regions that are prone to severe or frequent disturbances (e.g., from coral bleaching, storms, cold water upwelling, or outbreaks of A. planci) are more likely to be prone to degradation than rarely disturbed regions. This is because poor water quality often does not directly kill the adult coral populations, but retards coral recruitment and hence the speed of recovery from unrelated disturbances. It has also been shown that exposure to one form of stress may decrease the resilience of an ecosystem to another stressor (Hughes et al., 2003; Wooldridge et al., 2005). In summary, degradation from poor water quality is most likely to occur on deeper reef slopes, in locations with weak currents, in places where fish abundances are low, and in regions that are frequently affected by other forms of disturbance. In contrast, well-flushed locations with strong currents, shallow reef crests surrounded by a deep water body, and reefs inhabited by healthy populations of fishes are likely to have the highest levels of resistance and resilience. Several of these factors are identical to those that determine the resistance and resilience of coral to bleaching caused by warming oceans. For example, topography, fast currents, proximity to deep water, and a diverse community with abundant herbivores are considered reliable factors in predicting the likelihood of coral communities dying as a result of bleaching (West and Salm, 2003) – as they are in predicting death arising from eutrophication. One difference is that reefs in shallow waters are – relative to deeper reef slopes – tolerant of the turbidity associated with eutrophication but sensitive to bleaching. A better understanding of the additive or interactive effects between eutrophication and climate change is clearly needed.

Conclusions Although disturbances are a normal and important aspect of their environment, coral reefs are inherently quite stable over time, with coral cover and composition often not changing for many years to decades (Connell, 1997). After severe acute and short-term disturbances, coral in shallow, well-lit windward reefs with fast currents can recover within 10–15 years, if larvae are plentiful (Connell, 1997). Recovery takes longer – possibly up to 50 years – on deeper reef slopes and in poorly flushed settings such as lagoons and areas with weak water flow and poor connectivity. Recovery from chronic and humaninduced disturbances that alter the physical environment is also slower and less commonly observed than recovery from fast and acute disturbances (Connell et al., 1997). If widespread reef degradation is to be avoided, it is essential that the average frequency of severe disturbances does not

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exceed the average recovery time in any one location and that chronic disturbances are minimized. In recent times, the frequency and nature of major disturbances, such as coral bleaching, outbreaks of A. planci, and severe storms, have exceeded the capacity of many reefs around the world to recover (Wilkinson, 2004). Such large-scale events are typically not controllable by management action. In contrast, eutrophication is much more manageable and can often be prevented by preservation of vegetation cover on land, reduction of fertilizer loss into the sea, and restriction of aquaculture facilities to well-flushed locations where dilution is rapid and the resistance and resilience of the marine ecosystems are greatest. With increasing reef disturbances as a result of global warming and seawater acidification, management of water quality and healthy fish abundances will be critically important to the future of coral reef ecosystems.

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Cross-references Adaptation Algae, Coralline Algae-Macro Algae,Turf Bioerosion Bioturbation Climate Change and Coral Reefs Conservation and Marine Protection Areas Corals: Biology, Skeletal Deposition, and Reef-Building Corals: Environmental Controls on Growth Darwin, Charles (1809–1882) El Niño, La Niña, and ENSO FORAM Index Lagoon Circulation Microbes Ocean Acidification, Effects on Calcification Octocorallia Reef Classification by Fairbridge (1950) Reef Classification by Hopley (1982) Reef Classification by Maxwell (1968) Reef Restoration Sediments, Properties Scleractinia, Evolution and Taxonomy Sediment Durability Symbiosis; Zooxanthellae Temperature Change: Bleaching West Indian Coral Reef Classification

O

OCEAN ACIDIFICATION, EFFECTS ON CALCIFICATION Joan A. Kleypas National Center for Atmospheric Research, Boulder, CO, USA

Synonyms Lowered pH of seawater Definition Aragonite: A calcium carbonate (CaCO3) mineral with an orthorhombic crystal lattice structure; a polymorph of calcite. Calcification: The process by which an organism secretes calcium carbonate. Calcite. A calcium carbonate (CaCO3) mineral with a trigonal (rhombohedral) crystal lattice structure; a polymorph of aragonite. High-Mg calcite: Calcite containing at least 4 mol% MgCO3. Ocean acidification: The decrease in the pH of the Earth’s oceans caused by the uptake of carbon dioxide from the atmosphere. Rhodoliths: Coralline algae that grow in rounded, freeliving forms that are unattached to the substrate. Saturation state: A measure of the thermodynamic potential of a mineral to precipitate or dissolve. Introduction Technically, the term “ocean acidification” refers to any process that causes a decrease in seawater pH. Today, however, the term refers almost exclusively to the process by which oceanic water absorbs carbon dioxide from the atmosphere, causing a decrease in ocean pH and changes

in other chemical properties of seawater. Ocean acidification is not a consequence of climate change, but the two share a common cause: increasing concentration of carbon dioxide in the atmosphere. Ocean acidification can affect coral reefs in multiple ways: effects on biogeochemical processes, physiology of organisms, and even such things as sound transmission, have been demonstrated. However the best-documented effects are those on organisms that secrete calcium carbonate, most notably corals, coralline algae, and foraminifera, which are major contributors to reef growth.

Chemistry of ocean acidification The chemical reactions involved in ocean acidification are well documented, albeit somewhat complicated. Carbon dioxide reacts with seawater to form carbonic acid: CO2 þ H2 O ! H2 CO3 :

(1)

Most of the carbonic acid dissociates quickly to hydrogen ions (a contributor to acidity) and bicarbonate ions: H2 CO3 ! HCO3  þ Hþ ;

(2)

and some of the bicarbonate also dissociate to hydrogen ions and carbonate ions: HCO3  ! CO3 2 þ Hþ :

(3)

When carbon dioxide is added to seawater, the relative proportions of the carbonic acid complex (dissolved CO2 and H2CO3), bicarbonate, and carbonate, shift to maintain charge balance in seawater. This comprises the buffering capacity of seawater to maintain pH. For example, the addition of carbon dioxide to seawater causes an increase in bicarbonate ion concentration, a decrease in carbonate ion concentration, and a decrease in pH. The following

David Hopley (ed.), Encyclopedia of Modern Coral Reefs, DOI 10.1007/978-90-481-2639-2, # Springer Science+Business Media B.V. 2011

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further illustrates the reaction of carbon dioxide and water in the presence of calcium carbonate (CaCO3) minerals: CO2 þ H2 O þ CaCO3 ! Ca2þ þ 2HCO3  :

(4)

Calcium carbonate production thus releases CO2 to the water column while its dissolution removes CO2. The calcium carbonate saturation state (O) is a measure of the ion activity product of Ca2þ and CO32 relative to the apparent solubility product (K0 ) for a particular calcium carbonate mineral (e.g., calcite or aragonite). The surface ocean is supersaturated (>1) nearly everywhere (Figure 1), but spontaneous precipitation of carbonate minerals is constrained by kinetic barriers. Calcium is much more abundant than carbonate in seawater, so changes in

O mostly reflect changes in the carbonate ion concentration. The calculated Oaragonite for the tropical oceans, based on a combination of measurements taken in the 1990s (GLODAP, World Ocean Atlas), ranged between about 3.2 and 4.5 (Figure 1a). Most of this variation is due to the 18–30 C temperature range. First order calculations that take into account the progressive increase in atmospheric CO2 concentration (with an associated temperature increase) show that aragonite saturation state will decrease to levels outside the preindustrial range (Figure 1b).

Effects on formation and dissolution of calcium carbonate Calcium carbonate production in the oceans is almost entirely biogenic, and most is precipitated as one of two

Ocean Acidification, Effects on Calcification, Figure 1 (a) Aragonite saturation state of tropical oceans estimated by combining 1  1 gridded fields of total alkalinity and total dissolved inorganic carbon normalized to 1990 conditions (GLODAP data base, Sabine et al., 2005) with 1  1 gridded fields of average annual temperature, salinity, phosphate, and silicate (World Ocean Database, 2005; Boyer et al., 2006). Note that these values are highly interpolated and should not be used for research purposes. (b) Aragonite saturation state as a function of atmospheric CO2 concentration. Lower and upper solid lines are calculated for 18 and 30 C seawater, respectively, with total alkalinity of 2,300 meq kg1, and negligible phosphate and silicate. Parallel dashed lines indicate general range of aragonite saturation state in reef areas prior to the Industrial Revolution. Vertical lines show range of aragonite saturation states at preindustrial atmospheric CO2 concentration (280 ppm), the present decade (390 ppm), and 2 preindustrial (560 ppm).

OCEAN ACIDIFICATION, EFFECTS ON CALCIFICATION

polymorph minerals: calcite, which has a rhombohedral crystal structure; and aragonite, which is orthorhombic. The calcite crystal structure allows some substitution of magnesium ions (Mg) for calcium ions, and calcite with >4 mol% MgCO3 is called “high-Mg calcite.” Normally, calcite is much less soluble in seawater than aragonite, but high-Mg calcite with 12–16 mol% MgCO3 is about 20% more soluble than aragonite (Morse et al., 2006). Many other factors (impurities, structural disorder) can equally affect the solubilities of calcite, high-Mg calcite, and aragonite. The saturation states of the various minerals affect the rates of crystallization and dissolution. The rate of calcium carbonate formation, for example, is often expressed as the equation: R = k(O1)n, where k is the rate constant and n the reaction order. The solubilities of the calcium carbonate minerals have been of interest to Earth scientists for some time, mainly because interpreting Earth’s history requires an understanding of the processes controlling production and preservation of calcium carbonate deposits. Waters of low pH, such as those in the deep sea, have long been known to be corrosive to calcium carbonates, as first noted by Murray and Renard (1891). Marine calcifying organisms typically precipitate either one mineral or the other, and only a few species have the capacity to precipitate two. In general, calcite secretors (e.g., coccolithophorids) occur in higher latitudes where colder waters can hold more CO2 and thus have lower carbonate saturation. Organisms that precipitate the more soluble minerals aragonite (e.g., scleractinian corals) and high-Mg calcite (e.g., crustose coralline algae) are favored in low latitudes, which have higher saturation. Over geologic time, the dominant mineralogy of shallow water carbonates, as well as their preservation potential, has also fluctuated with shifts in seawater carbonate chemistry (Sandberg, 1983). The preferential disappearance of corals and other marine reef-builders in several major extinction events is coincident with rapid ocean acidification (Veron, 2008; Knoll et al., 2007).

Effects on calcification of marine organisms Controlled laboratory experiments have provided most of the evidence that ocean acidification will reduce calcification in corals and coralline algae. Results of these experiments indicate that calcification rates of most tropical coral species will decline significantly as ocean acidification proceeds (reviewed in Kleypas and Langdon, 2006; Langdon and Atkinson, 2005). A wide range of responses has been observed, but on average, experimental results under conditions of 2 preindustrial CO2 concentrations fall within two response classes: (1) >40% reduction in calcification rate and (2) 5 Myr), volcanism occurs or has recently occurred. Introduction Volcanism occurs mainly at or near active plate boundaries, in particular at mid-ocean ridges. However, there are important exceptions where volcanic activity occurs within plate interiors. In the ocean basins, the prime observational evidence comes from the presence of linear chains of islands and seamounts, sometimes extending for thousands of kilometers across the ocean basins. These island and seamount chains exhibit a general progression of decreasing elevation along the chain from volcanic island to fringing reef, to atoll, and finally to a submerged flat-topped seamount (guyot). An active or dormant volcano usually occupies the young end of an island chain, with progressively older and extinct volcanoes occurring along the rest of the chain. Such topographic features were

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suggested to have formed as the plate moved over a relatively stationary “hot spot” in the Earth’s mantle, where an upwelling mantle plume brings hot magma from great depths to the surface (Morgan, 1971; Wilson, 1963). The archetypal example of this process is the HawaiianEmperor seamount chain (Figure 1). The mantle plume that produced this 6,000 km long chain is presently situated beneath the Big Island of Hawaii, the site of active volcanism today. Extending 3,500 km to the northwest lies a linear chain of islands, atolls, and seamounts; there, near the Diakakuji Seamount the chain exhibits a sharp 60 change in orientation and continues another 2,500 km north-northwest to the Aleutian trench. The age of the volcanism gets progressively older with increasing distance from Hawaii, which is seen as a key observation that validates the hot spot hypothesis. Although predominantly an oceanic phenomenon, hot spot volcanism is not restricted to the ocean basins; volcanism of a similar type occurs within continents, one example being the activity of Yellowstone National Park in the western North America.

Origin of oceanic hotspots The hotspot hypothesis states that seamount chains and oceanic islands are the surface manifestation of impinging mantle plumes. These upwelling mantle plumes are thought to originate either at the core-mantle boundary (2,900 km depth) or the boundary between the lower and upper mantle (670 km depth). One theory suggests that a “plume head” develops, above a “plume stem” (e.g., Richards et al., 1989) but other scenarios of plume development with double plume heads (Bercovici and Mahoney, 1994) and time varying magma output (Coffin et al., 2002) have also been considered. As originally proposed (Morgan, 1971; Wilson, 1963), these plumes were considered stationary relative to the lithospheric plates that move over them, but recent evidence suggests that some plumes may have experienced considerable drift over their active lifespan. In particular, deep-sea drilling of seamounts in the northern Emperor seamount chain recovered oriented samples the magnetic memories of which imply that they must have formed at latitudes some 10–15 further north than the present location of the Hawaii hot spot. These systematic discrepancies have lead to suggestions that the mantle plume was further north in the past and consequently has drifted south since the formation of these seamounts (e.g., Tarduno et al., 2009). It is not entirely clear how mantle plumes manage to reach the surface. Apparently weakened by the impinging plume, the lithosphere allows magma to migrate through cracks and fractures and eventually reach the surface. There, a volcano builds upon the surface of the plate directly above the plume. The constant motion of the plate, however, will eventually carry the volcano too far from the source of magma and the volcano becomes extinct. As extinct volcanoes no longer have the

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Oceanic Hotspots, Figure 1 Geometry and ages of the Hawaii-Emperor seamount chain. The major bend in the chain occurred at 47–50 Ma (Sharp and Clague, 2006). Hotspot location (star) is located near Kilauea. Triangles indicate position of samples and their radiometric age is indicated by the color. White arrow shows the approximate Pacific absolute plate motion since 47 Ma.

regenerative volcanism to combat erosion, they subsequently erode as they cool and subside to form fringing reefs and atolls. With time, they eventually sink below the surface of the sea to form a seamount. Wave action will often flatten the top of such seamounts, leading to the classical truncated guyot (Hess, 1946). As long as the mantle plume supplies heat and magma, new active volcanoes will form directly above the mantle plume and the process continues. Mantle plumes may stay active from a few tens to a few hundred million years. The relative motion between plates and plumes causes the formation of lines of extinct volcanoes; these exhibit monotonic age progressions and reflect the history of past plate motions. Since the early formulation of the hotspot hypothesis, numerous hotspots have been proposed for sites of unusual volcanic activity (e.g., Burke and Wilson, 1976; Clouard and Bonneville, 2001; Sleep, 1990) yet conclusive imaging of the underlying mantle plumes using seismic tomography remains elusive (Nataf, 2000), perhaps with Iceland (Wolfe et al., 1997) as an exception. For instance, the archetypal strong plume that many believe has formed the Hawaii-Emperor seamount chain and currently thought to underlie the southeast end of the Big Island of Hawaii is not well resolved, whereas other, less productive hotspots (e.g., Easter, Ascension, Azores) appear more dominant in the tomographic images (Montelli et al., 2004). However, recent three-dimensional seismic imaging beneath the Hawaiian Islands now seem to require an upwelling mantle plume (Wolfe et al., 2009). Although the simple age progressions predicted by the hotspot hypothesis are borne out by observations for several seamount chains (such as the Hawaii-Emperor and Louisville chains), others exhibit a more complex age pattern which casts some doubt on the hotspot theory being the only explanation for such volcanism (e.g., McNutt et al., 1997).

Mantle plumes and hotspot swells Large bathymetric swells may form around and downstream from the site of active hotspots (e.g., McNutt, 1998). While seamount chains are usually confined within a relatively narrow zone (100–200 km), the broader effect of plumes on the lithosphere and upper mantle can extend for 1,000 km or more. Again, the most wellknown case of a hotspot swell is associated with the Hawaiian chain where the seamounts are centrally arranged on top of a broad (1,200 km) swell extending for 3,000 km (e.g., Van Ark and Lin, 2004). Despite this classic example, similar thermal swells are found in all oceans and always associated with hotspot chains. The amplitude and extent of such swells correlate strongly with the age of the underlying lithosphere; hence Hawaii (located in an area with seafloor of around 100 Ma) is more elevated than similar swells found in French Polynesia where the seafloor is only half that age (Cazenave et al., 1988). The earliest studies of hotspot swells concluded that the shallow bathymetry simply reflected reheating of the entire lithosphere (Detrick and Crough, 1978). However, others have pointed out the likely effect of compositional buoyancy (e.g., Jordan, 1979; Phipps Morgan et al., 1995) and dynamical uplift (e.g., Olson, 1990). In summary, hotspot swells appear to require the presence of a deep-seated upwelling plume as such features can explain both a broad swell and narrowly focused volcanism (e.g., Ribe and Christensen, 1999). Intraplate seamounts produced at oceanic hotspots Seamounts formed by hotspot volcanism can accumulate into the largest seamounts present in the oceans. Specifically, intraplate seamounts built on old (and hence thick

OCEANIC HOTSPOTS

and strong) oceanic lithosphere may in some circumstances attain heights of almost 10 km (measured from the seafloor to the tallest peak on an island). This is the case for Mauna Kea, one of five volcanoes that form the Big Island of Hawaii, which by this definition is the tallest mountain on Earth. Given that the smallest features considered by most geologists to be seamounts are 50– 100 m tall, the sizes of observed seamounts span almost three orders of magnitude. Studies have found that the number of seamounts varies considerably across the oceans and that they tend to form both linear and random constellations. Furthermore, their sizes and distributions provide invaluable information about their origins. Ranging from single-beam echosounder profiles, via multibeam surveys, to satellite altimetry, studies have found that the distribution of seamounts can be reasonably well explained by an exponential or power-law model (e.g., Craig and Sandwell, 1988; Smith and Jordan, 1987; Wessel, 2001). Such models support the observation that most seamounts are fairly small. Extrapolations from the power-law trends obtained for large seamounts suggest that perhaps over 100,000 seamounts of heights 1 km may be present in the oceans. Extrapolating further down to the smallest sizes observed (a few tens of meters) would predict a population over one million. However, sediments will most likely have buried the bulk of those seamounts, considering typical 100– 200 m sediment thicknesses in the ocean basins (e.g., Ludwig and Houtz, 1979). Consequently, we find that most of the smallest seamounts reside on young seafloor where the sediment cover is modest or nonexistent. We do not know why seamount abundances vary spatially. One factor may be the underlying distribution of mantle plumes, which seem to be found in higher numbers beneath plates with the largest seamount abundances. Another factor may be systematic variations in plate stresses. We note that smaller plates are possibly in compression, which could prevent the intrusion of magma. Smaller plates are also less likely to have a directional regional stress dominating the state of stress, and often are relatively young and buoyant (e.g., the Cocos plate); during subduction, such plates would only be associated with slightly negative buoyancy forces. On the other hand, the large Pacific plate, and specifically its equatorial region, appears to be under tension from the slab pull forces at the distant subduction zones, as evidenced by widespread extensional volcanism that is neither associated with hotspots nor mid-ocean ridges (e.g., Sandwell et al., 1995; Wessel and Kroenke, 2007). Finally, plates that move the fastest over the underlying mantle appear to have the highest seamount abundances provided they share at least one spreading plate boundary.

Summary Oceanic hotspots are the surface expression of rising mantle plumes from the Earth’s interior and are responsible for much of the intraplate volcanism observed in the ocean

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basins. Being relatively stationary with respect to the surrounding mantle, the oceanic hotspots produce linear chains of islands and seamounts as tectonic plates move over these sites.

Bibliography Bercovici, D., and Mahoney, J., 1994. Double flood basalts and plume head separation at the 660-kilometer discontinuity. Science, 266, 1367–1369. Burke, K. C., and Wilson, J. T., 1976. Hot spots on the Earth’s surface. Journal of Geophysical Research, 93, 7690–7708. Cazenave, A., Dominh, K., Rabinowicz, M., and Ceuleneer, G., 1988. Geoid and depth anomalies over ocean swells and troughs: evidence of an increasing trend of the geoid to depth ratio with age of plate. Journal of Geophysical Research, 93(B7), 8064–8077. Clouard, V., and Bonneville, A., 2001. How many Pacific hotspots are fed by deep-mantle plumes? Geology, 29, 695–698. Coffin, M. F., Prince, M. S., and Duncan, R. A., 2002. Kerguelen hotspot magma output since 130 Ma. Journal of Petrol, 43, 1121–1139. Craig, C. H., and Sandwell, D. T., 1988. Global distribution of seamounts from Seasat profiles. Journal of Geophysical Research, 93(B9), 10408–10420. Detrick, R. S., and Crough, S. T., 1978. Island subsidence, hot spots, and lithospheric thinning. Journal of Geophysical Research, 83, 1236–1244. Hess, H. H., 1946. Drowned ancient islands of the Pacific Basin. American Journal of Science, 244, 772–791. Jordan, T. H., 1979. Mineralogies, densities, and seismic velocities of garnet lherzolites and their geophysical implications. In Boyd F. R., and Meyer H. O. A. (eds.), The Mantle Sample: Inclusions in Kimberlites and Other Volcanics. Washington, DC: American Geophysical Union, pp. 1–14. Ludwig, W. J., and Houtz, R. E., 1979. Isopach Map of the Sediments in the Pacific Ocean Basin, Color Map with Text. Tulsa, OK: American Association of Petroleum Geologists. McNutt, M. K., 1998. Superswells. Reviews of Geophysics, 36(2), 211–244. McNutt, M. K., Caress, D. W., Reynolds, J., Jordahl, K. A., and Duncan, R. A., 1997. Failure of plume theory to explain midplate volcanism in the southern Austral Islands. Nature, 389(6650), 479–482. Montelli, R., et al., 2004. Finite-frequency tomography reveals a variety of plumes in the mantle. Science, 303(5656), 338–343. Morgan, W. J., 1971. Convection plumes in the lower mantle. Nature, 230, 43–44. Nataf, H. -C., 2000. Seismic imaging of mantle plumes. Annual Review of Earth and Planetary Sciences, 28, 391–417. Olson, P., 1990. Hot spots, swells and mantle plumes. In Ryan M. P. (ed.), Magma Transport and Storage. New York: John Wiley, pp. 33–51. Phipps Morgan, J., Morgan, W. J., and Price, E., 1995. Hotspot melting generates both hotspot volcanism and a hotspot swell? Journal of Geophysical Research, 100(B5), 8045–8062. Ribe, N. M., and Christensen, U. R., 1999. The dynamical origin of Hawaiian volcanism. Earth and Planetary Science Letters, 171, 517–531. Richards, M. A., Duncan, R. A., and Courtillot, V., 1989. Flood basalts and hot spot tracks: plume heads and tails. Science, 246, 103–107. Sandwell, D. T., et al., 1995. Evidence for diffuse extension of the Pacific plate from Pukapuka ridges and cross-grain gravity lineations. Journal of Geophysical Research, 100(B8), 15087– 15099.

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Sharp, W. D., and Clague, D. A., 2006. 50-Ma initiation of HawaiiEmperor bend records major change in Pacific plate motion. Science, 313, 1281–1284. Sleep, N. H., 1990. Hotspots and mantle plumes: some phenomenology. Journal of Geophysical Research, 95(B5), 6715–6736. Smith, D. K., and Jordan, T. H., 1987. The size distribution of Pacific seamounts. Geophysical Research Letters, 14, 1119–1122. Tarduno, J. A., Bunge, H. -P., Sleep, N. H., and Hansen, U., 2009. The bent Hawaiian-Emperor hotspot track: inheriting the mantle wind. Science, 324, 50–53. Van Ark, E., and Lin, J., 2004. Time variation in igneous volume flux of the Hawaii-Emperor hot spot seamount chain. Journal of Geophysical Research, 109(B11401), doi:10.1029/ 2003JB002949. Wessel, P., 2001. Global distribution of seamounts inferred from gridded Geosat/ERS-1 altimetry. Journal of Geophysical Research, 106(B9), 19431–19441. Wessel, P., and Kroenke, L. W., 2007. Reconciling late neogene Pacific absolute and relative plate motion changes. Geochemistry Geophysics and Geosystems, 8(Q08001), doi:10.1029/ 2007GC001636. Wilson, J. T., 1963. A possible origin of the Hawaiian islands. Canadian Journal of Physics, 41, 863–870. Wolfe, C. J., Bjarnason, I. T., VanDecar, J. C., and Solomon, S. C., 1997. Seismic structure of the Iceland mantle plume. Nature, 385(6613), 245–247. Wolfe, C.J. et al., 2009. Mantle shear-wave velocity structure beneath the hawaiian hot spot. Science, 326, 1388–1390.

Cross-references Atolls Plate Tectonics Subsidence Hypothesis of Reef Development Volcanic Disturbances and Coral Reefs Volcanic Loading and Isostasy

OCTOCORALLIA Katharina Fabricius Australian Institute of Marine Science, Townsville, QLD, Australia

Synonyms Alcyonaria (this term is now rarely used). Common names: Gorgonians; Octocorals; Sea fans; Sea pens; Sea whips; Soft corals Definition Octocorallia (also known as octocorals, or in earlier times “Alcyonaria”) are a subclass of the class Anthozoa, in the phylum Cnidaria. They are sessile polyp-bearing animals with a mobile larval phase that are only found in marine systems. The distinguishing characteristic of this subclass is that their polyps always bear eight tentacles (hence octocoral), which are usually (but not always) fringed along both edges by one or more rows of pinnules. The gastrovascular cavities are typically subdivided by eight septa.

Introduction Octocorals are a major component of the sessile benthic fauna of many coral reefs. Like Scleractinian corals, they are modular (colonial) organisms which develop through asexual replication and specialization of polyps. In contrast to Scleractinian corals (the primary reef building corals), most octocorals do not usually deposit a rigid calcium carbonate exoskeleton, and they therefore tend to attach to reefs rather than contribute substantially to their framework or to sedimentary deposits. However, there are a number of exceptions that do contribute substantially to framework accretion, sedimentary deposits, and sediment stabilization (see below). Classification Octocorallia are a subclass of the class Anthozoa, phylum Cnidaria. Like other anthozoans, octocorals bear polyps that consist of a tubular body, terminating in a mouth that is surrounded by hollow tentacles. Unlike other anthozoans, octocoral polyps always bear eight tentacles, which are usually fringed by rows of pinnules. Polyps are internally compartmentalized by eight mesenteries. The ectoderm of the tentacles contains simple nematocysts (cnidae or stinging cells) that are unable to paralyze large zooplankton or sting other larger animals. The ectodermal and endodermal cell layers are often connected by a relatively thick and cellular mesoglea, which often contains calcium carbonate sclerites and collagen. Most octocorals are colonial, with the exception of one deepwater species with a single solitary polyp (Taiaroa tauhou). A colony develops from a single founder polyp by asexual propagation (polyp budding). The resulting multiple polyps in a colony are often embedded into coenenchymal tissue, but remain connected by canals (solenia). There is substantial taxonomic uncertainty associated with Octocorallia. At present, the subclass is divided into three orders (Figure 1):  Helioporacea  Pennatulacea (sea pens)  Alcyonacea (soft corals and gorgonians) The Helioporacea are represented by a single species, the blue coral, Heliopora coerulea (family Helioporidae) in the Indo-Pacific, and the genus Epiphaxum (family Lithotelestidae) in the Caribbean. H. coerulea has no calcium carbonate sclerites, instead colonies consist of a massive aragonite skeleton perforated by wide cylindrical cavities containing the polyps, connected by narrow solenial tubes. The Pennatulacea (sea pens) presently distinguish around 200 species in 32 genera (15 families; Williams, 1995, 1999). The body of pennatulaceans consists of a single large primary polyp, called the oozooid, with a basal fleshy peduncle for anchorage in soft substratum. Through lateral budding of the upper body wall of the oozooid (rachis), two types of secondary polyps are being

OCTOCORALLIA

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Octocorallia, Figure 1 Representatives of the three orders of Octocorallia: (a) Helioporacea; (b) Pennatulacea; (c and d) two common morphological groups of Alcyonacea, soft corals, and gorgonians.

formed, autozooids and siphonozooids, with mesozooids as an additional third type of polyps in a few taxa. Pennatulaceans typically live in soft bottom habitats, being completely retracted during the day and emerging in the dark. The greatest pennatulacean diversity is found in the tropical Indo-Pacific (nine genera in five families). The Alcyonacea (soft corals and gorgonians) contain the large majority of octocoral species. The number of species is unknown, as numerous species await taxonomic

description and many genera need urgent revision. In shallow tropical and subtropical Indo-Pacific reefs alone, about 100 alcyonacean genera in 23 families are currently described (Fabricius and Alderslade, 2001). Their growth forms range from small colonies with few polyps connected by stolons, to fleshy soft corals to up to 3 m large sea fans (Figure 2). Within the order Alcyonacea, the term “soft coral” is commonly used to only refer to octocorals without internal axis or solid skeleton. In

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Octocorallia, Figure 2 Representative morphological groups of alcyonacean octocorals: (a) single polyps connected by stolons; (b) tall axial polyps that bud off lateral polyps; (c) polyps embedded in a fleshy tissue mass (“soft coral”); (d) gorgonian sea fan with upright growth form and internal axis (modified from Hyman, 1940).

Octocorallia, Figure 3 A representative set of octocoral sclerites of the genus Acanthogorgia. Sclerites are the main feature used for taxonomic analyses (drawing by P. Alderslade, from Fabricius and Alderslade, 2001).

contrast, the terms gorgonian, sea fan, sea rod, or sea whip are generally used to refer to octocorals that arise from the substratum with the support of an internal axis (but excluding sea pens and blue coral). However, in the Caribbean where gorgonians dominate the octocoral fauna, the term soft coral is often used for all octocorals. While the terms soft coral and gorgonian continue to be used to differentiate between morphological groups (and continues to be used in some classification schemata), continua of intermediate forms exist, and earlier taxonomic separations of the Alcyonacea into different groups have been abolished by most classifications (Bayer, 1981). Phylogenetic research suggests that further significant changes to the current classification may be warranted.

Most octocoral colonies are supported by sclerites embedded in the tissue, and a hydroskeleton or a proteinaceous and/or calcareous axis. Sclerites are minute calcareous structures typically less than 0.3 mm in length (range about 0.02 to >12 mm; Figure 3). Some species have neither sclerites nor other calcareous structures (e.g., some Xeniidae). A few species (e.g., H. coerulea, Tubipora musica, and Corallium) have a solid skeleton. The rigid structure of these species contributes to habitat complexity on the reef during their life, and they may also contribute to reef framework and sedimentary deposits on their death. Some species of Sinularia also contribute to reef structure and sediment stabilization by depositing massive trunks of consolidated sclerites that can reach decimeters across and meters in height.

OCTOCORALLIA

Among the gorgonians, a wide variety of additional supporting structures are found. Their internal axes may consist of fused or unfused sclerites, sometimes bound by amorphous calcareous material, solid calcium carbonate, and/or a combination of collagen and gorgonin (a dark, hard proteinaceous material related to horn). For example, sections along the axis may alternate between horny and calcium carbonate materials (family Isididae), nodes and internodes (family Melithaeidae), or layers across the axes may consist of a gorgonin core surrounded by large amounts of calcareous material (family Ellisellidae). The shape and size of sclerites, the structure and material composition of the internal axes, and colony growth forms are the main features used for taxonomic analyses (Fabricius and Alderslade, 2001). Sclerites may be investigated by placing a small amount of tissue (1–3 mm3) onto a microscopic slide, and digesting the tissue with a few drops of concentrated bleach (sodium hypochlorite). Once the tissue is digested and bubble formation has ceased after a few minutes, a few drops of a glycerol and cover slip are applied for casual microscopic examination. To produce permanent slides, the sclerites are rinsed in freshwater and then in 95% ethanol. When all liquid is completely evaporated and the sclerites dry, an acid-free mounting medium and a cover slip are applied.

Evolution and biogeography Due to a lack of preservable features, very little is known about the evolution of octocorals. The Pennatulacea are an exception, with phylogenetic analyses presented by Williams (1993a, 1995). Williams postulates that the Pennatulacea as a group initially differentiated from shallow-water alcyoniid-like ancestors in the shallow waters of tropical oceans, and subsequently diversified and dispersed to all depths of the temperate and polar regions, as well as the tropics. The most primitive pennatulacean group Veretillidae show a range of similarities with some alcyonacean taxa, and like the alcyonaceans are most diverse in the Indo-Pacific at 100 mm), and can have a patchy (e.g., shelter and intraparticle) to uniform (e.g., interparticle) distribution (Lǿnǿy, 2006). Where affected by karst dissolution, pore types include moldic macropores (>30–30 mm), vugs, and solution-enlarged fractures. Porosity values range from low (30%), as does permeability (50% of global reserves), and reef reservoirs are a significant proportion of these reservoirs. They are particularly abundant in the Siluro-Devonian, Cretaceous, and Neogene time periods. The Neogene-aged reservoirs are commonly coral reef and associated

750

OIL AND GAS RESERVOIRS AND CORAL REEFS

Oil and Gas Reservoirs and Coral Reefs, Figure 5 (a) Reflection and (b) acoustic impedance section through Jintan field. Synthetic seismic logs show good match with seismic. Horizons (H1B, H2A, etc.) define reservoir zones shown on acoustic impedance (Z1a, Z2a, etc.). Low impedance areas (blue/green) (i.e., higher porosity) show both vertical and lateral connectivity (from Vahrenkamp et al., 2004).

grainsupportstone deposits. Reef reservoir pore systems are generally characterized by some combination of primary depositional, and diagenetic pore types. Porosity ranges from patchy to uniform, resulting in vertical and horizontal heterogeneity. Climate controls the effects of early diagenesis. Greenhouse reservoirs show less effect of subaerial dissolution and greater internal continuity. Icehouse reservoirs that have been affected by humid karst conditions show significant pore system modification. The Neogene coral/algal

reservoirs are most similar to modern reef systems and exhibit dissolution modification of their pore systems.

Bibliography Abdullah, M., and Jordan, C. F., 1987. The geology of Arun field, Miocene reef complex. Proceedings of the Indonesian Petroleum Association, 16th Annual Conference, pp. 65–96. Beach, D. K., 1982. Depositional and diagenetic history of Pliocene-Pleistocene carbonates of northwestern Great Bahama

OIL AND GAS RESERVOIRS AND CORAL REEFS

Bank; evolution of a carbonate platform. PhD dissertation, University of Miami, p. 447. Choquette, P. W., and Pray, L. C., 1970. Geological nomenclature and classification of porosity in sedimentary carbonates. AAPG Bulletin, 54, 207–250. Ehrlich, R. N., Longo, A. P. Jr., and Hyare, S., 1993. Response of carbonate platform margins to drowning: evidence of environmental collapse. In Loucks, R. G., and Sarg, J. F. (eds.), Carbonate Sequence Stratigraphy. AAPG Memoir, 57, 241–266. Epting, M., 1980. Sedimentology of Miocene carbonate build-ups, Central Luconia, offshore Sarawak. Geological Society of Malaysia Bulletin, 12, 17–30. Epting, M., 1989, Miocene carbonate build-ups of central Luconia, offshore Sarawak. In Bally, A. W. (ed.), Atlas of Seismic Stratigraphy. AAPG Studies in Geology, 27, 168–173. Eyles, D. R., and May, J. A., 1982. Exploration of the L-Structure, Natuna D-Alpha Block, Offshore Indonesia: CCOP-ASCOP seminar on hydrocarbon occurrence in carbonate rocks. Surabaya, Indonesia, 15p. Fulthorpe, C. S., and Schlanger, S. O., 1989. Paleo-oceanographic and tectonic settings of early Miocene reefs and associated carbonates of offshore southeast Asia. AAPG Bulletin, 73, 729–756. Goldhammer, R. K., Dunn, P. A., and Hardie, L. A., 1990. Depositional cycles, composite sea level changes, cycle stacking patterns, and the hierarchy of stratigraphic forcing: examples from Alpine Triassic carbonates. Geological Society of America Bulletin, 102, 535–562. Grötsch, J., and Mercadier, C., 1999. Integrated 3-D reservoir modeling based on 3-D seismic: the Tertiary Malampaya and Camago buildups, offshore Palawan, Philippines. AAPG Bulletin, 83, 1703–1728. Gucci, M. A., and Clark, M. H., 1993. Sequence stratigraphy of a Miocene carbonate buildup, Java Sea. In Loucks, R. G., and Sarg, J. F. (eds.), Carbonate Sequence Stratigraphy. AAPG Memoir, 57, 291–304. Ho, K. F., 1978. Stratigraphic framework for oil exploration in Sarawak. Geological Society of Malaysia Bulletin, 10, 1–13. James, N. P., 1983. Reef environment. In Scholle, P. A., Bebout, D. G., and Moore, C. H. (eds.), Carbonate Depositional Environments. AAPG Memoir, 33, 345–440. Jordan, C. F., and Abdullah, M., 1992. The Arun field – Indonesia, North Sumatra basin, Sumatra. In Beaumont, A. F., and Foster, N. H. (comps.), Stratigraphic Traps III. AAPG treatise of Petroleum Geology, Atlas of Oil and Gas Fields, pp. 1–39. Koerschner, W. F., and Read, J. F., 1989. Field and modeling studies of Cambrian carbonate cycles, Virginia Appalachians. Journal of Sedimentary Petrology, 59, 654–687. Longman, M. W., Maxwell, R. J., Mason, A. D. M., and Beddoes, L. R., 1987. Characteristics of a Miocene intrabank channel in Batu Raja limestone, Ramba field, south Sumatra, Indonesia. AAPG Bulletin, 71, 1261–1273. Lǿnǿy, A., 2006. Making sense of carbonate pore systems. AAPG Bulletin, 90, 1381–1405. Loucks, R. G., 1999. Paleocave carbonate reservoirs: origins, burial-depth modifications, spatial complexity and reservoir implications. AAPG Bulletin, 83, 1795–1834. Lucia, F. J., 1999. Carbonate Reservoir Characterization. Berlin: Springer, p. 226. Maliki, M. A., and Soenarawi, S., 1991. South Lho Sukon-D1 discovery, north Sumatra. Proceedings of the Indonesian Petroleum Association, 20th Annual Convention, pp. 235–254. May, J. A., and Eyles, D. R., 1985. Well log and seismic character of Tertiary Terumbu carbonate, South China Sea, Indonesia. AAPG Bulletin, 69, 1339–1358. Melim, L. A., Swart, P. K., and Maliva, R. G., 2001. Meteoric and marine-burial diagenesis in the subsurface of Great Bahama Bank. In Ginsburg, R. N. (ed.), Subsurface Geology of

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a Prograding Carbonate Platform Margin, Great Bahama Bank: Results of the Bahamas Drilling Project. SEPM special publication number, 70, pp. 137–161. Moldovanje, E. P., Tanner, H. C., and Zhang, J. Y., 1995. Regional exposure events and platform evolution of Zhujiang Formation carbonates, Pearl River Mouth basin: evidence from primary and diagenetic seismic facies. In Budd, D. A., Saller, A. H., and Harris, P. M. (eds.), Unconformities and porosity in carbonate strata. AAPG Memoir, 63, 125–140. Perkins, R. D., 1977. Depositional framework of Pleistocene rocks in South Florida. Geological Society of America Memoirs, 147, 131–198. Read, J. F., 1995, Part 1. Overview of carbonate platform sequences, cycle stratigraphy and reservoirs in greenhouse and ice-house worlds. In Read, J. F. et al. (eds.), Milankovitch Sea-Level Changes, Cycles, and Reservoirs on Carbonate Platforms in Greenhouse and Ice-House Worlds. Society for Sedimentary Geology (SEPM) Short Course Number, 35, pp. 1–102. Roehl, P. O., and Choquette, P. W., 1985. Introduction. In Roehl, P. O., and Choquette, P. W. (eds.), Carbonate Petroleum Reservoirs. New York: Springer, pp. 1–15. Rudolph, K. W., and Lehmann, P. J., 1989. Platform evolution and sequence stratigraphy of the Natuna Platform, South China. In Crevello, P. D., Wilson, J. L., Sarg, J. F., and Read, J. F. (eds.), Controls on Carbonate Platform and Basin Development. SEPM special publication number, 44, pp. 353–361. Saller, A., Armin, R., Ichram, L. O., and Glenn-Sullivan, C., 1993. Sequence stratigraphy of aggrading and backstepping carbonate shelves, Oligocene, central Kalimantan, Indonesia. In Loucks, R. G., and Sarg, J. F. (eds.), Carbonate Sequence Stratigraphy. AAPG Memoir, 57, 267–290. Sulaiman, M., 1995. Cyclic Carbonate Deposition, Facies Succession, and Diagenesis of a Central Luconia buildup, Offshore Sarawak. MS thesis, University of Brunei Darussalam, 73 p. Sun, S. Q., and Esteban, M., 1994. Paleoclimatic controls on sedimentation, diagenesis, and reservoir quality: lessons from Miocene carbonates. AAPG Bulletin, 78, 519–543. Trice, R., 2005. Challenges and insights in optimizing oil production from Middle Eastern karst reservoirs. SPE Middle East Oil and Gas Show and Conference, Bahrain, SPE Paper 93679, 26 p. Vahrenkamp, V. C., David, F., Duijndam, P., Newall, M., and Crevello, P., 2004. Growth architecture, faulting, and karstification of a middle Miocene carbonate platform, Luconia Province, offshore Sarawak, Malaysia. In Seismic Imaging Of Carbonate Reservoirs and Systems. AAPG Memoir, 81, 329–350. Withjack, E. M., 1985. Analysis of naturally fractured reservoirs with bottom water drive: Nido A and B fields, offshore northwest Palawan, Philippines. Journal of Petroleum Technology, SPE Paper, 12019, 1481–1490. Wright, V. P., 1992. Speculations on the controls on cyclic peritidal carbonates: ice-house versus greenhouse eustatic controls. Sedimentary Geology, 76, 1–5. Yaman, F., Ambismir, T., and Bukhari, T., 1991. Gas exploration in Parigi and pre-Perigi carbonate buildups, NW Java Sea. Proceedings Indonesian Petroleum Association, 20th Annual Convention, IPA 91–11.20, 319–346.

Cross-references Diagenesis Density and Porosity: Influence on Reef Accretion Rates Porosity Variability In Limestone Sequences Reef Drilling Sea Level Change and Its Effect on Reef Growth Seismic Reflection Solution Processes/Reef Erosion Submarine Lithification

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OOIDS

OOIDS Maurice Tucker Durham University, Durham, UK

Synonyms Ooliths Definition Ooids are spherical-subspherical grains, consisting of one or more regular concentric lamellae around a nucleus, usually a carbonate particle or quartz grain. Sediment composed of ooids is referred to as an oolite. The term ooid has been restricted to grains less than 2 mm in diameter and the term pisoid (formerly pisolith) is used for similar grains of a larger diameter. Coated grain is a general term frequently used for ooids and pisoids, and includes oncoids, grains with a microbial coating. Ooids typically form in agitated shallow waters of subtropical seas where they are frequently moved as sandwaves, dunes and ripples by tidal and storm currents, and wave action. They are commonly found in association with reefs and reefal limestones; thus they often form along shelf-margins (e.g., the Bahamas), but they also occur in tidal-deltas and beach-barriers of shallow ramps (e.g., Trucial Coast, Arabian Gulf ). The majority of modern ooids range from 0.2 to 0.5 mm in diameter, are composed of aragonite and have a high surface polish. The characteristic microstructure is a tangential arrangement of acicular crystals, 2 mm in length. Rarely ooids (and pisoids) form in quieter-water marine locations, such as lagoons and tidal flats, where a radial fabric is more common. Ooids can form in areas of strong wave action in lakes, where they usually have a dull, commonly, cerebroid (bumpy) surface. They may be composed of aragonite or calcite, or be bi-mineralic (low-Mg calcite–aragonite). These may be associated with stromatolite bioherms. Ancient marine ooids Ooids in the rock record are generally composed of calcite (low Mg), but of these some were originally calcite (possibly high Mg), whereas others were originally aragonite. Primary calcite ooids, whether in high-energy or lowenergy facies, typically have a radial texture of wedgeshaped, fibrous crystals, with an extinction cross under

crossed polars. Ancient ooids originally of aragonite will have been altered during diagenesis to a greater or lesser extent. Commonly, the aragonite has been dissolved out completely, to leave oomoulds, or these holes may be filled with calcite cement. Some ooids have a fine-grained micritic texture, mostly the result of micritisation by endolithic microbial organisms.

Origin of ooids Current ideas for the origin of ooids invoke inorganic or biochemical processes. For the former, the sea water in shallow tropical areas is supersaturated with respect to CaCO3, so that this, together with water agitation, CO2 degassing and elevated temperature, is considered sufficient to bring about carbonate precipitation on nuclei. In a biochemical origin, the microbial processes in the organic mucilage (EPS) that coats and permeates the ooids create a microenvironment conducive to carbonate precipitation. The factors determining the primary mineralogy of ooids are water chemistry, especially PCO2, Mg/Ca ratio, carbonate saturation, and possibly the degree of water agitation. It is believed that aragonite and high-Mg calcite ooids are precipitated when PCO2 is low and Mg/Ca ratio high, with opposite conditions for low-Mg calcite ooids. High carbonate supply, as would occur in high-energy locations, is thought to favour aragonite precipitation over high-Mg calcite. There is a secular variation in the original mineralogy of ooids through the Phanerozoic, with aragonitic ooids, which may be associated with calcitic ooids (likely highMg calcite originally) in the late Precambrian/early Cambrian, mid-Carboniferous through Triassic and Tertiary to the Recent, and calcitic ooids (presumed to have low-to-moderate Mg content) dominant in the midPalaeozoic and the Jurassic–Cretaceous. This pattern, which correlates with the first-order, global sea-level curve, suggests a geotectonic mechanism(s) causing subtle variations in PCO2 and/or Mg/Ca ratio in seawater. High sea-level stands, correlating with calcite seas, are times of high rates of sea-floor spreading, and high PCO2 from increased metamorphism at subduction zones, and low Mg/Ca ratio from increased extraction of Mg2þ at midocean ridges through seawater pumping. Bibliography Tucker, M. E., and Wright, V. P., 1990. Carbonate Sedimentology. Oxford: Blackwell Science.

P

PACIFIC CORAL REEFS: AN INTRODUCTION James E. Maragos1, Gareth J. Williams2 1 U.S. Fish and Wildlife Service, Honolulu, HI, USA 2 Victoria University of Wellington, Wellington, New Zealand

Introduction This entry introduces coral reefs and associated archipelagos within or facing the Pacific Ocean, including those not sheltered behind adjacent western seas (Japan, Yellow, South China, Sulu, Celebes, Banda, Arafura); continents (Australia, Asia), and sub-continents (SE Asia, New Guinea). The Great Barrier Reef of Eastern Australian is within the Pacific Ocean but, like other regions, is covered in more detail in separate entries. A total of 47 archipelagos and isolated reefs and islands under the jurisdiction of 31 nations and territories in the Pacific are covered here and listed in Table 1. The location and current names of most of the archipelagos and some of the nations are shown in Figure 1. It is impossible to list and show all the governments and archipelagoes on one map or accurately estimate the number of tropical islands and reefs in the Pacific. However, there are likely more than 10,000 islands and associated reefs, far more than reported in Table 1, which focuses mostly on the larger inhabited Pacific islands. The Pacific supports the largest tropical and subtropical habitat for coral reefs in the world at about 58% of the global total, amounting to approximately 88 million km2 between latitudes of 30
N to 30
S, based upon distance and area calculations of Google Earth imagery. This compares to about 35 million km2 of habitat for the Indian Ocean, about 17 million km2 of habitat for the Gulf of Mexico-Caribbean Sea, and 13 million km2 of habitat for SE Asia. Along the Equator, the Pacific stretches from

127
E (off the eastern coast of Haimahera Island, north of New Guinea Island) to 80
W (off the western coast of Ecuador near Quito), a distance of 16,900 km and accounting for 42% of the circumference of the earth. Although the Pacific’s huge size supports much larger numbers of coral reefs, tropical islands, large reefs, and habitat compared to those in the Indian Ocean and Gulf of Mexico-Caribbean Sea, the Pacific island number and size totals are less than those within SE Asia; much of the tropical and subtropical habitat in the Pacific is devoid of islands and reefs.

Categories of reefs in the Pacific Ocean All coral reefs begin as coral and/or coralline algal communities either along the sides and tops of emergent volcanoes near or above the sea surface within the Pacific Ocean basin or on preexisting substrates of the continents or continental islands bordering the Pacific. Periods of subsequent subsidence or uplift, tectonic plate movements, sea-level rises and falls, and other epochs of unfavorable ocean temperatures and chemistry during the Mesozoic and Cenozoic likely led to massive drowning, dissolution, and subduction of many ancient reefs. Not all of these were restored following the return of favorable conditions for their growth and survival. This chapter covers only living Pacific reefs comprising: 1. Fringing reefs – coral reefs that grow out and upwards to the sea surface directly from the shoreline of mostly volcanic or continental islands (Figure 2), starting out as aprons that may eventually encircle the fringes of islands and form shallow reef crests or flats. 2. Barrier reefs – linear offshore reefs that encircle subsiding or emerging volcanic islands offshore or on continental shelves beyond shallow lagoons that separate the barrier reefs from the island or continent (Figure 3).

David Hopley (ed.), Encyclopedia of Modern Coral Reefs, DOI 10.1007/978-90-481-2639-2, # Springer Science+Business Media B.V. 2011

Jurisdiction

Papua New Guinea France FSM & Palau France Costa Rica New Zealand Australia Papua New Guinea Fiji Ecuador France Mexico Kiribati Tonga United States Indonesia New Zealand Kiribati & US Papua New Guinea United States

Archipelago

Admiralty – Bismarck Austral Caroline Clipperton Is. Cocos & Puerto Quepos Cooks Coral Sea & Middleton D’Entrecasteaux. Fiji – Lau Galapagos Gambiers Guadalupe – Socorro Gilbert Ha’apai Group Hawaii Indonesia – West Pacific Kermadecs Lines Louisiades Marianas 462,841 3,265 1,193 9 25 241 354 3,100 18,274 7,880 31 406 810 see Tonga 16,629 nd 33 515 1,790 1018

Land – km2 3,120,000 5,030,000 3,608,193 330,000 nd 1,830,000 nd see PNG 1,290,000 45,000 see Austral nd 3,550,000 see Tonga 2,147,985 nd nd nd 26,000 2,041,000

Ocean – km2 0 12 4 0 0 1 2 0 0 0 0 0 0 2 32 0 0 2 0 1

Bank 22 3 37 0 0 3 1 1 22 0 1 0 12 4 10 7 0 6 18 3

Lag 0 0 3 0 0 0 8 50 33 0 2 0 0 0 1 0 0 1 0 0

Areef

1 2 5 0 0 2 0 1 22 0 0 0 0 2 2 nd 0 0 12 2

Barr

24 12 13 0 0 2 5 1 15 0 0 0 6 6 1 9 0 4 60 2

Low

26 1 34 0 0 7 4 7 9 0 1 0 10 0 5 7 0 5 17 0

Atol

9 2 523 1 0 4 1 0 10 0 0 1 2 15 2 0 0 2 2 5

Lime

18 14 38 0 12 6 2 5 810 125 13 20 0 41 24 nd 15 0 21 12

Volc

60 31 608 1 12 19 12 13 844 125 14 21 18 62 32 16 15 11 100 19

All is

1 0 1 1 1 0 0 1 1 1 0 1 0 1 0 1 1 0 1 1

Tren

1 0 0 0 1 0 1 1 1 1 0 1 0 1 0 1 1 0 1 1

Cont

0 1 0 1 0 1 0 0 0 1 1 1 1 0 1 0 0 1 0 1

Bas

0 1 1 0 0 1 0 0 0 0 1 1 0 0 1 0 0 0 0 0

Hot

Pacific Coral Reefs: An Introduction, Table 1 Comparing 47 coral reef archipelagos and single islands in the Pacific Ocean. Explanations: bank = submerged reef, lag = lagoon or marine lake, areef = atoll reef, barr = barrier reef, low = low reef island, atol = atoll, lime = raised limestone island or atoll, and closed atolls, volc = volcanic island with fringing reef, all is = total of all main islands, tren = tectonic trench nearby and likely contributing to island’s formation, cont = continent or continental island, bas = basalt or oceanic island, hot = island or reef near to and likely formed over an oceanic hotspot, 1 (or more) = present, 0 = absent, and nd = no data. Area units = km2. Jurisdictions listed are at the highest independent national levels. Jurisdiction and archipelago names often differ, and archipelago names in brackets are those used during the twentieth century before political independence. Total number of islands and reefs are underestimated and focus on main inhabited islands, especially in the Solomon, Papua New Guinea, New Hebrides, Vanuatu, and Fiji Islands. (Sources: Bryan (1953), UNEP (1988), Dahl (1991), Maragos and Holthus (1998) National Geographic Society maps, Google Earth (2009), Reef Base, and Wikipedia web notes)

754 PACIFIC CORAL REEFS: AN INTRODUCTION

France Marshall Is. & US Japan Nauru France New Zealand Japan Panama Kiribati – US United Kingdom Chile Chile Japan Samoa & US France Solomon Is. Papua New Guinea Peoples Rep. China Republic of China NZ & US Tonga Australia & PNG France Tuvalu Tonga Vanuatu France

47

Marquesas Marshalls Minami Torishimas Nauru New Caledonia Nuie I Ogasawara (Bonin) Perlas & Chiriqui Phoenix Is Pitcairn & Henderson Rapa Nui Sala y Gomez Ryukyus Samoa Society Solomon Solomon Sea South China Sea Taiwan Tokelaus Tongatapu Group Torres Strait Tuamotu Tuvalu (Ellice) Vava’u Group Vanuatu (New Hebrides) Wallis & Futuna

Totals

625,451

1,049 181 1.2 21 19,060 259 73 394 31.4 43 166 0.15 nd 3240 1,680 28,450 see PNG nd 35,980 11 718 nd 885 26 see Tonga 12,190 278 35,299,966

see Austral 2,131,000 nd 320,000 1,740,000 390,000 2,300 nd 413,788 800,000 320,000 320,000 nd 510,000 see Austral 1,340,000 see PNG 729,000 nd 290,000 nd 48,000 see Austral 900,000 see Tonga 710,000 300,000 75

0 0 0 0 4 0 0 0 2 0 0 0 0 10 0 1 1 nd 0 0 0 nd 0 1 0 0 0 323

0 30 0 1 14 0 0 0 3 2 0 0 0 3 13 13 6 nd 0 3 3 2 72 6 0 1 1 148

0 0 0 0 38 0 1 0 0 0 0 0 0 0 0 9 0 1 0 0 0 0 0 0 0 1 0 85

0 0 0 0 4 0 0 0 0 0 0 0 0 0 8 19 1 nd 0 0 0 0 1 0 0 0 1 284

1 5 1 0 16 0 1 0 7 1 0 0 0 1 8 16 3 22 0 1 13 3 5 3 2 15 0 303

0 30 0 0 19 0 0 0 3 1 0 0 0 1 5 19 6 3 0 3 0 2 71 6 0 1 0 708

0 0 0 1 4 1 1 0 0 1 0 0 0 1 1 8 4 nd 0 0 53 0 1 0 15 38 0 2394

16 0 0 0 5 0 30 47 0 1 1 2 61 16 9 955 2 nd 6 0 5 nd 1 0 24 34 3 3,943

17 35 1 1 44 1 32 47 10 4 1 2 61 19 23 998 15 25 6 4 71 274 78 9 41 88 3 25

0 0 0 0 1 0 1 1 0 0 0 0 1 1 0 1 1 0 0 0 1 1 0 0 1 1 1 24

0 0 0 0 1 0 1 1 0 0 0 0 1 0 0 1 1 1 2 0 1 1 0 0 1 1 0 26

1 1 1 1 1 1 0 0 1 1 1 1 0 1 1 0 0 0 0 1 0 0 1 1 0 0 1 14

1 0 0 0 0 1 0 0 0 1 1 1 0 1 1 0 0 0 0 0 0 0 0 0 0 0 1

PACIFIC CORAL REEFS: AN INTRODUCTION 755

756

PACIFIC CORAL REEFS: AN INTRODUCTION

Pacific Coral Reefs: An Introduction, Figure 1 Mercator projection map of the marine geology and named archipelagos, trenches, and fracture zones in the Pacific Ocean Basin between Latitudes 30
N and 30
S (Source of map: National Geographic Society).

Pacific Coral Reefs: An Introduction, Figure 2 Fringing reefs encircling the Whitsunday Islands (20
19’S, Latitude and 145
56’E, Longitude) inside the Central Great Barrier Reef off north east Australia (Source Map Data Sciences Pty. Ltd. PSMA Image (2009) : Digital Globe (2009), Europa Technologies (2009) and Google Earth (2009).

PACIFIC CORAL REEFS: AN INTRODUCTION

3. Submerged reefs – living reefs on the tops of subsided volcanoes that are within water depths of 0–30 m, that may cause waves to break upon them, but that do not support emergent reefs or islands (Figure 4). 4. Atoll reefs – similar to the above except the reefs reach near the sea surface and encircle a protected lagoon without supporting emergent vegetated islands (Figure 5). 5. Atolls – similar to the above except that the atoll supports one or more vegetated islets and its reefs partially or wholly encircle a lagoon and allow regular surface interchange of waters between the ocean and lagoon (Figure 6). 6. Low reef islands – Similar to atolls in that they grow upwards off the crests of subsiding volcanoes, reach the sea surface, and support one or more emergent vegetated reef islets, but lack the size or shape to support a perimeter reef encircling a lagoon (Figure 7). 7. Raised limestone islands and atolls – similar to atolls and reef islets except that living reefs are limited to the submerged perimeter fringes facing the ocean. The interior parts are raised above sea level to elevations where reef builders (such as corals, coralline algae) cannot survive due to constant aerial exposure, causing the emergent part of the reef to weather and erode, often to limestone, and often creating inland marine lakes that lack regular surface water exchange with the ocean (Figure 8). 8. Closed atolls – similar to raised atolls except that they are near sea level but with interior marine lakes completely land-locked by the islands and blocking all surface seawater exchange with the ocean (Figure 9). The archipelagos and isolated islands in the Pacific can be grouped into about nine clusters based upon biogeography, geology, and cultural history (Figure 10). Beginning at due north and moving clockwise, these consist of: 1. Subtropical Hawaiian Islands – consisting of Hawaii and Johnston Atoll, including two dozen small to large volcanic basalt islands, several atolls, and two low reef islands. 2. Eastern tropical continental border – consisting of one closed atoll (Clipperton) and several volcanic island clusters off the continental shelves and slopes of the Americas from Mexico through Costa Rica and Panama to Ecuador and including the Galapagos. 3. Southeast tropical insular Pacific – consisting of scattered basalt volcanic islands, a hundred atolls and reef islets, and one raised limestone island (Makatea), and one raised atoll (Henderson) collectively within French Polynesia, Pitcairn, Rapa Nui, and Sala y Gomez. 4. Central tropical Pacific – consisting of about 50 atolls and reef islets and a few raised limestone and small basalt volcanic islands south of Hawaii and north of Samoa, and including the Phoenix, Line, Cook, Tokelau, Ellice [Tuvalu] archipelagos, Niue Island, and Rose Atoll.

757

5. Southern subtropical Pacific – consisting of small volcanic and reef islands north of New Zealand and east of Australia, including the Kermadec, Elizabeth, Middleton, Norfolk, and Lord Howe islands and reefs. 6. Southwestern continental border – consisting of the continental shelf islands and reefs from the southern end of the Great Barrier Reef to the Torres Strait. 7. Southwestern tropical Pacific – dominated by many large and small volcanic continental islands, numerous raised reef islands, and approximately 100 atolls and reef islets in Papua New Guinea, Solomon Islands, Vanuatu, Fiji, Tonga, Samoa, and New Caledonia. 8. Northwestern tropical Pacific – consisting of two dozen volcanic islands, several hundred raised reef islands, and about 100 atolls and reef islets including the Marshall Islands, Mariana Islands, Gilbert [Tungaru] Islands of Kiribati including Nauru and Banaba, Caroline Islands (Palau and Federated States of Micronesia), and a few offshore reefs and atolls east of the southern Philippines and Indonesia. 9. Northwestern subtropical continental border – including the Ogasawara [Bonin], Ryukyu, and Minami Torishima [Marcus] islands of Japan; Taiwan; and other small islands and reefs in open South China Sea north of the Philippines.

Geology Most of the contemporary Pacific Ocean is bounded by a “Ring of Fire” (Figure 11a), with dozens of volcanoes, deep ocean trenches, and mountain ranges lining the border that separates the oceanic (basalt) crust of the central Pacific Ocean basin from adjacent continental crust of islands and continents. About 200 million years ago, the Pacific Ocean was part of a single ancient global ocean before tectonic activity during the Permian began to fragment it along with the single continent (Pangaea) into smaller continental and ocean plates that have continued to collide and/or drift apart to this time. Eventually this led to segregation of the Pacific from the Indian Ocean and dissecting and fragmenting continental and oceanic crusts between the two emerging oceans to create what is now SE Asia. Over the past 100 million years, the massive Pacific Ocean plate has continued to produce and move new crust in a northwestern direction via convection flow beneath the East Pacific Rise and collisions with the adjoining Nazca Plate off the west coast of South America. The Pacific plate is continuing to move slowly towards the trenches of NE Asia, Siberia, and the Aleutians where eventually it is subducted under them. Additionally, the Asian continental plate is encroaching on the Pacific Ocean Plate from the NW, the Australian continental plate is encroaching from the west, and the two American continental and Nazca plates from the east, creating massive continental mountain ranges and trenches bordering the Pacific plate. These movements and collisions generate earthquakes, landslides, volcanism, tsunamis, uplift, and

758

PACIFIC CORAL REEFS: AN INTRODUCTION

Pacific Coral Reefs: An Introduction, Figure 3 Taha‘a Island and surrounding barrier reef (June 2000 image; 16
380 S Latitude and 151
300 W Longitude) and 10 km scale bar, Society Islands, French Polynesia (Source: Europe Technologies Image (2009), DigitalGlobe (2009), and Google Earth (2009)).

other seismic events. As the broad Pacific plate creeps towards the NW, parts of it move at different rates and directions, creating fracture zones, volcanic hotspots, earthquakes, subsidence, emergence, and flexures that have further contributed to the rise and fall of dozens of volcanic archipelagos at hotspots on the plate over the past 80–100 million years (Figures 11a and b). Overall, central Pacific archipelagos are now generally older towards the NW and younger towards the SE margins of the plate. Conversely, larger and more numerous islands are concentrated in the SW tropical Pacific but are virtually absent in the tropical and subtropical East Pacific. Deep ocean trenches separate the continents from the Pacific Ocean around most of the latter’s perimeter. Moreover, much of the southwestern Pacific region is bisected by several trenches stretching from the western end Samoa through Tonga, Vanuatu, Solomons, etc. (see Schellart et al., 2006).

Oceanographic and climatic processes affecting Pacific coral reefs A system of circular surface ocean currents, called gyres, dominates circulation in the Pacific and other large oceans

(Figure 12). The gyres move clockwise in the northern hemisphere and counterclockwise in the southern hemisphere, and greatly influence the diversity of reef life and distribution of Pacific reefs. In the northern hemisphere, the cold California Current runs south along the west coast of North America from Alaska to southern Mexico and then turns west where it is named the North Equatorial Current. As this currents runs west along the broader Pacific in low latitudes, it heats up before turning north, named the Kuroshio Current that runs along the east coast of Japan and Siberia, bathing the reefs at higher latitudes with warmer waters. Then the current turns east off Siberia where it is named the Subarctic Current that cools off as it travels east at high latitudes towards Alaska before it turns south as the California Current again, continuing its circular clockwise movement in the northern Pacific. The gyre in the southern hemisphere is a mirror image of the northern gyre. The cold Peru Current runs north along the west coast of South America, and turns west off Ecuador where it is named the South Equatorial Current that heats up along its long journey across the South Pacific at low latitudes. Some of this current moves west through the Torres Strait, but most runs through

PACIFIC CORAL REEFS: AN INTRODUCTION

Pacific Coral Reefs: An Introduction, Figure 4 Large complex of submerged reefs with 50 km scale bar (8
460 S Latitude and 149
590 E Longitude), off the Western D’Entrecasteaux Islands, Papua New Guinea (Source: U.S. Geological Survey (2009), Europa Technologies (2009), and Google Earth (2009)).

Pacific Coral Reefs: An Introduction, Figure 5 Example of atoll reef at Kingman Reef National Wildlife Refuge, Line Islands (6
240 N Latitude and 162
240 W Longitude). Length of reef is 15 km (Source of image: NASA and Lead Dog Consulting).

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Pacific Coral Reefs: An Introduction, Figure 6 Example of an atoll at Hereheretue Atoll, October 2006 (20
530 S Latitude and 144
580 W Longitude) with a 3 km scale bar, Tuamotu Islands, French Polynesia (Source of image: Europa Technologies (2009), DigitalGlobe (2009), and Google Earth (2009)).

Pacific Coral Reefs: An Introduction, Figure 7 Example of a low reef islet at Flint Island, March 2009 (11
260 S Latitude and 151
490 W Longitude), with a 5 km scale bar at the southern end of the Line Islands (Source of image: DigitalGlobe (2009), Lead Dog Consulting (2009), Geoeye (2009), and Google Earth (2009)).

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Pacific Coral Reefs: An Introduction, Figure 8 Example of a raised limestone island at Banaba [Ocean] Island, April 2005 (00
520 S Latitude and 169
320 E Longitude), west of the Gilbert Islands and with a 3 km scale bar (Source: DigitalGlobe (2009) and Google Earth (2009)).

the large islands of the SW Pacific before turning south off NE Australia and is known as the East Australia current that helps to extend corals and reefs as far south as Brisbane. Then at high latitudes, the current turns east to join the broad, cold West Wind Drift that encircles the globe north of Antarctica. A branch of the West Wind Drift moves towards South America and turns north off Chile and is known as the Peru Current again, continuing the gyre’s circular counterclockwise cycle. The forces that drive these currents are heat exchange between low and high latitudes, the NE trade-winds in the northern tropics, the SE trade-winds in the southern tropics, and the rotating Earth’s Coriolis Effect (describing mathematically the deflection of moving formations and masses, i.e., winds, ocean currents, clouds, aircraft, etc., caused by their inertia relative to the earth, a non-inertial uniformly rotating frame of reference). Trade-winds at the sea surface act to drag and mix the surface waters, pushing up sea level on the western ocean boundary to a slightly higher elevation vis-à-vis the eastern ocean boundary of the Pacific. The movement of both the NE and SE trade-wind systems towards low latitudes eventually causes them to collide and then rise skyward along an atmospheric band at about 5
N Latitude. These air masses then cool off and drop large amounts of rain in

a zone where surface winds are weak. This band of wet calm surface waters is the Inter-tropical Convergence Zone (ITCZ). The ITCZ weakens, strengthens, and shifts north and south depending on the strength of the tradewinds and the seasons. Because the surface trade-wind drag is weak or absent beneath the ITCZ, a return flow of ocean water moving “down-slope” from the elevated western Pacific to the lower eastern Pacific is maintained as the North Equatorial Countercurrent. There is also a smaller South Equatorial Countercurrent at approximately 8
S Latitude and a subsurface Equatorial Undercurrent that also runs west to east at the Equator. The eastward flow of water in these currents transports the larvae of many reef species from the biodiversity-rich western Pacific to the central and eastern Pacific. The warm water currents along the western boundary of the Pacific help to extend the range of many reef species to higher latitudes. Conversely, the cooler waters off the eastern boundary of the Pacific off the Americas may be another reason why reef development is generally poor in the East Pacific. All of the above current and wind systems are essential for exchanging heat between cooler polar areas and warmer tropical areas of the ocean and maintaining steady and predictable temperature regimes for corals and other reef life. Conversely,

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Pacific Coral Reefs: An Introduction, Figure 9 Example of a closed atoll, Clipperton Island (10
180 N Latitude and 109
130 W Longitude) and its marine lake SW of Mexico. Clipperton is one of the most northeastern reefs in the Pacific Ocean (Source: Europa Technologies (2009), DigitalGlobe (2009), NASA (2009), and Google Earth (2009)).

when the wind systems break down during El Nino Southern Oscillations (ENSO) at low latitudes, water temperatures rise, causing corals to bleach and/or die, and slowing down the flow of currents vital for survival of many migratory marine organisms including fish and seabirds (see entry “Climate Change and Coral Reefs”). Tropical cyclones are also concentrated along specific “alleys” in the western and eastern central Pacific, and their effects are sufficient to limit the three-dimensionality, growth, and biodiversity of many reefs, especially in the Philippine, Taiwan, Mariana, Caroline, Papua New Guinea, Solomon, Samoa, and Cook archipelagos, and lesser so in most other Pacific archipelagos.

Geography and biogeography of the tropical Pacific Ocean The oceanic, climatic, and geological history of the Pacific Ocean has led to the contemporary distributional patterns of its islands, reefs, and dominant reef life as showed by Figures 11a and b: 1. Mostly, atolls and reef islands dominate the central Pacific along an axis from the Pitcairn and Tuamotu Islands at the SE and Marshall Islands and Marcus Island at the NW end of the axis (Figure 13). This region of the ocean is geologically stable, allowing

ancient volcanic islands to subside and allow formation of atolls, a dozen barrier reefs, and reef islands in accordance with Darwin’s theory. 2. Numerous volcanic islands of continental origin are to the west and south of oceanic basalt islands to the north and east. Many of the SW Pacific islands are large, representing the consequence of considerable tectonic activity over long time periods (Figure 1). 3. Numerous raised limestone islands along the west Pacific boundary and isolated volcanic islands with limestone caps in the central Pacific populate the Pacific Ocean. This is likely the result of collisions between the Pacific, Asian, and SE Asian plates, causing the Pacific crust to buckle and push coral reefs above sea level, creating volcanic island arcs above the subduction zones (Figure 14). 4. Smaller high-volcanic island chains and seamounts are scattered haphazardly over the tropical Pacific crust near a combination of hotspots, fracture zones, and flexures that allow volcanism to penetrate upwards through to Pacific crust to form islands that may later subside and serve as the substrates for a variety of coral reefs (Figures 11a and b). At a Pacific-wide scale are vast expanses of the tropical Pacific that lack islands and coral reefs, particularly the eastern Pacific (Figures 11a and b), while the SW Pacific

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Pacific Coral Reefs: An Introduction, Figure 10 The location of the Eastern Pacific Barrier and the nine clusters of island archipelagos and coral reefs in the Pacific Ocean as explained in the text. (1) Subtropical Hawaiian Islands, (2) Eastern Continental Tropical Border, (3) Southeast Tropical Insular Pacific, (4) Central Tropical Pacific, (5) Southern Subtropical Pacific, (6) Southwest Tropical Continental Border, (7) Southwest Tropical Pacific, (8) Northwest Tropical Pacific, and (9) Northwest Subtropical Continental Border.

archipelagos have exceptionally high densities and larger coral reef habitats. The large timeframe and spatial distance of separation of East Pacific reefs from the rest of the reefs in the Central and Western Pacific has led to differentiation in evolution of species of corals and perhaps other reef-dwelling organisms. The broad zone between the two subregions lacking islands and reefs is referred to as the East Pacific Barrier. Additionally, there are still substantial distances between many of the archipelagos in the central and western Pacific, and their separate histories have also led to lesser levels of differentiation of their biological and geological characteristics and biodiversity. Most conspicuous is that the richness of coral reef and related marine species increases when moving from east to west across the Pacific. This increasing trend towards the western Pacific applies to all coral reef and related shallow-water tropical biota for which sufficient data are available, including: reef building corals, reef fishes (Figure 15), other reef invertebrates, algae, sea grasses, and mangroves. Ongoing research, including that of Coral Geographic scientists are now compiling patterns for coral species richness at the archipelagic scale and have further proposed a “coral triangle” area where coral diversity reaches its highest global levels, including SE Asia and the SW Pacific areas of the Solomons and Papua New Guinea. The pattern of increasing species that totals towards the west is generally maintained at the archipelagic level throughout the Pacific (Figures 16). However, recent

research and analyses by the present authors covering more than 2,200 site surveys at 70 reefs of various types in the Central Pacific reveal that there is significant small-scale heterogeneity between reefs based upon the coral species richness and the complement of coral species reported at each (Figures 17 and 18). Among these reefs were found 25 significantly different clusters, including substantial differences in species compositions and richness between adjacent reefs. Although data analyses had not been completed, the principal factors explaining these clusters appear to be:  Latitudinal     

proximity to the North Equatorial Countercurrent Longitude (distance from the “coral triangle”) Size of each reef or island Proximity to the nearest reef Size of the nearest reef, and Type of reef

These factors may help explain how especially geographic isolation, reef size, proximity to other reefs, and boundary currents can contribute substantially to our understanding of the biogeography of corals and their patterns of biodiversity, dispersal, and endemism throughout the tropical Pacific. However, these same factors may have different effects or be less influential for other reef biota, especially fishes that have greater mobility during both their adult and larval stages, and to algae that produce spores that remain viable over long time periods and distances. In

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the past, theories of radiation and accumulation have been put forth to explain the broader scale biodiversity levels in the western Pacific, but these are insufficient for explaining patterns at smaller spatial scales (Maragos and Williams, in preparation). Certainly there is mounting evidence that the “coral triangle” region has contributed many of the species in the Pacific, but it is also clear that localized endemism and speciation at isolated reefs have also contributed additional species within the distant and isolated corners of the Pacific Ocean.

Early cultural settlement of the central Pacific Ocean Linguistics, oral history, legends, field archaeology, radiocarbon dating, artifacts (fish hooks, pottery, etc.), pollen

analyses, and molecular analyses suggest that there are three broad cultural groups that settled within the insular tropical Pacific, also known as Oceania, after the last glacial recession about 10,000 year ago. These are still the principal cultures resident in most of the islands of Oceania today (Figure 19):  Melanesia. Settlers first reached the SW Pacific Islands

of Fiji, New Caledonia, Papua New Guinea, Solomon Islands, Vanuatu more than 10,000 years ago from Southeast Asia and the Indian Ocean regions (Bellwood 1978; Rainbird 1994).  Micronesia. Settlers from the Philippines first reached the NW Pacific islands about 5,000 years ago beginning with the Mariana Islands and then to the Caroline,

Pacific Coral Reefs: An Introduction, Figure 11 (a) Broad map of the Pacific Ocean plate tectonics showing the “Ring of Fire” (red lines) where the Pacific Plate is colliding with several continental plates, and about ten hot spots (yellow circles). A hotspot is a superheated stationary spot deep in the earth’s mantle resulting in upward movement of volcanic magma and penetration through the oceanic crust of the Pacific Plate and causing submarine volcanic eruptions and formation of island archipelagoes, in the Pacific Crust that have spawned volcanic archipelagos. The East Pacific Rise is at the SE corner of the Pacific where the Pacific and Nazca Plates collide, and about ten named longitudinal fracture zones are to the northwest. (Source of map: National Geographic Society). (b) Marine geological floor of the Pacific Ocean showing vertically exaggerated island groups, seamounts, fracture zones, trenches, and the East Pacific Rise along the SE boundary of the Pacific Plate with the Nacza Plate (Source of map: National Geographic Society).

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765

Pacific Coral Reefs: An Introduction, Figure 11 (Continued)

Marshall, and Gilbert (Tungaru) archipelagos and the two raised islands of Banaba (Ocean Islands) and Nauru, near the Gilberts.  Polynesia. The migrants departed from what is now the Maluku Province of Indonesia about 5,000 years ago and eventually settling in the Pacific about 3,000 years ago at Tonga and Samoa and then slowly moving eastward to what is now the Society Islands in the SE Central Pacific. Then between 1,100 and 600 years ago, a second, broader series of migrations occurred in several waves, beginning with Tahitians and later generations of evolving Polynesian cultural groups who moved: 1. Westward to what are now known as the Cook, Niue, Tokelau, Tuvalu, New Zealand Islands 2. Northward to the Marquesas and Hawaiian Islands 3. Eastward to what are now known as the Tuamotu, Pitcairn, and Easter [Rapa Nui] Islands 4. Southward to the Austral and Gambier Islands, and then 5. Northwestward towards the Carolines, Fiji, northern Solomon Islands, and northern Papua New Guinea, including islands previously settled by Melanesians and Micronesians. Samoans also reached the previously settled Gilbert and Tokelau Islands within the last 800 years. It is also likely

that Polynesians reached South America before the era of European rediscovery (Figure 20). Several island groups (Phoenix, Line) in the central Equatorial Pacific were visited by both Polynesians and Micronesians but were not permanently settled at the time of first European explorers 500 years ago, although some were later settled during the past century of European colonial rule.

Political history of the Pacific since the European rediscovery Spanish explorers were the first Europeans to discover and navigate across the Pacific Ocean in the early sixteenth century. Spain was also the first to establish colonies at Guam and several Northern Mariana Islands. Other Spanish, French, and British explorers during the sixteenth to eighteenth centuries were mostly exploring, seeking, or transporting gold, silver, pearls, spices, jewels, silks, and trepang. The French, British, and Americans eventually established colonies and/or developed plantation agriculture, forestry, ranching, or mining on some of the larger islands by the mid-nineteenth century (Fiji, New Caledonia, New Hebrides [Vanuatu], Tahiti, New Zealand, Australia, Hawaii). In 1876, Germany began establishing large colonies in the Pacific including Nauru, Papua New Guinea, Western Samoa, Northern Mariana, Caroline, and Marshall Islands. In turn, this prompted other

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Pacific Coral Reefs: An Introduction, Figure 12 Global projection of the Pacific Ocean showing the major surface currents in the northern and southern hemispheres. Red – warm western ocean currents, blue – cold eastern ocean currents, and yellow – cross-ocean currents at intermediate temperatures that are either warming up or cooling off (Source of base map: D Mapas (2009), Europa Technologies (2009), Tele Atlas (2009), and Google Earth (2009)).

world powers to annex islands or establish colonies in the rest of the strategically important Pacific. Of special importance was mining of guano because of its agricultural value as a fertilizer and ingredient for gunpowder. France and Great Britain established territories in French Polynesia, Wallis and Futuna, New Hebrides [Vanutau], Solomon Islands, Gilbert Islands, and Ellice [Tuvalu] Islands. The US annexed the Philippines, Hawaii, American Samoa, Guam, Wake Atoll, Midway Atoll, and the Panama Canal Zone between 1898 and 1903, and the US and Great Britain also claimed smaller islands previously mined for guano. After Germany declared war at the outbreak of World War I (WWI) in Europe in 1914, Japan seized German colonies in the Marshall, Caroline, and Mariana Islands while New Zealand, Australia, and Great Britain took custody of the remaining German possessions after WWI and the defeat of Germany. After 1935, Japan began fortifying its Pacific possessions, and the USA, Great Britain, and Australia claimed and/or fortified additional islands in the Line, Tuvalu [Ellice], Solomon, Gilbert, and Phoenix Islands and Papua New Guinea. After WWII and creation of the United Nations, a dozen inhabited trust territories were designated (Figure 20) with the goal of preparing them

for self rule and democracy; these primarily included inhabited islands claimed, captured, and colonized before and between the two world wars. Most of the Pacific islands are now independent republics or democratically governed territories and protectorates as voted upon and chosen by the resident populations (see Table 1). Only the Kingdom of Tonga has been spared the colonialism experienced by all other Pacific cultures during the past five centuries. The economies of the islands have now shifted to fisheries, tourism, agriculture, ranching, mariculture, mining, and military defense. The emerging island nations also continue to respect their traditional cultures and control of lands and marine areas for subsistence and conservation to maintain valued resources, including coral reefs.

Status of and threats to Pacific coral reefs The massive size of the Pacific Ocean combined with low human population densities and high buffering capacity has to date spared most of its coral reefs from the degradation experienced by reefs in the other oceans and seas where human populations are much denser. The largest marine-protected areas (MPAS) in the world are now established in the Pacific including the Great Barrier Reef

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PACIFIC CORAL REEFS: AN INTRODUCTION

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Bonin Midway Laysan

Marcus

Hawaii

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Johnston

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Pacific Coral Reefs: An Introduction, Figure 13 Distribution of low reef islands and atolls and about a dozen raised limestone islands in the central Pacific (Source: Stoddart (1992)).

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Pacific Coral Reefs: An Introduction, Figure 14 Distribution of high limestone islands primarily along in the western Pacific, with many to the west of the solid line. Black circles are Makatea Islands or volcanic islands with substantial Makatea limestone (Source: Stoddart (1992)).

PACIFIC CORAL REEFS: AN INTRODUCTION

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Pacific Coral Reefs: An Introduction, Figure 15 Lines of increasing families of warm water fishes towards the western Pacific Ocean (Source of data: Springer (1982) in Stoddart (1992)).

Pacific Coral Reefs: An Introduction, Figure 16 Coral species richness at Pacific archipelagos showing increasing trends towards the western Pacific and the Coral Triangle. Color legend shows white for lower levels of diversity through increasing shades of yellow, orange, and red for higher levels of diversity. Map after Coral Geographic, J.E.N Veron et al. (2009).

Pacific Coral Reefs: An Introduction, Figure 17 Color-based coral species richness levels at 70 individual islands and reefs in the Pacific. Color legend follows the same scale as in Figure 16, with white lowest and red highest in species richness per island or reef. Reef codes (in parentheses): A = atoll, B = barrier reef, C = closed atoll, I = volcanic island, L = raised limestone island, R = reef, and S = submerged reef. Explanations for island names, in alphabetical order for both Figures 17 and 18: AIL = Ailinginae (A), ANG = Angaur (I), ARN = Arno (A), AUN = Aunu‘u (I), BAB = Babeldaob (B), BAK = Baker (I), BIKA = Bikar (A), BIKI = Bikini (A), BIR = Birnie (I), BOK = Bokaak (A), CHU = Chuuk (B), ENE = Enewetak (A), ERI = Erikub (A), FAN = Fanna (I), FFS = French Frigate Shoals (A), FLI = Flint (I), Gardiner Pinnacles (I), Hawai‘i (I), HEL = Helen Reef (A), HOW = Howland (I), JAR = Jarvis (I), JEM = Jemo (I), JOH = Johnston (A), KAN = Kanton (A), KAU = Kaua‘i (I), KIM = Kimbe Bay (B), KIN = Kingman Reef (R), KIR = Kiritimati (A), KOS = Kosrae (I), KUR = Kure (A), LIS = Lisianski-Neva (A), MAJ = Majuro (A), MAL = Malden (L), MAR = Maro (R), MAU = Maui (I), MER = Merir (I), MID = Midway (A), MIL = Millennium (A), MIN = Minto (R), MOL = Moloka‘i (I), NEC= Necker (I), NGE= Ngeruangl-Velasco (R&S), NIH= Nihoa (I), NUK = Nukunonu (A), OAH = O‘ahu (I), OFU = Ofu-Olosega Islands (I), OROL = Oroluk (A), ORON = Orona (A), P&H = Pearl and Hermes (A), PAL = Palmyra (A), PEL = Peliliu (I), POH = Pohnpei (B), PUL = Pulo Anna (I), RAIT = Raita Bank (S), RON = Rongerik (A), ROS = Rose (A), SON = Sonsorol (I), STA = Starbuck (L), SWA = Swains (C), TAB = Tabuaeran (A), TAK = Taka (A), Ta‘u (I), TER = Teraina (I), TOB = Tobi (I), TUT = Tutu‘ila (I), VOS = Vostok (L), WAK = Wake (A), WOT = Wotto (A), and YAP = Yap Is (B). (Source: Maragos and Williams (in preparation)).

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Pacific Coral Reefs: An Introduction, Figure 18 Results of multivariate similarity profile analyses comparing the coral species compositions among the same 70 islands and reefs of Figure 17. Results show 25 statistically distinct clusters among the 70 reefs (green = < 60%, blue = 60–80%, red = > 80% similarity). Locations left unmarked form a “cluster” in isolation (e.g., JOH.) Lines connect spatially separated locations that form an individual cluster (e.g., POH and YAP). Five locations that are extremely spatially separated form an individual cluster, indicated by “þ”. Preliminary multivariate analyses indicate that longitude, proximity to the North Equatorial Countercurrent, size of reef, distance to nearest reef, and size of neighboring reefs all may be important contributing factors. Island codes listed in caption for Figure 20 (Source: Maragos and Williams (in preparation)).

of Australia, Northwestern Hawaiian Islands, several northern Line Islands, Phoenix Islands, much of New Caledonia’s reefs, and marine areas around the Mariana Trench. Several more of the northern Marshall Islands and the southern Line Islands may also be formally protected soon as MPAS, and there are hundreds of smaller MPAS now scattered throughout the Pacific. The major existing impacts to Pacific reefs are sea surface temperature anomalies, tropical cyclones, tsunamis, urban pollution, coastal construction, and over-fishing near population centers, exacerbating coral disease, bleaching and excessive predation by crown-of-thorn sea stars (Acanthaster), and localized dominance of invasive species. Additionally there are residual effects of WWII construction of military bases and related battles, and the postwar impacts of the USA, British, and French nuclear testing programs in the Pacific.

The future of Pacific coral reefs Despite the generally excellent status of many of its reefs at this time (Riegl and Dodge 2008), Pacific coral reefs

and related islands face an uncertain future. Past development proposals that would have adversely affected Pacific reefs may arise again: a sea-level canal through Central America; use of atolls and islands for garbage disposal, storage of nuclear waste, commercial missile-launching sites, oil transshipment facilities, and desalinization plants. The unauthorized poaching of fish and use of destructive fishing methods is already a major concern, especially near uninhabited islands and reefs where fish stocks are plentiful. Also, incidents of grounded and wrecked fishing vessels on remote reefs are on the rise. Emerging Asian economies and demand for fish are depleting tuna and other commercial fishing stocks in other oceans and East Asian seas, and compelling Pacific nations to enter into unenforceable fishery treaties that support foreign plans to establish fish processing bases throughout the tropical Pacific. During the past two centuries, wooden sailing vessels with poor navigation capabilities avoided remote reefs and islands for fear of crews being wrecked, grounded, stranded, or eaten by sharks. Now, thousands of modern steel-hulled fishing vessels with GPS are plying the Pacific in search of islands, reefs,

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Pacific Coral Reefs: An Introduction, Figure 19 Map showing the migration and settlement patterns of the Pacific islanders throughout Oceania over the past several millennia (Source of map: ORSTOM 1981).

sharks, and other fishing grounds, often far from the watchful eyes of residents and the limited enforcement assets of their governments.

Responses to global climate change Sea level, tropical cyclone, and temperature rises and possibly ocean acidification pose as serious threats to reefs worldwide (see entry “Climate Change and Coral Reefs”). However, many remote Pacific reefs may harbor greater resilience to these impacts due to the lack of present-day direct anthropogenic impacts. Although sealevel rise may destroy islands and force islanders off low islands and atolls, the reefs themselves may continue to survive and grow upwards. However, endemic and indigenous wildlife (nesting seabirds and sea turtles, resting shorebirds, Hawaiian monk seals) on low uninhabited and undisturbed atolls and reef islands will be most as risk as their island habitats are washed away. Emergency habitat development and translocations may be needed for many species to save them. Many of these islands now have alien rats and cats that prey on smaller ground-nesting seabirds, especially petrels, shearwaters, and terns. Many undisturbed Pacific reefs would serve as excellent field laboratories to assess the effects of global climate change not complicated by other anthropogenic effects. Appendix A is a listing and catalogue of intact reefs

(1) at the margins in the Pacific, and (2) along the Equator that could serve as excellent field laboratories for climatechange research, spanning the spectrum of Pacific climatic and oceanic habitats.

Summary and conclusions The Pacific Ocean is the World’s last true frontier on Earth. We are still in the discovery and inventory phase as to understanding the wealth and diversity of coral reefs and the reef life that inhabits or depends upon them. Pacific reefs remain in excellent condition at this time, although Asian economic expansion and global climate change are now beginning to threaten these coral reefs and their inhabitants. Islanders that dwell on low reef islands and atolls will likely need to be evacuated, while many terrestrial wildlife species on low uninhabited islands may need to be saved from extinction. The nations and cultures within Oceania should consider forming their own federation to protect and care for islanders who will need assistance and new homes, and save island wildlife species threatened with extinction. Some of the key elements of such a federation would be to:  Promote ecological and climate-change research to better respond to the future  Research and develop remote surveillance technology to discourage unauthorized fishing and access

Pacific Coral Reefs: An Introduction, Figure 20 Map showing post-WWII Pacific Ocean archipelagos, governments, and Trust Territories. Water current patterns are also shown (Source of map: National Geographic Society).

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Pacific Coral Reefs: An Introduction, Figure 21 (a) Kure Atoll (2004 image) at the northern tip of the Northwestern Hawaiian Islands, and the northernmost atoll in the Pacific Ocean (28
250 N Latitude and 178
200 W Longitude). Maximum diameter of atoll is 7 km. Image processed by D. Siciliano. (Source of image: Quickbird Imagery and DigitalGlobe (2009)). (b) Aerial image of Ducie Atoll (December 2004), the most southeastern atoll in the subtropical Pacific (24
410 S Latitude and 124
470 W Longitude) with a 2 km scale bar. (Source: Google Earth, Europa Technologies, and DigitalGlobe (2009)). (c) Aerial Image of Minami Torishima [Marcus Island], September 2003, the most northwestern low reef island in the Pacific (24
170 N Latitude and 153
590 E Longitude), with a 1 km scale bar. (Source: Google Earth and DigitalGlobe (2009)). (d) A partial aerial image of Helen Reef, one of the most western atolls in the tropical Pacific (2
540 N Latitude and 131
510 E Longitude) and showing a 10 km scale bar (Source: Google Earth (2009), and DigitalGlobe (2009)).

 Restore reefs degraded by WWII construction and

battles  Treaties with neighboring developed nations to assist in enforcement of territorial seas and access to remote surveillance technology  Establish a ocean-wide system of marine protected areas and employment opportunities for islanders to manage, monitor and restore habitats and species, and promote compatible visitation and education  Negotiate fishery treaties and other foreign development that would be in the best interests of all the peoples and resources of Oceania

Appendix A: Gallery of coral reefs at the margins and Equator of the Pacific Ocean The following islands and atolls are a few of the many coral reef formations at the margins of the Pacific Ocean basin, and all are in relatively pristine states and potentially important for various aspects of climate change and ecological research.  Kure Atoll, North Pacific, northernmost atoll in the Pacific (Figure 21a)  Clipperton Island, NE Pacific (Figure 21b), most northeastern reef in the Pacific

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PACIFIC CORAL REEFS: AN INTRODUCTION

a

b

c

d

Pacific Coral Reefs: An Introduction, Figure 22 (a) Google Ocean image of the Haimahera Sea, north of New Guinea where a dozen small and two large islands cross the Equator at 129
360 E Longitude. (Source: Europa Technologies (2009), Tele Atlas (2009), and Google Earth (2009)). (b) Google Ocean image of the Equator between several islands (Kuria, Aranuka, Abatiku, and Nonouti) in the Southern Gilbert Island of Kiribati at Longitude 173
490 E. (Source: Europa Technologies (2009) and Google Earth (2009)). (c) Google Ocean image of the Equator crossing northern Isla Isabela and close to the rest of the Galapagos Islands at 90
320 W Longitude. (Source: Europa Technologies (2009), Lead Dog Consulting (2009), and Google Earth (2009)). (d) Aerial image, Dec 2007, of Kapingamarangi with a 5 km scale bar, an atoll settled by Polynesians about 600 years ago. The atoll is within 1
N of the Equator at Longitude 154
460 E, and south of the Pohnpei State of the Caroline Islands (Source: DigitalGlobe (2009) and Google Earth (2009)).  Ducie Atoll, SE Pacific (Figure 10), most southeastern

atoll in the Pacific  Minami Torishima (Marcus Island), NW Pacific (Figure 21c), most northwestern reef in the Pacific where Japan maintains a weather station, and  Helen Reef (Figure 21d), one of the westernmost atolls in the Pacific One island in the eastern border of the Pacific (Isabela in the Galapagos) and a dozen islands in the Haimehera Sea, NW of New Guinea, at the opposite end of the Pacific

support the only reefs crossing the Equator within the Pacific Ocean. Several other reefs in the Northern Phoenix, Southern Gilberts, Southern Lines, and Kapingamarangi Atoll in the Caroline Islands fall within 1
of latitude of the Equator. Some of these islands are uninhabited, many are in healthy condition, and most could serve as important sites for climate-change research.  Jarvis Island National Wildlife

Refuge (NWR), central Line Islands (Figure 15.13, in Maragos et al., 2008)

PACIFIC CORAL REEFS: AN INTRODUCTION

 Banaba and Nauru Islands, west of the Gilbert Islands      

(see Figure 9) A dozen Haimahera Sea islands, north of New Guinea (Figure 22a) Several islands in the southern Gilberts (Figure 22b) Isla Isabela in the main Galapagos Islands in the Eastern Pacific (Figure 22c) Kapingamarangi Atoll south of the Caroline Islands (Figure 22d) Baker Island NWR in the northern Phoenix Islands (Figure 15.11, in Maragos et al., 2008), and Howland Island NWR in the northern Phoenix Islands (Figure 15.12, in Maragos et al., 2008).

Bibliography Allen, J., 1997. The impact of Pleistocene Hunter Gatherers on the Ecosystems of Australia and Melanesia. In Kirch, P., and Hunt, T. (eds.), Historical Ecology in the Pacific Islands: Prehistoric Environmental and Landscape Change. New Haven, CT: Yale University Press. Atwater, T., Sclater, J., Sandwell, D., Everinghaus, J., and Marlow, M. S., 1993. Fracture zone traces across the North Pacific Cretaceous Quiet Zone and their tectonic implications. In Pringle, M. S., Sager, W. W., Sliter, W. V., and Stein, S. (eds.), The Mesozoic Pacific: Geology, Tectonics and Volcanism. Geophysical Monograph Series, 77, Washington, DC: AGU, pp. 137–154. Bellwood, P., 1978. Man’s Conquest of the Pacific: The Prehistory of Southeast Asia and Oceania. Auckland, New Zealand: Collins. Bryan, E. H., Jr., 1953. A checklist of atolls. Atoll Research Bulletin, 19, 1–39. Coudray, J., and Montaggioni, L. F., 1983. Coraux et recifs coralliens de la province indopacifique: repartition geographique et altitudinale en relation avec la tectonique globale. Bulletin De La Societe Geologique De France, 24, 981–993. Dahl, A. L., 1980. Regional ecosystems survey of the South Pacific area. South Pacific Commission Technical Paper No. 179, Noumea. Dahl, A. L., 1991. Island directory. UNEP Regional Seas Directories and Bibliographies No. 35. United Nations Environment Programme, Nairobi, Kenya, 573 p. Daly, R. A., 1915. The glacial control theory of coral reefs. Proceedings of the American Academy of Arts and Science, 51, 155–251. Davis, W. M., 1928. The coral reef problem. Special Publications American Geographical Society, 9, 1–596. Ekman, S., 1953. Zoogeography of the Sea. London: Sedgwick & Jackson, xiv þ 417 p. Fischer, S., 2002. A History of the Pacific Islands. New York: Palgrave. Hezel, F. X., 2000. The First Taint of Civilization: A History of the Caroline and Marshall Islands in Pre-colonial Days, 1521–1885. Honolulu, HI: Unverisity of Hawaii Press, 372 p. IUCN (International Union for the Conservation of Nature and Natural Resources, Commission on National Parks and Protected Areas), 1986. Review of the Protected Areas System in Oceania. United Nations Environment Program/IUCN, Cambridge and Gland. Irwin, G. J., 1992. The Prehistoric Exploration and Colonisation of the Pacific. Cambridge: Cambridge University Press. Kleypas, J. A., Feely, R. A., Fabry, V. J., Langdon, C., Sabine, C. L., and Robbins, L. L., 2006. Impacts of Ocean Acidification on

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Coral Reefs and Other Marine Calcifiers: A Guide for Future Research. Report of a workshop held 18–20 April 2005, St. Petersburg, FL, sponsored by National Science Foundation, National Oceanic and Atmospheric Administration, and the US Geological Survey, 88 p. Koppers, A. A. P., Staudigel, H., Wijbrans, J. R., and Pringle, M. S.,1998. The Magellan seamount trail: implications for Cretaceous hotspot volcanism and absolute Pacific plate motion. Earth and Planetary Science Letters, 163, 53–68. Lessios, H. A., Kessing, B. D., and Robertson, D. R., 1998. Massive gene flow across the world’s most powerful biogeographic barrier. Proceedings of the Royal Society of London Series B, 265(1396), 583–588. Maragos, J. E., and Holthus, P. E., 1998. A status report on the coral reefs of the insular tropical Pacific. In Eldredge, L. C., Maragos, J. E., Holthus, P. F., and Takeuchi, H. F. (eds.), Marine and Coastal Biodiversity of the Tropical Island Pacific Region, Program on Environment, East-West Center, and Pacific Science Association c/o B.P. Bishop Museum, Honolulu. Maragos, J. E., and Williams, G. J., Biogeography of Central Pacific Corals, in prep. Maragos, J., Miller, J., Gove, J., DeMartini, E., Friedlander, A. M., Godwin, J., Musburger, C., Timmers, M., Tsuda, T., Vroom, P., Flint, E., Lundblad, E., Weiss, J., Avotte, P., Sala, E., Sandin, S., McTee, S., Wass, T., Siciliano, D., Brainard, R., Obura, D., Ferguson, S., and Mundy, B., 2008. US coral reefs in the Line and Phoenix Islands, Central Pacific Ocean: history, geology, oceanography, and biology. In Riegl, B., and Dodge, R. E. (eds.), Coral Reefs of the USA. Coral Reefs of the World. New York: Springer, Vol. 1, pp. 595–641. ORSTOM - L’office De La Recherche Scientifique et Technique de la Outre Mer, 1981. Atlas de la Nouvelle Caledonie et Dependances, Paris, Reimpression 1983, ISBN 3-7099-0601-5, planche 16, Archeologie et Prehistoire. Overpeck, J. T., Otto-Bliesner, B. L., Miller, G. H., Muhs, D. R., Alley, R. B., and Kiehl, J. T., 2006. Paleoclimatic evidence for future ice-sheet instability and rapid sea-level rise. Science, 311, 1747–1750. Rainbird, P., 1994. Prehistory in the northwest tropical Pacific: the Caroline, Mariana, and Marshall Islands. Journal of World Prehistory, 8(3), 293–349. Riegl, B., and Dodge, R. E. (eds.), 2008. Coral Reefs of the USA. Coral Reefs of the World. New York: Springer, Vol. 1. Schellart, W. P., Lister, G. S., and Toy, V. G., 2006. A Late Cretaceous and Cenozoic Reconstruction of the Southwest Pacific Region: tectonics controlled by subduction and slab rollback processes. Earth-Science Reviews, 76, 191–233, doi: 10.1016/ j.earscirev.2006.01.002. Schlanger, S. O., Jackson, E. D., Boyce, R. E., Cook, H. E., Jenkyns, H. C., Johnson, D. A., Kaneps, A. G., Kelts, K. R., Martini, E., McNulty, C. L., and Winterer, C. L., 1976. Initial Reports of the Deep Sea Drilling Project 33. US Government Printing Office, Washington, DC. Scott, G. A. J., and Rotondo, G. M., 1983. A model for the development of types of atoll and volcanic islands on the Pacific lithospheric plate. Atoll Research Bulletin 260, 1–33. Spalding, M. D., Ravilious, C., and Green, E. P., 2001. World Atlas of Coral Reefs. World Conservation Monitoring Center and United Nations Environment Program. Berkeley, CA/London: University of California Press, 424 p. Springer, V. G., 1982. Pacific plate biogeography, with special reference to shorefishes. Smithsonian Contributions to Zoology, 367, 1–182. Stehli, F. G., and Wells, J. W., 1971. Diversity and age patterns in hermatypic corals. Systematic Zoology, 20(2), 115–126. Stoddart, D. R., 1992. Biogeography of the tropical Pacific. Pacific Science, 46, 276–293.

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UNEP/IUCN (United Nations Environment Program and International Union for the Conservation of Nature and Natural Resources), 1988. Coral reefs of the world; Vo1. 1: Atlantic and Eastern Pacific; and Vo1. 3: Central and Western Pacific. UNEP Regional Seas Directories and Bibliographies. Switzerland/Cambridge, UK: IUCN, Gland; and Nairobi, Kenya: UNEP. Veron, J. E. N., 1986. Corals of Australia and the Indo-Pacific. Australian Institute of Marine Science. Honolulu, HI: University of Hawaii Press, 644 p. Veron, J. E. N., 1993. A biogeographic database of hermatypic corals. Australian Institute of Marine Science Monographs Series, 10, 433 p. Veron, J. E. N., 2000. Corals of the world. Australian Institute of Marine Science, Cape Fergeson, Australia, 3 vols. Veron, J., Devantier, L., Turak, E., and Kininmonth, S., 2009. Coral Geographic: A. progress report (abstract). www.coralreefresearch. org/html/crr_cg.htm Vitousek, P. M., 1994. Beyond global warming: ecology and global change. Ecology, 75, 1861–1876. Wiens, H. J., 1962. Atoll Environment and Ecology. New Haven, CT: Yale University Press.

Cross-references Acanthaster planci Antecedent Platforms Atoll Islands (Motu) Atolls Barrier Reef (Ribbon Reef ) Coral Cay Classification and Evolution Climate Change and Coral Reefs Daly, Reginald Aldworth (1871–1957) Darwin, Charles (1809–1882) Davis, William Morris (1850–1934) East Indies Triangle of Biodiversity Engineering On Coral Reefs With Emphasis On Pacific Reefs Fringing Reefs Glacial Control Hypothesis Global Ocean Circulation and Coral Reefs Great Barrier Reef Committee Indonesian Reefs Makatea Mass Extinctions, Anoxic Events and Ocean Acidification Oceanic Hotspots Plate Tectonics Ryukyu Islands Submerged Reefs Subsidence Hypothesis of Reef Development Tahiti/Society Islands Tethys Ocean Vanuatu

PACKSTONE Peter Flood University of New England, Armidale, NSW, Australia Packstone is a type of limestone recognized in the Dunham (1962) Classification Scheme that is grain supported and contains some sand-sized particles.

Bibliography Dunham, R. J., 1962. Classification of carbonate rocks according to depositional texture. In Ham, W. E. (ed.), Classification of Carbonate Rocks: American Association of Petroleum Geologists Memoir, pp. 108–121.

Cross-references Porosity Variability In Limestone Sequences

PALAEOSOLS Colin D. Woodroffe University of Wollongong, Wollongong, Australia

Synonyms Paleosols; Terra Rossa soils Definition A palaeosol (or paleosol) is, as its name suggests, an old soil. The term is particularly used to describe the reddish layers, also called terra rossa, that occur over the surface reef limestones and within fossil dune sequences. Soils form on subaerial exposures, such as on reef islands, or on the surface of reef limestones when these are exposed at times of lower sea level. Palaeosols, the remnants of former soils, are encountered where the soil has become lithified and has been preserved. The upper surface of reef limestones often has an incomplete cover of these former soils, together with other diagenetic evidence of exposure to terrestrial processes. These formed at times when the carbonates have been emerged, particularly during glaciations. Pleistocene reef limestones are often capped by a reddish brown crust, also called “caliche,” or calcrete, as across many of the islands in the Bahamas. A similar surface is encountered where drill cores intersect the upper surface of buried Pleistocene surfaces, as at discontinuities beneath Eniwetok, Mururoa, and Fangataufa, and on reefs on the Great Barrier Reef. These former soils are marked by micrite and sparry cements, as well as reddish-brown colouration. On several emergent islands, such as Nauru and Niue, there are extensive deposits of phosphate that are interpreted to have accumulated from the droppings of seabirds, and which is known as guano. Phosphate-rich soils occur on some reef islands where there are presently large numbers of seabirds, and cemented phosphate deposits on Holocene-age islands are similarly considered to be the remnants of soils developed in this manner, lithified, and preserved as palaeosols. Particularly well-developed palaeosols are associated with carbonate dune deposits. Fossil dunes, termed as aeolianite (or eolianite), are composed of sand-sized carbonate sediments, also known as calcarenite. On large calcarenite islands, such as Bermuda and the islands of

PALEOCLIMATE FROM CORALS

the Bahamas, the separate phases of dune accumulation are bounded by prominent red palaeosols. A lesser interruption in dune accretion is marked by a less welldeveloped soil horizon known as a protosol, or a hard layer of calcrete. Palaeosols formed during periods when the dunes were stabilized by vegetation. Despite an early view that the fossil dunes formed during periods of low sea level, such as the last glacial maximum, it is now known from extensive dating on many eolianite islands that most dunes accumulated during sea-level highstands. The clay-rich palaeosols mark those periods during which the sea was lower, the glaciations. Prominent in many palaeosols are fossils of land snails, such as Cerion in the case of the Bahamas, and Placostlyus on Lord Howe Island. Also widespread are the bones of seabirds, as well as rich assemblages of other fossils. Identification and dating of palaeosols and their correlation across islands enables discrimination of the intervening fossil dune formations. These sequences have been particularly effectively mapped and researched on the island of Bermuda.

Bibliography Baker, J. C., Jell, J. S., Hacker, J. L. F., and Baublys, K. A., 1998. Origin of insular phosphate rock on a coral cay – Raine Island, northern Great Barrier Reef, Australia. Journal of Sedimentary Research, 68, 1001–1008. Brooke, B., 2001. The distribution of carbonate eolianite. Earth Science Reviews, 55, 135–164. Hopley, D., Muir, F. J., and Grant, C. R., 1984. Pleistocene foundations and Holocene growth of Redbill Reef, south central Great Barrier Reef. Search, 15, 288–289. Vacher, H. L., and Quinn, T. M. (eds.), 1997. Geology and Hydrogeology of Carbonate Islands. Developments in Sedimentology 54. Amsterdam: Elsevier. Vacher, H. L., Hearty, P. J., and Rowe, M. P., 1995. Stratigraphy of Bermuda: nomenclature, concepts, and status of multiple systems of classification. Geological Society of America Special Paper, 300, 271–294.

Cross-references Calcrete/Caliche Eolianite Soils of Low Elevation Coral Structures

PALEOCLIMATE FROM CORALS Helen V. McGregor University of Wollongong, Wollongong, NSW, Australia

Synonyms Coral-climate proxies; Coral palaeoclimatology/ paleoclimatology; Paleoceanography; Past climates from corals

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Definition Coral paleoclimatology is the use of geochemical records from the skeletons of fossil or modern corals to reconstruct tropical climate variability during the time the coral lived. Introduction Ocean–atmosphere interactions in the tropics have farreaching consequences for climate variability across the globe. The tropics drive heat transfer to the poles, and tropical inter-annual oscillations such as the El Niño-Southern Oscillation (ENSO) and Indian Ocean dipole (IOD), via atmospheric teleconnections, affect rainfall patterns and climate conditions in areas far beyond the tropics (Ropelewski and Halpert, 1987), causing major socioeconomic impacts. Monitoring efforts have focused on improving observations and understanding of tropical climate variability, with the view to refining modeling of the tropical oceans and atmosphere. Despite these efforts, most instrumental records span only the past few decades and do not capture the full range of tropical climate variability, limiting our ability to model future changes. Coral paleoclimatology offers the prospect to extend instrumental records of tropical climate variability and can provide unique insights into tropical ocean–atmosphere interactions. Long-lived, massive corals record climate changes in the geochemistry of their skeletons. As corals grow, they deposit an aragonitic (calcium carbonate) skeleton, usually as one high and low density band per year, visible by x-ray of the coral skeleton (Barnes and Lough, 1993). Incorporated in the coral skeletons are varying proportions of geochemical elements, depending on the prevailing environmental and climatic conditions in the ambient seawater in which the corals live, and on the coral’s own physiology (see “Stable Isotopes and Trace Elements”). The geochemical, or the so-called “proxy,” records derived from the coral skeletons can be measured at sub-annual resolution and can be empirically related to a given climate parameter (e.g., Figure 1). Widely used coral proxies include the ratio of strontium to calcium (Sr/Ca), a proxy for sea surface temperature (SST), and the coral oxygen isotope ratio (d18O). The coral d18O is a function of SST and the d18O of seawater, where the d18O of seawater correlates with changes in sea surface salinity (SSS), which in turn may respond to changes in rainfall. Using the Sr/Ca SST proxy, the SST component of the coral d18O signal can be removed leaving the oxygen isotope residual (Dd18O) a SSS-only proxy (McCulloch et al., 1994; Gagan et al., 1998; Gagan et al., 2000). With the most commonly used coral genus Porites often living for a century or more, measurement of the coral skeletal geochemistry can provide quantitative, seasonally resolved, century-length records of climate variability for the time the coral lived. Reconstructing climate records from corals General approach. There are two main approaches to reconstructing past climates from corals using modern or

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Paleoclimate from Corals, Figure 1 IOD climate anomalies. (a) SST anomalies during November of the 1997 IOD event (Reynolds and Smith, 1994), when anomalous cooling in the east and warming in the west produced a reversal of the equatorial SST gradient across the Indian Ocean. Boxes mark the eastern and western sectors used to define the dipole mode index (Saji et al., 1999), and the white circle shows the location of the Mentawai Islands. (b) In the Mentawai Islands the strong IOD events of 1994 and 1997 were characterized by cool SST anomalies (black curve; Saji et al., 1999) and drought (gray bars; Xie and Arkin, 1996). These distinct IOD SST and rainfall anomalies are preserved, respectively, in coral Sr/Ca SST (red circles and curve) and Dd18O (blue circles and curve) anomalies (Gagan et al., 1998). Coral time series between July 1993 and February 1997 are based on the average of two coral records, with error bars showing the difference between the coral records for each monthly data point. (Reprinted by permission from Abram et al. [2007].)

fossil corals. The most common approach is to collect a core from a live coral on a reef. The chronology is provided from the coral density growth bands and/or the seasonal cycle recorded in the coral geochemical records. Annual coral density bands can be counted; beginning at the tissue layer at the uppermost surface of the coral, the paired high and low density bands, revealed by x-ray of the coral, are counted back in time. Coral density bands are not always well defined, so alternatively the annual peaks or troughs in the coral geochemical climate proxies can themselves be counted. Using either method, or a combination of the two, it may be possible to count back several centuries with errors of 1 or 2 years (Lough and Barnes, 1997; Guilderson and Schrag, 1999; Hendy et al., 2002; Hendy et al., 2003). The advantage of using a modern coral for climate reconstruction is that the start age is known and the coral geochemical proxies can be

directly calibrated against climate parameters as measured in the instrumental records increasing the robustness of the reconstruction. In addition, replicate coral cores can be taken from different coral heads from the same reef, enhancing reproducibility. The disadvantage is that the reconstructed record is limited by the length of time the coral lived, usually no longer than a few centuries. Dating errors may increase further back in time. There are now significant numbers of modern coral records extending back several centuries from the present day from various locations throughout the tropics such that the records are being combined to examine the climate connectivity of the tropical oceans (Kaplan et al., 1998; Evans et al., 2000; Hendy et al., 2002; Charles et al., 2003; Wilson et al., 2006). An emerging alternative approach is to use wellpreserved fossil corals to reconstruct climate further back

PALEOCLIMATE FROM CORALS

Paleoclimate from Corals, Figure 2 Fossil Porites sp. coral from Muschu Island, Papua New Guinea. The coral, once living underwater on the reef, has been uplifted by tectonic activity and is now preserved as an isolated coral head within the intertidal zone, and is exposed at low tide. Note the concentric growth bands in this coral and that the top of the coral has been planed flat.

in time (Figure 2). The age of the fossil coral can be precisely determined using radiocarbon or U-series dating, and the proxy climate record from the coral skeleton provides a “snapshot” of past variations (e.g., Beck et al., 1997; Gagan et al., 1998; Tudhope et al., 2001; Woodroffe et al., 2003; Corrège et al., 2004; Felis et al., 2004; McGregor and Gagan, 2004; Abram et al., 2007). It is also essential to analyze a modern coral that overlaps instrumental record to quantify the coral geochemistry–climate parameter relationship since the same climate proxies are used for the fossil and modern corals. In this case, the modern records for calibration need not extend back several centuries. The major advantage of using fossil corals is that it is possible to look at climate variations further back in time when the climate boundary conditions were different from today, giving a perspective on the range of possible, natural climate modes. The disadvantage of using fossil coral proxy records is that they are almost inevitably disjointed. However, an exciting new development has been to use high-precision U-series dating of multiple fossil coral and look for age overlaps (Cobb et al., 2003; Zhao et al., 2009). Age dating errors are typically þ/0.5%, equivalent to less than a decade for Holocene-aged corals, and where individual fossil corals overlap in age their proxy records can be pieced together, where the climate “wiggles” from each individual coral are matched, to produce a longer record (Cobb et al., 2003). Coral-climate signals and reproducibility. For both fossil and modern coral studies, it is essential to establish the climate factors controlling the signal at a given location, quantify this relationship, and determining the magnitude of associated errors. For example, the commonly use d18O proxy reflects changes in both SST and SSS. However at some locations, such as in PNG in the western tropical Pacific SSS changes dominate the signal (Tudhope et al.,

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2001; McGregor, 2004), whereas at locations such as the central Pacific SST exerts the major control (Evans et al., 2000; Woodroffe and Gagan, 2000). In addition, the aim for coral paleoclimatology is often to be able to use the coral proxy record from a single location to infer regional climate variability. Thus, the climate factors that control the coral proxy climate signal must be quantifiably related to regional climate variations (Guilderson and Schrag, 1999). Some locations appear to be “nodes” for particular tropical climate oscillations (e.g., ENSO or the North Atlantic Oscillation), and coral records from these locations can be representative of climate variability over very large spatial areas. One example of this is where coralclimate records from Kiritimati Island were used to construct a coral-C INDEX, equivalent to the NINO3.4 Index used to define ENSO events (Evans et al., 1998). From this point of view, corals from well-flushed reef settings are preferred. For in situ fossil corals it is usually possible to establish the paleo-reef morphology; however, for transported corals (dislodged either by storm activity or post-uplift erosion) establishing provenance is almost impossible (Cobb et al., 2009). Several studies have suggested that proxy climate signals from corals living on the same reef may show higher between-coral (inter-reef) differences than that expected from climate variability alone (Guilderson and Schrag, 1999; Linsley et al., 1999; Cohen et al., 2002; Felis et al., 2003). The origins of these offsets are not well understood, and, in addition, there are studies that suggest minimal between-coral offsets (Gagan et al., 1998; Hendy et al., 2002; Stephans et al., 2004). But between-coral differences appear to affect the coral oxygen isotope ratios more than Sr/Ca. Coral growth form, growth rate, reef setting, and diagenesis have all been suggested as possible causes of offsets (Cobb et al., 2009). Corals can also show within-coral variations (McConnaughey, 1989; de Villiers et al., 1995; Alibert and McCulloch, 1997; Cohen and Hart, 1997). Regardless of the origin and magnitude of between-coral and within-coral offsets, reproducing proxy climate signals from the same location, using at least two corals, will reduce uncertainty and quantify errors in climate reconstructions from both modern and fossil corals (Lough, 2004; Stephans et al., 2004; Abram et al., 2009). Diagenesis. An underlying assumption in the use of corals to reconstruct climate is that the corals are pristine, that is, that original coralline aragonite is preserved. However, coral skeletons are susceptible to a process known as diagenesis. Diagenesis is the precipitation of secondary aragonite or calcite in skeletal voids, or the replacement of skeletal aragonite, usually with calcite. If even a small amount of diagenetic material is included in a sample, then the resulting climate reconstruction may be rendered inaccurate. This is because isotopes and trace elements are exchanged and removed during the diagenetic transformation, changing the geochemistry of the coralline matrix. For example, secondary calcite can lead to “warm” SST artefacts (SST appears warmer than was the case), particularly in Sr/Ca SST reconstructions, of 1
C or more

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(McGregor and Gagan, 2003). Secondary aragonite, typical for early marine diagenesis of modern corals, causes considerable alteration of coral geochemistry and creates “cool” SST artefacts (Bar-Matthews et al., 1993; Enmar et al., 2000; Müller et al., 2001; Lazar et al., 2004; Müller et al., 2004; Allison et al., 2007). Dissolution was shown to create “cool” anomalies in a range of trace element SST proxies (Hendy et al., 2007). Diagenesis has been observed in living corals and corals just a few decades old (Enmar et al., 2000; Müller et al., 2001; Hendy et al., 2007; Nothdurft et al., 2007); however, fossil corals are more susceptible to diagenetic processes as they are exposed to seawater, sea spray, rainfall, and/or groundwater for centuries to millennia or longer. Numerous studies have called for screening for diagenesis in modern and fossil corals to become standard procedure (McGregor and Gagan, 2003; Gagan et al., 2004; Quinn and Taylor, 2006; Allison et al., 2007; Hendy et al., 2007; Cobb et al., 2009). Diagenesis that produces an addition of 10% or more secondary material to the skeletal bulk density can be detected in coral density profiles, X-radiographs, and UV luminescence photos (Hendy et al., 2007). However, just 1% calcite diagenesis would alter the coral climate signal (McGregor and Gagan, 2003). X-ray diffraction (XRD) has been the most commonly used diagenesis screening method in coral paleoclimate studies. However, XRD detects secondary calcite, not secondary aragonite or dissolution, and the XRD detection limit is effective down to around the 1% calcite (McGregor and Gagan, 2003; Allison et al., 2007). For lower levels of diagenesis, thin section is proving a highly effective means of screening samples for all types of diagenesis (McGregor and Gagan, 2003; McGregor and Abram, 2008; Cobb et al., 2009). One thin section per 20–30 years of coral growth is recommended (Cobb et al., 2009) and a guide has been developed to assist coral paleoclimatologists in screening using thin sections (McGregor and Abram, 2008). Scanning electron microscopy (SEM) is also a highly effective tool for diagenetic screening (Nothdurft et al., 2007).

Coral paleoclimate records One of the major strengths of corals as paleoclimate archives is their ability to record monthly or finer resolution (e.g., weekly) climate information. This level of resolution means corals are ideally suited to investigating inter-annual climate phenomena, such as the El NiñoSouthern Oscillation and the Indian Ocean Dipole. Approximately 90 coral proxy records based on corals that were alive at the time of collection have been published, with around 30 of those extending from the late twentieth century back prior to 1900 (Jones et al., 2009). These longer coral records are able to provide a wealth of information on multidecadal to centennial timescale variability. At least one coral record exists from every tropical ocean, with the vast majority of records from the tropical Pacific. There are significantly fewer records from

older fossil corals. Coral paleoclimatology has made a significant contribution to understand a number of facets of tropical climate variability and a snapshot of these contributions will be discussed in the following sections.

Interannual and multi-decadal variability in the tropical Pacific: variations past and present A key focus of coral paleoclimate research has been in extending records of the El Niño-Southern Oscillation system. ENSO, with its origins in the equatorial Pacific, is the largest source of interannual climate variability across the planet. It is a coupled climate system between the atmosphere and the ocean, and it oscillates irregularly on a timescale of 2–7 years. The average state of the equatorial Pacific Ocean involves strong easterly trade winds pushing warm water to the west, which then brings (upwells) cool subsurface water in the east. This sea surface temperature gradient then reinforces the easterly winds. A La Niña event is an enhancement of the average state, where the easterly winds and temperature gradient strengthen. In an El Niño event, the easterly winds slacken and reduce the ocean temperature gradient, allowing warm water to flow back to the central and eastern Pacific. Atmospheric convection and precipitation follows the warmest water, and in an El Niño, above average rainfall is deposited over the central Pacific. El Niño events have become stronger and more frequent since the mid-1970s (McPhaden et al., 2006). However, there is debate about whether strengthened El Niño is a result of global warming or whether the 1970s intensification resulted from a decadal or longer-term cycle of ENSO variability (Fedorov and Philander, 2000; Cane, 2005; McPhaden et al., 2006). The relatively short instrumental record does not give a complete picture of ENSO behavior, and models are unable to simulate ENSO fully, limiting our ability to predict future ENSO scenarios. Corals from optimal locations across the equatorial Pacific are able to capture the SST and rainfall/SSS variations that result from ENSO oscillations (Cole et al., 1993; Dunbar et al., 1994; Evans et al., 1998; Urban et al., 2000), and have complemented and extended instrumental records of ENSO variability. Spectral analysis of one of the earliest published coral d18O records, that from Tarawa Atoll, western Pacific, showed changing dominance of seasonal and interannual variability through the twentieth century, suggesting a change in the ENSO “pulse” (Cole et al., 1993). Further analysis of the Tarawa d18O record and comparison with a 155 year d18O record from Maiana Atoll, just south of Tarawa, revealed, among other things, that the 1976 shift in ENSO variability coincided with changes in the mean background climate of the tropical Pacific, with the ENSO period shifting from 2.9 to 4 years across the 1976 change (Urban et al., 2000). The record also revealed that during the late nineteenth century, ENSO cycles lasted 10–15 years. These results are significant because they show that the length of the ENSO cycle varies with small changes in tropical climate.

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Furthermore, the results suggest that there is substantial natural variability in ENSO, but that the most recent period may be unique in that the changes in the mean climate due to anthropogenic greenhouse gases may also influence ENSO (Dunbar, 2000). The network of coral-climate records from across the Pacific, although still sparse, does appear to capture large-scale, multi-decadal climate variance, and in combination with coral records from other ocean basins, is revealing ocean–atmosphere connections right across the tropics. Consensus is emerging for significant decadal variability in tropical Pacific SSTs at the 9–14 year period (Cobb and Charles, 2001; Holland et al., 2007; Ault et al., 2009). The decadal variability displays an ENSOlike spatial and temporal pattern suggesting that the Pacific decadal-scale variance is directly related to ENSO (Urban et al., 2000; Holland et al., 2007; Ault et al., 2009). Decadal-scale variability appears to be stronger in the late nineteenth century, and analysis based on only twentieth century records may underestimate decadal-scale variability (Ault et al., 2009). Synthesis of coral SST records spanning the whole tropics suggests that the late twentieth century is the warmest period for the past 250 years, which can be explained by the increase in anthropogenic greenhouse gases in the atmosphere (Wilson et al., 2006). A consistent antiphased correlation between the South Pacific Convergence Zone, a low pressure trough extending from around the Solomon Islands to French Polynesia, and the decadal variability in the central equatorial Pacific has been documented back to 1650 AD (Linsley et al., 2008). At the decadal-scale the Atlantic Ocean may be directly influenced by SST anomalies in the central tropical Pacific (Cobb and Charles, 2001), and central equatorial Pacific SSTs also correlate on a variety of timescales with Indian Ocean coral oxygen isotope records, suggesting a decadal-scale connection between these ocean basins (Cobb and Charles, 2001; Charles et al., 2003). The longest, continuous coral proxy record published to date, based on eight coral cores from the Great Barrier Reef (GBR), sampled at 5-yearly resolution and totaling 420-years from present back to 1565 AD, showed that corals can also record centennial-scale shifts in tropical climate (Hendy et al., 2002). This study was unique in its approach in that it combined records in a similar manner to the approach used in tree-ring studies to extract the robust “common signal” of low-frequency SST and SSS variability in the records. The resulting common GBR SSS anomaly record showed increased salinity between 1565 and 1870, around the time of the Little Ice Age (LIA) in the northern Hemisphere (Hendy et al., 2002). The SST anomaly results, along with other long records from the tropical Pacific, suggest a stronger latitudinal temperature gradient during the LIA, enhancing winddriven evaporation, giving rise to the high salinity anomalies. The authors suggest that the high evaporation equated to a net export of moisture from the tropics contributing to glacial advance during the LIA (Hendy et al., 2002).

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Also, drawing on approaches adapted for tree-ring studies, Cobb et al. (2003) pieced together numerous fossil and modern corals from ENSO-sensitive Palmyra Atoll, central Pacific, to give the first picture of ENSO variability for the past millennium. Their landmark reconstruction was based on dislodged coral heads deposited on Palmyra during storms events. The corals were precisely dated by the U-Th method to identify periods of overlap, and then the monthly resolved d18O records were “wiggle matched” together. In total, this resulted in 430 years of monthly resolved record, for five intervals across the past millennium. The record showed a surprising degree of variability in ENSO strength, independent of changes in solar or volcanic forcing. In addition, ENSO strength does not appear to relate to changes in Northern Hemisphere climatic periods such as the Little Ice Age or Medieval Warm Period (Cobb et al., 2003). ENSO strength seemingly switches modes in a matter of decades, and the late twentieth century ENSO, although strong, is not unprecedented over the length of the record (Cobb et al., 2003). The results suggest that ENSO can change its character on its own, and in future, may shift with or without an additional push from anthropogenic greenhouse gases (Cobb et al., 2003; Tudhope and Collins, 2003). This presents a challenge for modeling future ENSO behavior, and testing global climate models against past ENSO variability provides a means to refine and improve these models. Coral paleoclimate reconstructions have been instrumental in understanding the origin and long-term evolution of ENSO (Hughen et al., 1999; Corrège et al., 2000; Tudhope et al., 2001; Woodroffe et al., 2003; Kilbourne et al., 2004; McGregor and Gagan, 2004; Sun et al., 2005). A major discovery from Indonesian and Papua New Guinea (PNG) fossil corals has been that ENSO has been a component of the tropical climate system for at least 130,000 years (Hughen et al., 1999; Tudhope et al., 2001), and has varied significantly in strength through glacial and interglacial cycles (Tudhope et al., 2001). Tudhope et al. (2001) proposed that ENSO strength varied as a result of changes in tropical Pacific seasonality related to the Earth’s orbital cycles (orbital forcing) plus dampening of ENSO strength during glacials. Further work on coral from PNG, and additional coral records from the central Pacific, are providing a more detailed picture of the evolution of ENSO for the most recent Holocene period (10,000 years ago to present; Corrège et al., 2000; Tudhope et al., 2001; Woodroffe et al., 2003; Gagan et al., 2004; McGregor and Gagan, 2004; Sun et al., 2005). These studies suggest that ENSO was less active compared to today, although still present, and may have been more active around 2,000 years ago. A key component of ENSO and the tropical Pacific climate system is the Indo-Pacific Warm Pool (IPWP). The IPWP is the warmest body of ocean water in the world having an average temperature of >28
C, and coupled ocean–atmosphere interactions in the IPWP are not only thought to trigger El Niño events, but also deliver large

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amounts of heat to the atmosphere via the global atmospheric Hadley-Circulation (Webster, 1994). Much coral evidence for the evolution of ENSO originates from the IPWP and coral data has also provided important information on the long-term climate variability of the warm pool. Sea surface temperature estimates, based on corals from Vanuatu, PNG, GBR, and Indonesia and utilising a revised coral Sr/Ca SST calibration, suggest that the early Holocene IPWP was 1–3
C cooler than present (Gagan et al., 2004 and references therein). Temperatures had reached modern values by 8.5 ka, in general agreement with planktonic foraminifera Mg/Ca and alkenone SST estimates, and mid-Holocene SSTs were 0.5–1
C warmer than present (Gagan et al., 2004). A more detailed picture of IPWP SST variations, based on 48 fossil corals from Indonesia and PNG, suggested that the southern margin of the IPWP cooled and warmed periodically through the mid-Holocene, associated with an expansion and contraction of the warm pool as the Asian summer monsoon weakened and strengthened, respectively, and points to the fundamental importance of the warm pool in propagating climate change (Abram et al., 2009).

Interannual modes in the Atlantic and Indian Ocean Understanding the influence of the tropical Pacific and ENSO on the Indian Ocean has been recognised as an important issue and corals have played a role in better characterising the interactions. For the past few decades, SSTs in the Indian Ocean have been associated with ENSO and the Asian monsoon via complex multifeedbacks (Webster et al., 1998). As with the tropical Pacific, however, a major limit in understanding climate variability in the tropical Indian Ocean, and the Indian, East African, and southeast Asian monsoons, on which millions of people are dependent for life-giving rain, has been the absence of records of more than a few decades in length. Multi-century coral oxygen isotope records from the central and western Indian Ocean however, reveal a clear ENSO teleconnection, whereby during El Niño events western Indian Ocean SSTs are warmer (Charles et al., 1997; Cole et al., 2000; Zinke et al., 2005; Pfeiffer and Dullo, 2006). The western Indian Ocean SSTs also display decadal-scale ENSO-like variability (Cole et al., 2000; Pfeiffer and Dullo, 2006), which may be linked to the Pacific decadal oscillation (Crueger et al., 2009). Although there may be a generally consistent teleconnection between the western Indian Ocean and ENSO, new perspectives from coral records from the eastern Indian Ocean show a changing ENSO–Indian Ocean–monsoon interaction, with potentially negative consequences for SE Asian rainfall under global warming scenarios (Abram et al., 2007). Recently identified IOD events are defined by a reversal of the equatorial Indian Ocean east–west SST gradient and zonal winds from their mean climatological state (Saji et al., 1999; Figure 1),

resulting in drought for western Indonesia and southern Australia (Overpeck and Cole, 2007). IOD events were reconstructed back to 1846 using a suite of corals from the eastern and western Indian Ocean (Abram et al., 2009). The coral IOD index showed an increase in the strength and frequency of IOD events through the twentieth century (Figure 1). The results also showed that despite the historical influence of ENSO in triggering IOD events, the twentieth century IOD intensification was a direct result of IOD–monsoon feedbacks, with consequences for rainfall distribution (Abram et al., 2009). Coral reconstructions of Holocene IOD events have likewise revealed changes in the IOD–monsoon–ENSO relationship over time (Abram et al., 2007). The results are significant, their strength being that they provide information on monthly scale timing and duration of individual IOD events (Figure 1), and show that 6,500 years ago IOD SST cooling lasted 5 months, up from 3 months for the present day. They also reveal that long-duration IOD events result in drought peaking later in the calendar year than for the present day, coinciding with what would normally be the maximum monsoon rainfall in western Indonesia. The mid-Holocene enhanced IOD cooling and drying is thought to result from strengthened Asian monsoon, and future Asian monsoon–IOD strength may extend drought through Australasia (Abram et al., 2007). Coral records are making a significant contribution to our understanding of the Atlantic Multidecadal Oscillation (AMO), hurricane activity, and the North Atlantic Oscillation/Arctic Oscillation (NAO/AO). The AMO is a decadal- to multidecadal-scale variation in SSTs across the Atlantic and may be related to hurricane activity in the northern tropical Atlantic (Goldenberg et al., 2001). A 440-year growth rate SST reconstruction from the Bahamas suggests that multidecadal variability may only be significant after 1730, limiting the accuracy of decadal climate forecasts (Saenger et al., 2009). In a related study, luminescence lines from Caribbean corals were used in combination with a marine sediment core proxy record to investigate hurricane frequency, vertical wind sheer, SST, and the AMO for the past 270 years (Nyberg et al., 2007). The results suggest that vertical wind sheer is more important than SST in controlling hurricane frequency, and controversially, that increased hurricane frequency since 1995 is not unusual compared to other periods of high hurricane activity in the record (Elsner, 2007; Nyberg et al., 2007; Neu, 2008; Nyberg et al., 2008). Coral records, particularly those from fossil corals, have made key inroads into our understanding of the NAO/AO. The NAO/AO is the dominant interannual atmospheric mode in the North Atlantic influencing the climate of much of Europe via its modulation of the strength of the subpolar westerlies (Hurrell, 1995; Thompson and Wallace, 2001). Oxygen isotopes in corals from the northern Red Sea reflect the influence of the NAO/AO and ENSO on the climate of the region (Felis et al., 2000). Sr/Ca SST records from 125,000 year old fossil corals show increased temperature seasonality

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(Felis et al., 2004). In an innovative approach, combining the seasonal-resolution coral data with coupled atmosphere–ocean model simulations, the results suggest that a more positive NAO/AO operated 125,000 years ago contributing to the increased seasonality in the Red Sea coral record, as a response to insolation changes from differences in Earth’s orbital configuration at that time (Felis et al., 2004). Such studies help to disentangle natural climate variability and will help us better understand the response of interannual climate oscillations to anthropogenic greenhouse warming.

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Linsley, B. K., Messier, R. G., and Dunbar, R. B., 1999. Assessing between-colony oxygen isotope variability in the coral Porites lobata at Clipperton Atoll. Coral Reefs, 18, 13–27. Linsley, B. K., Zhang, P., Kaplan, A., Howe, S. S., and Wellington, G. M., 2008. Interdecadal-decadal climate variability from multicoral oxygen isotope records in the South Pacific Convergence Zone region since 1650 A.D. Paleoceanography, 23, PA2219. Lough, J. M., 2004. A strategy to improve the contribution of coral data to high-resolution paleoclimatology. Palaeogeography, Palaeoclimatology, Palaeoecology, 204, 115–143. Lough, J. M., and Barnes, D. J., 1997. Several centuries of variation in skeletal extension, density and calcification in massive Porites colonies from the Great Barrier Reef: a proxy for seawater temperature and a background of variability against which to identify unnatural change. Journal of Experimental Marine Biology and Ecology, 211, 29–67. McConnaughey, T., 1989. 13C and 18O isotopic disequilibrium in biological carbonates: I. patterns. Geochimica et Cosmochimica Acta, 53, 151–162. McCulloch, M. T., Gagan, M. K., Mortimer, G. E., Chivas, A. R., and Isdale, P. J., 1994. A high resolution Sr/Ca and d18O coral record from the Great Barrier Reef, Australia, and 1982–83 El Niño. Geochimica et Cosmochimica Acta, 58, 2747–2754. McGregor, H. V., and Abram, N. J., 2008. Images of diagenetic textures in Porites corals from Papua New Guinea and Indonesia. Geochemistry Geophysics Geosystems, 9, Q10013. McGregor, H. V., and Gagan, M. K., 2003. Diagenesis and geochemistry of Porites corals from Papua New Guinea: implications for paleoclimate reconstruction. Geochimica et Cosmochimica Acta, 67, 2147–2156. McGregor, H. V., and Gagan, M. K., 2004. Western Pacific coral d18O records of anomalous Holocene variability in the El Niño-Southern Oscillation. Geophysical Research Letters, 31, L11204. McPhaden, M. J., Zebiak, S. E., and Glantz, M. H., 2006. ENSO as an integrating concept in Earth science. Science, 314, 1740–1745. Müller, A., Gagan, M. K., and McCulloch, M. T., 2001. Early marine diagenesis in corals and geochemical consequences for paleoceanographic reconstructions. Geophysical Research Letters, 28, 4471–4474. Müller, A., Gagan, M. K., and Lough, J. M., 2004. Effect of early marine diagenesis on coral reconstructions of surface-ocean 13 12 C/ C and carbonate saturation state. Global Biogeochemical Cycles, 18, GB1033. Neu, U., 2008. Is recent major hurricane activity normal? Arising from: Nyberg et al. Nature 447, 698–701 (2007). Nature, 451, E5–E6. Nothdurft, L. D., Webb, G. E., Bostrom, T., and Rintoul, L., 2007. Calcite-filled borings in the most recently deposited skeleton in live-collected Porites (Scleractinia): implications for trace element archives. Geochimica et Cosmochimica Acta, 71, 5423–5438. Nyberg, J., Malmgren, B. A., Winter, A., Jury, M. R., Kilbourne, K. H., and Quinn, T. M., 2007. Low Atlantic hurricane activity in the 1970s and 1980s compared to the past 270 years. Nature, 447, 698–701. Nyberg, J., Malmgren, B. A., Winter, A., Jury, M. R., Kilbourne, K. H., and Quinn, T. M., 2008. Nyberg et al. reply, Replying to: U. Neu Nature 451, doi: 10/1038/nature06576 (2008). Nature, 451, E6. Overpeck, J. T., and Cole, J. E., 2007. Lessons from a distant monsoon. Nature, 445, 270–271. Pfeiffer, M., and Dullo, W.-C., 2006. Monsoon-induced cooling of the western equatorial Indian Ocean as recorded in coral oxygen isotope records from the Seychelles covering the period of 1840–1994 AD. Quaternary Science Reviews, 25, 993–1009.

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Quinn, T. M., and Taylor, F. W., 2006. SST artifacts in coral proxy records produced by early marine diagenesis in a modern coral from Rabaul, Papua New Guinea. Geophysical Research Letters, 33, L04601. Reynolds, R. W., and Smith, M. T., 1994. Improved global surface temperature analysis. Journal of Climate, 7, 927–948. Ropelewski, C. F., and Halpert, M. S., 1987. Global and regional scale precipitation patterns associated with the El Niño/Southern Oscillation. Monthly Weather Review, 115, 1606–1626. Saenger, C., Cohen, A. L., Oppo, D. W., Halley, R. B., and Carilli, J. E., 2009. Surface-temperature trends and variability in the low-latitude North Atlantic since 1552. Nature Geoscience, 2, 492–495. Saji, H. H., Goswami, B. N., Vinayachandran, P. H., and Yamagata, T., 1999. A dipole mode in the tropical Indian Ocean. Nature, 401, 360–363. Stephans, C. L., Quinn, T. M., Taylor, F. W., and Corrége, T., 2004. Assessing the reproducibility of coral-based climate records: a multi-proxy replication test using multiple coral heads from New Caledonia. Geophysical Research Letters, 31, L18210. Sun, D., Gagan, M. K., Cheng, H., Scott-Gagan, H., Dykoski, C. A., Edwards, R. L., and Su, R., 2005. Seasonal and interannual variability of the Mid-Holocene East Asian monsoon in coral d18O records from the South China Sea. Earth and Planetary Science Letters, 237, 69–84. Thompson, D. W., and Wallace, J. M., 2001. Regional climate impacts of the Northern Hemisphere Annular Mode. Science, 293, 85–89. Tudhope, A., and Collins, M., 2003. The past and future of El Niño. Nature, 424, 261–262. Tudhope, A. W., Chilcott, C. P., McCulloch, M. T., Cook, E. R., Chappell, J., Ellam, R. M., Lea, D. W., Lough, J. M., and Shimmield, G. B., 2001. Variability in the El Niño-Southern Oscillation through a glacial-interglacial cycle. Science, 291, 1511–1517. Urban, F. E., Cole, J. E., and Overpeck, J. T., 2000. Influence of mean climate change on climate variability from a 155-year tropical Pacific coral record. Nature, 407, 989–993. Webster, J. M., Yokoyama, Y., and Cotterill, C., 2009. Great Barrier Reef environmental changes: the last deglacial sea level rise in the South Pacific: offshore drilling northeast Australia. IODP Scientific Prospectus, 325, doi:10.2204/iodp. sp.325.2009. Webster, P. J., 1994. The role of hydrological processes in oceanatmosphere interactions. Reviews of Geophysics, 32, www. iodp.org/scientific-publications/. Webster, P. J., Magaña, V. O., Palmer, T. N., Shukla, J., Tomas, R. A., Yanai, M., and Yasunari, T., 1998. Monsoons: processes, predictability, and the prospects for prediction. Journal of Geophysical Research, 103, 14451–14510. Wilson, R., Tudhope, A., Brohan, P., Briffa, K., Osborn, T., and Tett, S., 2006. Two-hundred-fifty years of reconstructed and modeled tropical temperatures. Journal of Geophysical Research, 111, C10007. Woodroffe, C. D., and Gagan, M. K., 2000. Coral microatolls from the central Pacific record late Holocene El Niño. Geophysical Research Letters, 27, 1511–1514. Woodroffe, C. D., Beech, M. R., and Gagan, M. K., 2003. Mid-late Holocene El Niño variability in the equatorial Pacific from coral microatolls. Geophysical Research Letters, 30, 1358, doi:10.1029/2002GL015868. Xie, P., and Arkin, P. A., 1996. Analyses of global monthly precipitation using gauge observations, satellite estimates, and numerical model predictions. Journal of Climate, 9, 840–858. Zhao, J.-X., Yu, K.-F., and Feng, Y.-X., 2009. High-precision 238 U-234U-230Th disequilibrium dating of the recent past: a review. Quaternary Geochronology, 4, 423–433.

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Cross-references Aragonite Calcite Corals: Biology, Skeletal Deposition, and Reef-Building Corals: Environmental Controls on Growth Diagenesis El Niño, La Niña, and ENSO Microatoll Mid Holocene Radiocarbon (14C): Dating and Corals Sclerochronology Stable Isotopes and Trace Elements Uranium Series Dating

PATCH REEFS: LIDAR MORPHOMETRIC ANALYSIS John C. Brock1, Monica Palaseanu-Lovejoy2 U.S. Geological Survey, Reston, VA, USA 2 U.S. Geological Survey, St. Petersburg, FL, USA

1

Setting Alina Reef is one of several thousand patch reefs that lie across the shallow carbonate platform seaward of Hawk Channel off the northern Florida Keys. The site is near the northern latitudinal fringe of the late Holocene western Atlantic coral reef distribution (Figure 1). The area is covered by calcareous sand and discontinuous Thalassia testudinum seagrass meadows and is studded with numerous scattered Holocene patch reefs. Most of the patch reefs are found in water depths of 2–9 m, are subcircular, elliptical, or irregular in plan view, and range up to about 8 m in vertical relief and 700 m in width. Coring has demonstrated thicknesses of 4.5–6 m and has revealed frameworks built by large, massive head corals. Lidar surveys In August 2002, the Experimental Advanced Airborne Research Lidar (EAARL) system was used by NASA and the U.S. Geological Survey to survey the submarine topography of the broad swath of the reef tract seaward of Elliot Key and within Biscayne National Park (Figure 1). The NASA – USGS Airborne Lidar Processing System (ALPS) was used to interpret the EAARL laser soundings to create a spot-elevation data set that was subsequently subjected to triangulation and gridding to create a digital-elevation model (DEM) at 1-m cell resolution. Next, a 1-m-resolution slope map was constructed from the lidar-derived DEM based upon

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Patch Reefs: Lidar Morphometric Analysis, Figure 1 The northern Florida Keys reef tract depicted on a QuickBird satellite image.

the average maximum method (Burrough and McDonell, 1998). Lastly, a lidar-rugosity mapping procedure was applied to produce a 1-m-resolution digital rugosity map for the study area (Brock et al., 2006).

Morphometric analysis The results of the lidar morphometric analysis at Alina Reef are shown in Figure 2. Alina Reef is depicted on the color-coded digital-elevation map (DEM) as the

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Patch Reefs: Lidar Morphometric Analysis, Figure 2 A subregion of the 1-m-resolution lidar-based (a) topographic, (b) slope, and (c) rugosity maps that were used to define patch-reef boundaries and morphometric attributes. Black polygons depict boundaries of the full patch reef, blue polygons delimit patch-reef tops, and the area between these polygons on a given reef represents the patch-reef rim.

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largest of several surrounding patch reefs at the approximate center of a small subarea of the total region mapped by the EAARL (Figure 2a). The black dots signify the volumetric center points of the patch reefs, one example of the morphological metrics that may be obtained through quantitative analysis of the lidar DEM. An extensive morphometric analysis across the entire EAARL survey region revealed two morphologically different populations of patch reefs associated with two distinct depth intervals. Alina Reef is an excellent example of an identified group of shallow patch reefs that are generally much broader and flatter than the deeper population. Compared to the shallow reefs, the deep reefs were smaller in area and volume and showed no trend in topographic complexity relative to water depth. The fundamentally different morphologies of the shallow and deep reefs can be clearly seen on the slope map, (Figure 2b) which depicts a small steep-sided deep reef immediately adjacent and to the east of Alina Reef. The slope map also reveals scattered small rubble piles just off the western flanks of several reefs that may stem from the frequent hurricane crossings that impact the northern reef tract.

The roughly 200 m  300 m region around Alina Reef mainly contains reefs that are members of the shallow-reef population, revealed by the lidar to be, in comparison to the deep-reef group, more variable in area and volume. Moreover, lidar revealed that reefs in the shallow population became flatter and less topographically complex with decreasing water depth. The more uniform knoll-like morphology of the deep reefs is consistent with steady and relatively rapidly rising early Holocene sea level that may have restricted the lateral growth of reefs. The morphology of shallow “pancake-shaped” reefs (Figure 3) at the highest platform elevations may have been produced by cycles of growth and erosion driven by fluctuating sea level during the late Holocene (Balsillie and Donoghue, 2004). Although the ultimate cause for the morphometric depth trends sensed by lidar remains speculative, these interpretations are compatible with a recent eustatic sea-level curve that hindcasts fluctuating late Holocene sea level (Siddall et al., 2003). It is thus suggested that the morphological differences revealed by lidar represent two stages of reef accretion that occurred during different sea-level conditions.

Patch Reefs: Lidar Morphometric Analysis, Figure 3 Oblique perspective topographic depiction of Alina Reef, a shallow patch reef that exhibits the characteristic ‘‘pancake’’ morphology of the shallow patch-reef population in the northern Florida Keys reef tract.

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Summary Apart from providing evidence for oscillating late Holocene sea level, lidar mapping provided new information on regional reef ecology and benthic community structure. Previous work has shown that many patch reefs in Biscayne Bay are composites of several characteristic zones, namely (a) massive boulder-coral colony clusters, (b) a zone of scattered massive corals with abundant octocorals, and (c) peripheral rubble zones (Ginsburg et al., 2001). The cluster zones occur at the patch-reef margins and are mainly composed of 1–5 m-diameter coalesced colonies of Montastraea annularis and Montastraea faveolata, along with fewer Siderastrea siderea and Colpophyllia natans heads. The octocoral zone is basically a hard ground community with areas that are dominated by coral rubble and carbonate sand. The tropical storms and hurricanes that frequently track across the northern reef tract have been implicated in the formation of the rubble aprons that occur at the peripheries of many patch reefs and can readily be identified on the lidar-based slope map (Brock et al., 2006) (Figure 2b). Analysis of dense lidar-elevation soundings can also yield spatially continuous maps of coral reef topographic complexity, or rugosity. A lidar-based rugosity map (Figure 2c) for the Alina Reef region confirms previous work that established that the known characteristic zonation of Biscayne Bay patch reefs can be sensed remotely and mapped by the EAARL system (Brock et al., 2006). The high-rugosity flanks of Alina Reef and neighboring reefs generally correspond to cluster zones of live or degraded massive boulder-coral colony clusters, or derivative rubble piles. The mixed, low-rugosity upper platforms of these reefs were identified by field observations to represent octocoral zones. In summary, the EAARL lidar reef-mapping test conducted across a portion of the northern Florida Keys reef tract, and the sample presented here, demonstrated that lidar mapping can provide new and valuable information to reef-resource managers, carbonate geologists, and coral reef ecologists. Bibliography Balsillie, J. H., and Donoghue, J. F., 2004. High resolution sea-level history for the Gulf of Mexico since the last glacial maximum. Report of Investigations Number 103. Tallahassee: Florida Geological Survey. Brock J. C., Wright, C. W., Kuffner, I. B., Hernandez, R., and Thompson, P., 2006. Airborne lidar sensing of massive stony coral colonies on patch reefs in the northern Florida reef tract. Remote Sensing of Environment, 104, 31–42. Burrough, P. A., and McDonell, R. A., 1998. Principles of Geographical Information Systems. New York: Oxford University Press. Ginsburg, R. N., Gischler, E., and Kiene, W. E., 2001. Partial mortality of massive reef-building corals: An index of patch-reef condition, Florida reef tract. Bulletin of Marine Science, 69, 1149–1173. Siddall, M., Rhling, E. J., Almogi-Labin, A., Hemleben, C., Meischner, D, Scmelzer, I, and Smeed, D. A., 2003. Sea-level fluctuations during the last glacial cycle. Nature, 423, 853–858.

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Cross-references Florida Keys Postglacial Trangression Remote Sensing Swathe Mapping

PERMIAN CAPITAN REEF SYSTEM Rachel Wood University of Edinburgh, Edinburgh, UK The Permian Capitan Reef system (275–255 Myr), as exposed in the Guadalupe and Delaware Mountains of the Delaware Basin in New Mexico and west Texas, presents some of the finest outcrops of ancient reef and associated rocks known. Ochoan evaporites, infilled the Delaware Basin preventing extensive erosional modification and structural deformation, so preserving an entire depositional system from the back-reef across the reef margin to the deep basin. The Delaware Basin was situated close to the western margin of Pangaea, about 10
north of the palaeo-equator and was subjected to subsidence resulting in the accumulation of between 2,100 and 4,200 m of clastic, carbonate, and evaporite strata; 1,000 m alone was deposited during Guadalupian times. The depositional system can be broadly divided into back-reef, shelf margin, reef-slope and basin sediments. Leonardian and early Guadalupian strata show shelf-to basin ramp sequences, but these later evolved into a rimmed shelf where the margin was poised nearly 600 m above the basin floor. Topography was further enhanced by reciprocal sedimentation where considerable thicknesses of carbonates were deposited on the shelf area during transgressions and highstands, and clastic material accumulated in the basins during lowstands. Cyclic sedimentation is widespread in the Permian Basin and operated on many frequencies. Eustatic cycles of roughly a million years or less were superimposed on an apparently long-term drop in sea level throughout late Permian time. The Capitan reef shows a high biodiversity and is the main Guadalupian carbonate-producing facies, forming a clearly defined, largely continuous margin which surrounds the Delaware Basin. Many different reef-building communities have been documented, including various calcareous sponge-bryozoan assemblages, and phylloid algae (Newell et al., 1953). Some have re-interpreted the previously described reef framework as having lived within cavities (Wood et al., 1996). Many organisms, commonly still found in living position, are often encrusted by possible red or blue-green algae, particularly Archaeolithoporella as well as microbialite, which almost certainly represents secondary growth during burial (Figure 1). In addition to framework and encrusting

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Permian Capitan Reef System, Figure 1 Reconstruction of bryozoan-sponge community, Permian Capitan Reef, USA (courtesy J. Sibbick); 1. Frondose bryozoans, 2. Solitary sphinctozoan sponges, 3. Archaeolithoporella (encrusting alga), 4. Automicrite, 5. Synsedimentary cement botryoids, 6. Sediment.

organisms, the reefs contain diverse echinoderms, bryozoans, brachiopods, molluscs, ostracodes, scarce solitary corals and trilobites. On the shelfward side of the reef, sponge-algal rubble passes into Tubiphytes thickets, and Mizzia and Macroporella green algal grainstones with belerophont gastropods. Late Guadalupian reefs and near-back-reef grainstones are filled with large volumes of syndepositional marine cements (Mazzullo and Cys, 1977). Many reef cavities were rapidly filled with massive botryoids and radial-fibrous cements, partly intergrown with Archaeolithoporella and internal sediment. High rates of biological productivity produced more carbonate than could be accommodated given limited rates of subsidence and, excess material was transported into back-reef and particularly fore-reef environments. As a result, the Capitan reef shows considerable seaward progradation, varying from 10 km on the stable western margin of the Delaware Basin to 3 km on the more rapidly subsiding eastern margin (Ward et al., 1986). The rapid rates of synsedimentary cementation of the reef facies combined with progradation over largely unconsolidated and compactable debris, led to extensive syndepositional fracturing of the cemented reef slabs and differential subsidence of the reef complex as a whole (Saller, 1996).

Midland, TX: Permian Basin Section-SEPM Publication 77–16, pp. 151–200. Newell, N. D., Rigby, J. K., Fischer, A. G., Whiteman, A. J., Hickox, J. E., and Bradley, J. S., 1953. The Permian Reef Complex of the Guadalupe Mountains Region, Texas and New Mexico. San Francisco, CA: W.H. Freeman and Co., 236 pp. Saller, A. H., 1996. Differential compaction and basinward tilting of the prograding Capitan reef complex, Permian, west Texas and southeast New Mexico, U.S.A. Sedimentary Geology, 101, 21–30. Ward, R. F., Kendall, C. G., St. C., and Harris, P. M., 1986. Upper Permian (Guadalupian) facies and their association with hydrocarbons, Permian basin, west Texas and New Mexico. American Association of Petroleum Geologists Bulletin, 70, 239–262. Wood, R. A., Dickson, J. A. D., and Kirkland, B. L., 1996. New observations on the ecology of the Permian Capitan reef, Guadalupe Mountains, Texas and New Mexico. Palaeontology, 39, 733–762.

Bibliography

Definition The Persian/Arabian Gulf (herein referred to as the “Gulf ”) is a subtropical, epicontinental sea situated roughly between 23
500 N and 29
520 N and is therewith home to some of the northernmost coral reefs on the western boundary of the Indo-Pacific (Kharku Island, Iran, at

Mazzullo, S. J., and Cys, J. M., 1977. Submarine cements in Permian boundstones and reef-associated rocks, Guadalupe Mountains, west Texas and southeastern New Mexico. In Hileman, M. E., and Mazzullo, S. J. (eds.), Upper Guadalupian Facies, Permian Reef Complex, Guadalupe Mountains, New Mexico and West Texas (1977 Field Conference Guidebook):

Cross-references General Evolution of Carbonate Reefs

PERSIAN/ARABIAN GULF CORAL REEFS Bernhard Riegl, Samuel Purkis Nova Southeastern University Oceanographic Center, Dania, FL, USA

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29
200 N). The Gulf is 989 km long and 50 km wide and has an area of 251,000 km2. It has an average depth of 20 m and a maximum of 90 m. Its shallow nature, combined with the high-latitude geographical position and the presence of mountainous plateaus and deserts nearby, make the Gulf’s climate the most extreme endured by reef-building corals anywhere in the world. Nonetheless, the Gulf is home to about 40 species of scleractinian and 31 species of alcyonacean corals. The reef fauna represents a small but typical segment of that of the Indo-Pacific.

The distribution of coral reefs in the Gulf Coral growth occurs virtually throughout the entire Gulf (Figure 1), with best development on offshore shoals but important fringing systems even along the mainland shoreline (in particular, UAE, Qatar, Saudi Arabia). Relatively few records exist of coral assemblages on the Iranian mainland coast (Shokri et al., 2000; Maghsoudlou et al., 2008), which can be due to runoff from the mountainous hinterland creating unfavorable conditions as well as insufficient records in the literature available in the West. Pleistocene (MIS 7 and 5e) reefs are known from the Iranian islands of Kish (Preusser et al., 2003) and Qeshm (Pirazzoli et al., 2004). The Holocene transgression is relevant to understanding the present distribution of coral reefs in the region. Flooding initiated approximately 12 kybp (=thousands of years before present) and sea level rose rapidly until

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9 kybp from 50%) occurred in the centre of the islands. Vegetation was cleared, temporary buildings constructed, and even tramways and wharfs built on the most remote of islands. Environmental impacts were devastating and in some locations recovery has still not taken place (Hopley, 1988, 1989; Daley and Griggs, 2006). Pisonia forest has rarely returned but weeds, introduced during the period of exploitation, may now be a permanent feature of the flora. In fact, guano mining is probably the most lasting of all the anthropogenic impacts on GBR islands, having taken place on nine reef islands and one high island (Holbourne), which has an area of Holocene carbonate sediments, similar to the older cays including three

Phosphatic Cay Sandstone, Figure 2 The central guano flat on Raine Island still largely bare of vegetation 100 years after the end of phosphate mining. The central flat originally had a guano deposit about 1.5 m thick, overlain by a black soil.

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phosphatized beach ridge terraces (Hopley, 1975), which were mined between 1918 and 1921 (Daley and Griggs, 2006). The degraded quarry remains as a reminder of the impact of mining on the physical environment. However, on Raine Island and elsewhere on the GBR, the small decline in seabird roosting and nesting appears to be more recent than the time of phosphate mining (1890–1892). Batianoff and Cornelus (2005) relate this to climatic and oceanographic changes in surrounding waters rather than devastation caused by the mining (Figure 2). Sixteen species of birds (some uncommon or rare) still nest on Raine Island, which is also the most important nesting site of the green turtle Chelonia mydas in the world (Limpus et al., 2003).

References Baker, J. C., Jell, J. S., Hacker, J. L. F., and Baublys, K. A., 1998. Origin of recent insular phosphate rock on a coral cay – Raine Island, northern Great Barrier Reef, Australia. Journal of Sedimentary Research, 68, 1001–1008. Batianoff, G. N., and Cornelius, N. J., 2005. Birds of Raine Island: population trends, breeding behavior and nesting habitats. Proceedings of the Royal Society Queensland, 112, 129. Chen, D., and Krol, A., 1997. Hydrogeology of Heron Island, Great Barrier Reef, Australia. In Vacher, H. L., and Quinn, T. M. (eds.), Geology and Hydrogeology of Carbonate Islands. Developments in Sedimentology, 54, 867–884. Chivas, A., Chappell, J., Polach, H., Pillans, B., and Flood, P., 1986. Radio-carbon evidence for the timing and rate of island development, beach rock formation and phosphatization at Lady Elliott Island, Queensland, Australia. Marine Geology, 69, 273–287. Daley, B., and Griggs, P., 2006. Mining the reefs and cays: coral, guano and rock phosphate extraction in the Great Barrier Reef, Australia, 1844–1940. Environment and History, 12, 395–433. Fosberg, F. R., 1954. Soils of the Northern Marshall Atolls, with special reference to the Jemo Soils. Soil Sciences, 78, 99–107. Fosberg, F. R., 1994. Comments on atoll phosphate rock. Atoll Research Bulletin, 396, 1–5. Hopley, D., 1975. Contrasting evidence for Holocene sea levels with special reference to the Bowen-Whitsunday area of Queensland. In Douglas, J., Hobbs, J. E., and Pigram, J. J. (eds.), Geographical Essays in Honour of Gilbert, J. Butland. NSW, Australia: University of New England, 51–84. Hopley, D., 1988. Anthropogenic influences on Australia’s Great Barrier Reef. Australian Geographer, 19, 26–45. Hopley, D., 1989. The Great Barrier Reef: Ecology and Management. Melbourne, Australia: Longman Cheshire, 54 pp. Hopley, D., Smithers, S. G., and Parnell, K. E., 2007. The Geomorphology of the Great Barrier Reef: Development, Diversity and Change. Cambridge: Cambridge University Press, 532 pp. Limpus, C. J., Miller, J. D., Parmenter, E. J., and Limpus, D. J., 2003. The green turtle, Chelonia mydas, population of Raine Island and the northern GBR: 1843–2001. Memoirs of the Queensland Museum, 49(1), 349–440. Miller, M. W., Halley, R. B., and Gleason, A. C. R., 2008. Reef geology and biology of Navassa Island. In Riegl, R. M., and Dodge, R. E. (eds.), Coral Reefs of the USA. Coral Reefs of the World. The Netherlands: Springer, Vol. 1, pp. 407–433. Rodgers, K. A., 1992. Occurrence of phosphate rock and associated soils in Tuvalu, Central Pacific. Atoll Research Bulletin, 360, 31 pp. Rougerie, F., Fichez, R., and Déjardin, P., 1997. Geomorphology and hydrogeology of selected islands of French Polynesia: Tikehau (Atoll) and Tahiti (Barrier Reef ). In Vacher, H. L., and

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Quinn, T. (eds.), Geology and Hydrogeology of Carbonate Islands. Developments in Sedimentology, 54, 475–502. Stoddart, D. R., and Scoffin, T. P., 1983. Phosphate rock on coral reef islands. In Goudie, A. G., and Pye, K. (eds.), Chemical Sediments and Geomorphology. London: Academic, pp. 369–400. Stoddart, D. R., Gibbs, P. E., and Hopley, D., 1981. Natural history of Raine Island, Great Barrier Reef. Atoll Research Bulletin, 254, 70 pp.

Cross-references Beach Rock Coral Cay Classification and Evolution Coral Cays-Geohydrology Makatea Mining/Quarrying of Coral Reefs Nutrient Pollution/Eutrophication Palaeosols Soils of Low Elevation Coral Structures

PLATE TECTONICS Paul Wessel University of Hawaii at Manoa, Honolulu, HI, USA

Synonyms Continental drift Definition Absolute plate motions (APM): Motion of tectonic plates relative to a fixed point in the mantle, typically defined by one or more hot spots (i.e., the hotspot reference frame). Apparent polar wander (APW): Motion of the north pole as seen from a moving tectonic plate. Relative plate motions (RPM): Motion of tectonic plates relative to each other, typically inferred from marine magnetic anomalies and the geomagnetic timescale. True polar wander (TPW): Rotation of the entire solid Earth leading to a realignment of the rotation (and geomagnetic) axis. Introduction Reefs build on shallow-water geologic foundations (e.g., continental shelves, tectonically active oceanic islands) that slowly undergo horizontal (due to plate tectonics) and vertical (thermal and flexural subsidence) movements over time. As plates move laterally, coral-rich areas may move to latitudes where coral growth is inhibited (e.g., Grigg and Epp, 1989). Furthermore, long-term plate tectonic changes in the configuration of landmasses can alter the global oceanic circulation, thus affecting the colonization of coral (e.g., Grigg, 1997). Understanding absolute plate motions is thus important for understanding long-term coral variability and the history of global biodiversity patterns (e.g., Renema et al., 2008). The main purpose of this article is therefore to

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give an overview of absolute plate motions, the assumptions used, and the current status of plate tectonic research. Plate tectonics, the major paradigm for how the Earth works, was established in the 1960s following decades of observational research that culminated in the key discoveries of geomagnetic reversals, mid-ocean ridges, transform faults, and seafloor spreading. Collectively, these insights gave rise to the “new global tectonics” or theory of plate tectonics (e.g., Wessel and Müller, 2007). The history of RPM is inferred from the mapping of marine magnetic anomalies and seafloor morphology using implementations of Euler’s theorem for rotations on a sphere. Using improved observations of seafloor morphology derived from multibeam bathymetry and satellite gravity, combined with better coverage of magnetic anomalies, has allowed researchers to improve the resolution of RPM models. The global crustal age models derived from a worldwide compilation of RPM studies (Müller et al., 2008) perhaps best represent the state of the art (Figure 1). However, RPMs do not let scientists examine changes in latitude, and thus an absolute reference frame is needed. APMs typically refer to motion relative to a fixed mantle, and the most prominent mantle reference that has been proposed is the “fixed hotspot” reference frame. The hotspot reference frame is derived from the geometries of dated hotspot islands and seamount chains

Plate Tectonics, Figure 1 Relative motion between plates as evidenced by marine magnetic anomalies. Combined with a geomagnetic reversal timescale, one can create crustal age grids (Mu¨ller et al., 2008). The relative motions between plates are very well determined, but it is the absolute motions of plates relative to a fixed point in the mantle that let us examine any changes in latitude that may affect reef colonization.

(e.g., Duncan and Clague, 1985) and was introduced by the plate tectonics pioneers Wilson (1963) and Morgan (1971). The foundation of this absolute reference frame, that is, the fixed hot spots, has been the focus of much research during the last few years (e.g., Koppers et al., 2001; Steinberger and Torsvik, 2008; Wessel and Kroenke, 2008). The Hawaii-Emperor chain and its famous bend have been interpreted as the prime surface manifestations of the changing Pacific plate motions over a stationary mantle plume; in fact, this interpretation has reached textbook status. However, over the last decade, a more complicated picture has emerged. Paleolatitudes of the Emperor seamount chain have been inferred from the frozen-in paleomagnetic field, and these imply a birthplace significantly farther north than the present location of the Hawaiian hot spot. The most logical conclusion is that the Hawaiian plume must have been further north in the past and subsequently drifted south (Tarduno et al., 2003). The extent to which such drift contributed to the prominent Hawaii-Emperor bend, now believed to have formed 47–50 Ma (Chron 21–22), is the topic of ongoing research and has the potential to significantly revise established models of APMs.

Plate tectonics Seafloor spreading, a key component of plate tectonics, was developed in the early 1960s (Dietz, 1961; Hess, 1962) when it was first realized that new seafloor is created at mid-ocean ridges, spreads away from them as it ages, and is recycled at subduction zones. However, earlier and remarkably similar ideas had been suggested by Holmes (1944). A key insight came from Wilson (1965), who introduced the concepts of plates and transform faults. Wilson argued that the Earth’s active mobile belts were continuous, marked by active seismicity, and divided the Earth into a set of approximately a dozen rigid plates. These active mobile belts consist of ridges where plates are created, trenches where plates are destroyed, and transform faults that link the other two belts. Plate tectonics theory states that the Earth’s surface is an interlocking, internally rigid set of plates in constant relative motion. Plates are rigid except at their boundaries, which are weak lines between contiguous plates. Earthquakes occur due to the relative motion between plates; in fact, the earthquakes define the plate boundaries (Figure 2). Three types of plate boundaries are recognized: (1) divergent boundaries – where new crust is produced, (2) convergent boundaries – where crust is recycled as one plate subducts beneath another, and (3) transform boundaries – where crust is preserved as two plates slide horizontally past each other. Complications can arise in some areas. For instance, in places where the boundaries are not well defined, the deformation tends to extend over a broad plate-boundary zone (Gordon, 2000). Such plate-boundary zones tend to have complex geology and deformation, possibly involving one or more microplates.

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Plate Tectonics, Figure 2 Map showing idealized narrow plate boundaries (white lines, Coffin et al., 1997), shallow (0–70 km depth) seismicity (red dots, Engdahl et al., 1998), and regions of diffuse plate boundaries (areas of lighter illumination, Gordon, 2000). Plate abbreviations: B Borneo, AN Antarctica, AR Arabia, AU Australia, CA Caribbean, CAP Capricorn, CL Caroline, CO Cocos, EU Eurasia, I Indo-China, JF Juan de Fuca, NA North America, NB Nubia, NC North China, NZ Nazca, OK Okhotsk, PA Pacific, SA South America, SC Scotia Sea, SM Somalia, Y Yangtze, T Tarim Basin, PH Philippine.

The morphology of the ocean floor Wilson (1963) realized that some regions, such as Hawaii, have been sites of long-term volcanic activity. Such focused volcanism would seem to imply that relatively small, stable, and anomalously hot regions (hot spots) exist beneath the plates, supplying high heat energy (i.e., a thermal plume) to maintain the surface volcanism. Wilson suggested that the alignment of the Hawaiian Islands-Emperor seamounts chain must have resulted from the Pacific plate passing over a fixed hot spot deep in the mantle, now located beneath the Island of Hawaii. Morgan (1971) further elaborated on this idea and developed the fixed hotspot hypothesis. About 100 hot spots have been proposed to be active during the past 10 million years (e.g., Burke and Wilson, 1976). The majority are located beneath plate interiors (e.g., the African plate), but others can be found near a mid-oceanic ridge system, such as the hot spot beneath Iceland, the Azores, and the Galapagos Islands. The standard way of reconstructing plates relative to a fixed mantle relies on linear chains of volcanoes that exhibit a monotonic age progression and can be traced back to singular spots of melting in the upper mantle (see Oceanic Hotspots). Mapping seamounts thus became an important prerequisite to deciphering APMs. Even prior to Wilson’s realization, it was well known that the seafloor was littered with underwater volcanoes

known as seamounts. During the war, Hess (1946) found and mapped several flat-topped seamounts that he named guyots; these seemed to be former islands that had been worn down to sea level and eventually had drowned. After World War II, a flurry of ocean-going expeditions equipped with echo-sounding instruments heralded major advances in the mapping of seafloor morphology, including the discovery that the oceans possessed mountain chains of extraordinary lengths. These facts were popularized by the famous physiographic maps of Lamont’s Heezen and Tharp (1961, 1964), rendering ubiquitous abyssal hills and majestic seamounts. Surveys of seamounts in the Pacific suggested that there might be as many as 100,000 seamounts exceeding a height of 1 km in the Pacific alone (Menard, 1964). Nevertheless, given the vastness of the oceans and the sparse sampling provided by surface ships, the large majority of seamounts were never charted. Statistical analyses of seamount populations were limited to extrapolations from smaller, well-surveyed areas by dense, single-beam echo-sounding tracks or by the improved coverage of the newer multi-beam systems in the early 1980s (e.g., Smith and Jordan, 1988). Space-age technology brought satellite altimetry as a new technique for the study of oceanography and, indirectly, seamounts. The global gravity maps developed from Seasat altimeter data portrayed a previously unseen

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level of detail of seafloor fabric, including numerous underwater volcanoes (Haxby, 1987; Haxby et al., 1983). Using along-track profiles of sea surface heights from the Seasat mission, Craig and Sandwell (1988) pioneered a global investigation into the distribution of seamounts, discovering 8,500 individual edifices throughout the world’s oceans. Key findings were that the Pacific Ocean basin contains most of the seamounts and that in particular the Western Pacific displays an unusual high density of large seamounts. Unfortunately, the large track spacing (100 km) meant that many seamounts were only partially surveyed and still more were left undiscovered in the gaps between tracks. The later Geosat/ERS-1 altimetry missions greatly improved the coverage by having closer track spacing (a few kilometers) and resulted in much improved global gravity grids (Sandwell and Smith, 1997). Using such grids, Wessel and Lyons (1997) characterized 8,900 seamounts in the Pacific Ocean alone, with the majority of them being smaller, hitherto uncharted seamounts. Wessel (2001) expanded the analysis globally, finding almost 14,700 seamounts, now adjusted down to 11,800 due to duplicates (Wessel et al., 2010). The size-frequency distribution suggests that sizes of seamounts appears to follow a power-law relationship, which can be extrapolated to match Menard’s original prediction of 100,000 seamounts of at least 1 km in height (Figure 3). Scientists 60°E

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now routinely study even smaller seamounts (e.g., 50– 100 m tall), and should the inferred frequency-size relationship hold for such small seamounts one would predict a global population into the millions. If valid, this would make these morphological features second in ubiquity only to the abyssal hills (e.g., Macdonald et al., 1996).

The ages of seamounts and oceanic islands The geometries of seamount chains place strong constraint on plate motions. However, the hotspot hypothesis predicts there should be a monotonic age progression along seamount and island chains formed by hotspot volcanism. Hence, determining the ages from rock samples collected at these sites became another key undertaking. Dredging of seamount flanks and sampling from oceanic islands built up a valuable data set whose importance cannot be overstated (Duncan, 1981; Duncan et al., 1985; McDougall, 1971; Turner and Jarrard, 1982; Turner et al., 1980). Recently, Clouard and Bonneville (2005) reviewed a compilation of more than 1,500 individual age-determinations from 300 different Pacific volcanoes. The available data seem to suggest that monotonic age progressions are indeed observed along several of the Pacific island and seamount chains, but for others the pattern is less clear (e.g., Dickinson, 1998; McNutt et al., 1997). The key difficulty in assessing age-progressive 120°W

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Plate Tectonics, Figure 3 Distribution of seamounts inferred by satellite altimetry (Wessel, 2001). Colors reflect estimated seamount sizes (blue are 1–2 km, green are 1.5–4 km, and red are >3.5 km tal). Most seamounts are found in the Pacific basin, with significant populations in both the Atlantic and Indian oceans. Hot spots (yellow stars) are often found near the young end of chains. Large igneous provinces (LIPs; orange) are often associated with seamount provinces. Black arrows indicate current absolute plate motions (for scale, the Pacific plate arrow length represents a velocity of 10.8 cm/year).

PLATE TECTONICS

volcanism is the lack of available samples. Furthermore, as the field of radiochronology has advanced, there is a great need to reanalyze older samples. Finally, the discovery that volcanism at a given site can remain active over an extended period (e.g., Pringle et al., 1991) makes it problematic to determine if the observed age represents the feature as a whole. For instance, in some areas where new samples have been obtained and the latest dating techniques have been applied, significant discrepancies between the new and old dates were found (e.g., Koppers et al., 2004). While such reanalysis is necessary, one must be careful when combining old and new ages in ageprogression studies. Surprisingly, some Pacific chains that have geometries similar to the Hawaii-Emperor bend have recently been dated, yielding ages that are incompatible with a hotspot origin (Koppers and Staudigel, 2005). Perhaps the volcanism was instead triggered by transient stresses during times of plate motion changes, leading to decompressional melting that rises up through preexisting zones of weakness (e.g., Sager and Keating, 1984; Wessel and Kroenke, 2007). Such complications continue to make studies of age-progressive volcanism challenging. Outside the Pacific, the situation is similar: Apparent age progressions are often found for long seamounts chains, but unfortunately both the quality and quantity of samples are much lower that for the Pacific (e.g., Baksi, 1999). It is clear that recent advances in dating techniques promise to raise the quality of radiometrically determined ages to a higher level (e.g., Koppers et al., 2004). However, this will take time, as reanalysis is a time-consuming and costly endeavor. Furthermore, the acquisition of new samples, in particular for older chains, is also a difficult and expensive undertaking and may ultimately require ocean drilling for reliable results (Tarduno et al., 2003). Therefore, the study of APMs is a topic constantly in need of more and better data.

Studies of absolute plate motions In recent decades, our understanding of the Mesozoic and Cenozoic RPM between the major tectonic plates has improved considerably, even though APM relative to a “fixed” underlying mantle is still controversial. Marine magnetic anomalies, some as old as 165 million years, along with fossil seafloor fabric based on bathymetry and altimetry clearly constrain RPMs for most of the major plates (Müller et al., 2008). Paleomagnetic data and hotspot traces are among the concepts that have been used to attempt to constrain APM. Paleomagnetic data may provide the paleo-meridian orientation and paleolatitude of a plate; together these determine the paleopole for a given plate. However, because the Earth’s dipole field is axisymmetric, no paleo-longitudinal information is obtained from paleomagnetic data. Hotspot tracks with linear age progressions can be used to restore plates to their original positions by assuming that hot spots are approximately fixed relative to each other (i.e., the “fixed hotspot hypothesis”). However, an early indication

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of problems to come appeared when Molnar and Stock (1987) showed that during the Tertiary, the Hawaiian hot spot apparently had moved at 10–20 mm/year relative to the Iceland hot spot and to hot spots beneath the African and Indian plates. Recently, new analyzes have questioned the magnitude and significance of this conclusion (Andrews et al., 2006). Numerous paleomagnetic data sets and models for the APM of the North American, African, and Eurasian plates during the Mesozoic and Cenozoic have been published (e.g., Engebretson et al., 1985). Unfortunately, there are inconsistencies between different apparent polar wander paths and hotspot models. Especially for times prior to 47 Ma, large misfits are found between APM models based on hotspot tracks and those based on apparent polar wander paths from paleomagnetic data (e.g., Sager, 2007). The key plates whose APM forms the basis for our global understanding are the African and Pacific plates.

Motion of the African plate Assuming fixed hot spots, Müller et al. (1993) proposed a key African APM model based on a refined model for global RPMs, bathymetry, and radiometric age dates of major hotspot tracks. They combined the major hotspot tracks with observed age progressions from the Atlantic and Indian oceans and employed an interactive technique to derive a “best-fit” qualitative model for motions of the major plates in the Atlantic-Indian domain relative to those hotspot tracks that had clear age progression. Despite the popularity of this model, it had some wellrecognized limitations. The Late Tertiary portion of this model was largely unconstrained by radiometric ages, based on the lack of published age dates. Before 80 Ma, the only hotspot tracks with age progression in the Atlantic-Indian oceans are the New England seamount chain (tied to the Great Meteor hot spot on the Africa plate) and the Walvis Ridge/Rio Grande Rise (likewise linked to the Tristan da Cunha hot spot) in the Atlantic Ocean (Duncan, 1984; O’Connor and Duncan, 1990). Consequently, the APM of the Indian, Australian, and Antarctic plates must be computed by plate circuit closure for these times. When the Müller et al. (1993) model was constructed, pre-80 Ma RPMs in the Indian Ocean were poorly known, due to a lack of data in crucial areas, especially offshore Antarctica in the Enderby Basin and areas south of the Kerguelen Plateau. There, a sequence of Mesozoic magnetic anomalies was subsequently mapped and modeled, originating at 130 Ma (Gaina et al., 2007; Gaina et al., 2003). However, the Müller et al. (1993) model (incorrectly) assumed a post-120 Ma breakup between India and Madagascar. Disagreements between hot spot and published paleomagnetic reference frames have surfaced for India (Müller et al., 1994) and Australia (Idnurm, 1985), suggesting that the mantle beneath the Indian Ocean cannot be considered a fixed reference frame. This mismatch directly affects estimates of latitude. For instance, paleopoles for India from the

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Rajmahal Traps (Das et al., 1996; Rao and Rao, 1996) result in a 47
S paleolatitude of the traps at their time of formation (117 Ma) (Baksi, 1995), whereas the Müller et al. (1993) model places them at about 40
S; both estimates have an uncertainty of 400 km. Comparing mid-Cretaceous (122–80 Ma) paleolatitudes of North America and Africa with estimates from hotspot tracks (Van Fossen and Kent, 1992) indicates an 11–13
misfit. This discrepancy suggests that Atlantic hot spots were not fixed relative to the Earth’s spin axis before 80 Ma, but moved southwards by as much as 18
(Torsvik et al., 2002) during 130–100 Ma. Others have argued that this apparent southward movement was caused by true polar wander (TPW) (Prevot et al., 2002), but this has been refuted by Tarduno and Smirnov (2002). The Müller et al. (1993) APM model predicts a relatively sharp bend of plate motion directions (e.g., of Australia and Antarctica) at about 80 Ma. This property originates from the bend between the New England seamount chain and the Corner seamounts at roughly 80 Ma in the central North Atlantic. An equivalent bend in fracture zones is not found in either the Atlantic or Indian Oceans. Revisiting the Africa APM, O’Neill et al. (2005) allowed for moving hot spots in order to satisfy both paleomagnetic constraints and hotspot chain geometry and age progression. They found evidence for plume motion on the order of 5–10
prior to 80 Ma; however since that period any plume motion is less than the model uncertainties. These improvements place India farther north in the past than what conventional fixed hotspot models do.

Motion of the Pacific plate The Hawaiian chain is clearly the most studied hotspot chain and has as its most characteristic feature the prominent bend near longitude 172
E, here called the HawaiiEmperor Bend (HEB; see Oceanic Hotspots). Since Wilson (1963) first suggested it, the HEB has been explained by a 60
change in direction of Pacific plate motion over a stationary hot spot in the mantle (Morgan, 1971). In fact, the geometries of the Hawaii-Emperor chain and many coeval chains (Figure 4) have been used to model the Pacific APM (e.g., Duncan and Clague, 1985; Koppers et al., 2001; Wessel et al., 2006; Wessel and Kroenke, 2008; Yan and Kroenke, 1993). A puzzle, however, was represented by the age of the bend, initially dated to 43 Ma (Dalrymple et al., 1987). Scientists expected that such a large reorientation of the motion of the dominant Pacific plate should have left abundant evidence of contemporaneous tectonic and magmatic events along the plate boundary; however, careful assessments of the geologic record failed to uncover such evidence (Atwater, 1989; Norton, 1995). Furthermore, rock samples obtained from drilling the Emperor seamounts revealed a frozen-in paleomagnetic field best explained if the seamounts had formed significantly further north (5–10
) of the present-day location of Hawaii (Kono,

1980; Tarduno and Cottrell, 1997). Subsequent efforts to project the APM of Africa, via the global plate circuit, into the Pacific failed to reproduce the shape of the HEB (Cande et al., 1995; Raymond et al., 2000). These inconsistencies lead to alternative models where the plume sustaining the volcanism was no longer stationary (Steinberger, 2000; Steinberger and O’Connell, 1998). Such models could fit both the changing latitude of the hot spot (as suggested by paleomagnetics) and the geometry and age progression of the seamount chain. Some researchers concluded that no change in Pacific APM had taken place at all: Since the trail records the vector sum of plate and plume motion, it could be reconstructed by a rapid slowdown in the southward motion of the plume while Pacific plate motion remained unchanged in direction and magnitude (Tarduno et al., 2003). To reconcile APM models inferred from the Indian, Atlantic, and the Pacific oceans requires their propagation via the global plate circuit (Acton and Gordon, 1994). Because the history of all RPM changes between conjugate plate pairs is not known, the projection of the African APM into the Pacific is subject to uncertainties that cannot easily be quantified. For instance, by choosing a different plate circuit for connecting the Pacific to Australia via the Lord Howe Rise, Steinberger et al. (2004) showed that the HEB did seem to require a plate motion component and thus could not only be caused by plume drift alone. Meanwhile, a reanalysis of the rock samples taken from around the HEB region revealed that the HEB is apparently much older than originally determined. Sharp and Clague (2006) found that the rocks around the bend erupted during the 47–49 Ma interval (Chron 21–22) and, by allowing for 1–2 million year construction time, that the bend itself could have formed even as early as 50 Ma – fully 7 million year earlier than previous estimates. This older age complicates many of the previous conclusions about the lack of correlation between the HEB and plate boundary processes, since Chron 21–22 is generally recognized as an active tectonic period in the Pacific and elsewhere (Cande et al., 1995; Rona and Richardson, 1978). Sharp and Clague (2006) examined details of the age progression and concluded it was unlikely that the plume was moving during the formation of the bend, thus favoring the original explanation (purely a plate motion change over a stationary hot spot) for the origin of the HEB.

Discussion APM studies continue to be carried out in the aftermath of these discoveries. For instance, Whittaker et al. (2007) presented new tectonic evidence for a change in plate motion between Australia and Antarctica; this kink appears to coincide in time with the HEB. This change in plate motions may further affect the global plate circuit and results that are derived from its use. Recently, the case for moving hot spots was reviewed by Tarduno et al. (2009). They argued that mantle-scale flow models could explain the types of plume drift

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Plate Tectonics, Figure 4 Map of the Pacific hemisphere with location of dated rock samples (triangles) from the Hawaii-Emperor (HI; red), Louisville (LV; blue), Foundation (FD; green), Pitcairn (PC; magenta), Caroline (CR; yellow), Cobb (CB; black), and other Pacific seamount chains (white). Red (blue) stars represent presently active (extinct) hotspot locations, while solid lines represent the WK08-G APM model geometry (Wessel and Kroenke, 2008); ellipses indicate 95% confidence regions. White stars represent active hotspot locations for chains not used to constrain the APM model. Heavy dashed lines represent the 15 inter-hotspot separations between the six chains identified. Dark red lines are plate boundaries (redrawn from Wessel and Kroenke, 2009).

inferred from the Hawaiian plume by the observed paleolatitudes from the Emperor seamount chain. In fact, they suggested that hotspot tracks, rather than constraining the APM, may instead give insight into the patterns of past mantle convection. At the same time, the distances between same-age samples in the HawaiiEmperor and Louisville chains were analyzed, suggesting relatively minor changes in the separation of these two hot spots through time (Wessel and Kroenke, 2009). In the critical time period 80–47 Ma, this separation decreased by perhaps 3–5
(Figure 5), compared to 7–9
predicted by moving hotspot models (Koppers et al., 2004; Steinberger and Gaina, 2007). There appears to be several ways

to satisfy both the quantified distance constraints and the observed paleolatitudes: (a) if the Hawaii plume was farther north at 90 Ma but moving southward (to satisfy the paleolatitudes), then the Louisville plume must have been much farther north as well (to match the observed hotspot separations). Given the variability in modeling results of the Louisville plume motion (e.g., Antretter et al., 2004), it is possible that a combination of boundary conditions and mantle viscosities will be found that will predict plume motions for Hawaii and Louisville that satisfy both distance and paleolatitude constraints; (b) perhaps plume drift only accounts for a part (e.g., 5
) of the observed paleolatitudes and the remainder should be attributed to

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Plate Tectonics, Figure 5 Examination of distance deviations, D(t): the distance between coeval points on different chains minus their present (fixed) hotspot separation. The polygons represent continuous estimates of D(t) with assigned uncertainties. Triangles show D(t) for locations of dated rock samples and their estimated coeval locations on other chains. Dashed line indicates typical prediction of D(t) for Hawaii and Louisville from mantle dynamic modeling (Steinberger and Gaina, 2007), with dotted line the (t) for the model of Koppers et al. (2004). Scatter due to age uncertainties (not analytic errors, but uncertainties associated with the timescale of volcano building) seems contained within a 2
window (Redrawn from Wessel and Kroenke, 2009).

TPW (e.g., Duncan and Richards, 1991); or finally (c) the interaction between the Hawaii hot spot and the KulaPacific ridge may have lead to the entrapment of plume melt by the ridge (Wessel and Kroenke, 2009). This interaction may have forced surface volcanism due to the plume to erupt closer to the ridge axis (instead of vertically above the plume), particularly after the ridge overrode the plume at 78 Ma. The overall effect would be to distort the oldest age-progression data but also result in more northerly paleolatitudes. While option (c) is not new (Sleep, 1996; Small, 1995; Tarduno et al., 2003) and could perhaps account for 4
of deviations (which is close to the 5
we observe), it remains a scenario that is difficult to test. Note, however, that if seamounts formed closer to the ridge than the hotspot location would imply, then the observed age progression would at first be slower but then accelerate as seamounts again begin to form vertically above the hot spot; of course, this scenario reflects the observed Emperor age progression (Sharp and Clague, 2006) and sense of paleomagnetic latitude shifts (Tarduno et al., 2003). Option (b) is a compromise model that allows for most of the HEB change to originate from a plate rather than plume motion change, as supported by contemporaneous global tectonic events (Rona and Richardson, 1978; Wessel and Kroenke, 2008). The implied small amount of TPW (32
C), but also very cold (15
C) waters such as in the northern Arabo-Persian Gulf. It also tolerates permanently high (e.g., Gulf of Tomini, Sulawesi, ca. 30
C) or cool waters, down to 21–22
C sea surface temperature (SST) mean annual SST (e.g., Rapa, Austral Islands). In many high latitude reefs, however, the genus is either absent (Kermadec Islands, 29
120 S, mean annual SST 21
C) or represented by only one species: P. lichen at Lord Howe Island (31
300 S, mean annual SST 21
C) or P. lobata at Easter Island (27
070 S, mean annual SST 22
C). It is to be noted that both high and low extremes of seawater temperatures can trigger a bleaching event (discoloration of the living tissues due to a breakdown of the symbiotic association), which in many instances is followed by a recovery of the colonies affected. Severe bleaching associated with significant SST anomalies, however, may lead to partial or even total mortality of the colonies concerned. Salinity Like many other coral genera, Porites species better tolerate high salinity than low salinity, which limits coral development in the vicinity of estuaries and river mouths. Occasional or accidental decrease in salinity associated with freshwater plumes of flooding coastal streams or rivers may cause partial or total bleaching of Porites, which may lead to partial or total mortality of the colonies. A number of species can thrive in areas where salinity is permanently and significantly higher than normal, and up to at least to 42%, such as the Red Sea and the Arabo-Persian Gulf. Many large massive or columnar species are common throughout the Red Sea (where they were frequently used as building material) and colonies of P. lutea and P. harrisoni up to several meters across are very common in the coral reefs of Kuwait. The combined effect of temperature and salinity has a synergistic effect on the level of stress and survival of corals, and

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Porites tolerance to extremes for both factors makes the genus a dominant feature in coral reefs developed in harsh environments.

Light In the natural environment, light intensity and quality is associated with other environmental factors such as depth, water transparency, and emersion. Light requirements are, to a large extent, species-specific, but various species of Porites can be found at both ends of the light intensity range. Species such as P. murrayensis, P. stephensoni, P. lutea, P. lobata, or P. somaliensis tolerate exposure to full sunlight at low tide on the reef flats for several hours. Conversely, several large massive species including, inter alia, P. lutea and P. lobata can occur on the deep fore reef slope. Sedimentation In addition to the various mechanisms that have been developed by scleractinian corals to eliminate fine particles which deposit on the colony surface (ciliary currents, tentacular movements), several species in the genus Porites produce abundant sheets of mucus, which covers and protects the living tissue, and on which the fine sediment particles get trapped. After some time, generally up to a few weeks, the sheet of mucus is eliminated, thus removing the sedimented particles, and the process may start again. Overall, Porites species cope well – within limits – with significant levels of fine particle sedimentation, hence their dominance in areas such as back reef slopes, fringing reefs, or deep embayments. Water motion Despite the fact that many of the species build very large massive colonies, which seemingly would resist strong currents or pounding by breaking swell and waves, large specimens of massive Porites are rarely found in areas of very high hydrodynamic energy such as the shallow reef front. This situation may be due to a combination of factors: on the one hand, the area of attachment of massive species on the substratum is much smaller than the large size of the colonies would suggest, and further the colony is often not very firmly adherent to the reef; on the other hand, in such environments, Porites is out competed by other (r-strategist) genera such as Acropora and Pocillopora, or even by crustose coralline algae. As a result, large Porites specimens are more commonly found in more protected reef zones such as the reef flat and the reef slopes. Local and regional variability in the abundance and size frequency distributions of massive Porites colonies also reflects variability in both the prevailing wave climate and the return period between cyclonic waves strong enough to dislodge corals of different sizes at different depths (Massel and Done, 1993). Emersion A number of Porites species found of the reef flat may be fully emerged during low spring tides and exposed to full

sun during daytime, for up to several hours. They include species restricted to that type of environment (P. murrayensis, P. stephensoni) but also species with a larger cross-reef distribution such as P. lutea, P. lobata, and P. australiensis. The ability of such species to withstand exposure to air, and in particular desiccation, stems from the fact that the three-dimensional structure of the skeleton acts as a wick, pumping water from residual pools of seawater or from water-logged sediment and bringing it to the soft tissue which are as a result kept sufficiently moist. Thus, Porites species are often, in terms of vertical distribution, some of the highest corals to be found on the top of the reef buildup.

Population biology Reproduction The majority of the Indo-Pacific species of Porites are gonochoric (separate sexes), spawning gametes, although occasionally a few hermaphroditic polyps or colonies are found, and they reproduce only once a year (Harrison and Wallace, 1990). In contrast, the Atlantic representatives of the genus are hermaphroditic and up to nine reproductive cycles have been recorded for P. astreoides (Szmant, 1986). This situation can be explained by the fact that the faunas of the two regions have been separated for more than 10 million years, hence there has been evolutionary divergence resulting in contrasting sexual patterns. Little is known of size/age at first reproduction. Harriott (1983) gives a colony diameter of less than 8 cm for gravid P. australiensis and P. lutea (which have massive growth forms) For the branching P. astreoides in Panama, first reproduction occurs when the average branch length reaches 8.4 cm, the smallest reproductive colony having a dimension of 3  2 cm (Soong and Lang, 1992). However, polyp fecundity in female colonies is correlated with branch thickness but not with age (Chornesky and Peters, 1987). For branching species in general, the tips of the branches are infertile (juvenile polyps) as are polyps at the base of the colonies. There are no data specifically for Porites species on the duration of the planktonic larval stage and larval dispersal. With respect to larval settlement, laboratory experiments have shown that for P. porites, larval settlement is patchy as a result of aggregated settlement behavior (Goreau et al., 1981). Like many other corals, growth rates in the first year after settlement are low and mortality very high. Quantitative data are scarce: Vaughan (1912) gives a figure of 4  3 mm after a year for the slowest growing Porites spat, with up to 14  23 mm for an agglomeration of six planulae. Among the various modes of asexual reproduction, breakage, survival, and subsequent regrowth of coral fragments are common occurrences for scleractinian corals (Highsmith, 1982). This reproduction strategy occurs in large, massive, Porites species (Done and Potts, 1992) but it is probably more frequent in columnar species and branching species such as P. cylindrica and P. nigrescens on the Great Barrier Reef (Resing and Ayre, 1985) and

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P. compressa in Hawaii (Hunter and Kehoe, 1986). At population level, Hunter (1985) presented evidence of predominantly clonal populations for P. compressa in Hawaii, but at a broader geographic scale the relative importance of asexual reproduction has not been fully assessed for most of the dominant, reef-building species.

Growth and growth rates Growth in the genus Porites, in particular retrospective growth in massive species, has been the subject of particular attention. This attention reflects their paramount importance as primary reef builders, and hence their role in coral reef development and maintenance. Growth in massive Porites is particularly amenable to a study by radiographic techniques (“density banding”), which rely on the seasonality in the density of calcium carbonate deposition. In most cases, yearly skeletal accretion includes a dark band (period of intense CaCO3 deposition) and a lighter band (period of lower CaCO3 deposition). For a general review of coral growth, refer to Buddemeier and Kinzie (1976). Radiographic analyses of density banding in cores obtained from large massive Porites show that:  Many large colonies are more than 100 years old, and some even reach an age of 700 years or more.  Average growth rates (radial increment) vary between 8 and 18 mm year1 for unstressed corals, with extremes of ca. 3 mm year1 for P. lutea at Enewetak (Highsmith, 1979) and 22.7 mm year1 for Porites sp. on the Great Barrier Reef (Lough and Barnes, 2000). For massive Porites with a 0.2–0.5 m colony height range, growth rates along the vertical axis (13.0 þ/ 3.4 mm year1) are significantly higher than along the horizontal axis (10.9 þ/ 3.2 mm year1) (Lough and Barnes, 2000). In the particular case of colonies living on the reef flat, and especially in areas of significant tide range, however, the situation may be reversed. After some time, the vertical growth of the massive and often hemispherical colony becomes limited by the emersion of the top of the colony at low tide, while it continues to grow horizontally, taking an irregular disc shape. The upper surface dies and sometimes becomes the settlement site of other corals, or may be the subject of intense bioerosion, leading to an excavation in the central (older) part of the colony. Given suitable hydrodynamic conditions, the living tissue can grow again on the sides of the excavation. Because of its gross morphological similarity with that of an atoll (at a much smaller scale), such formations have been named “micro-atolls” (Figure 3). Environmental factors other than emersion that control the growth rates of Porites include light, temperature, turbidity, and depth (which itself integrates the action of light, water motion, rates of sedimentation – see “Corals: Environmental Controls on Growth”). In broad terms light enhances calcification, and its variations with season and depth account for some of the variability recorded in growth rates. The physiological

Porites, Figure 3 A micro-atoll of Porites lutea, emerged during a low spring tide on a reef flat. Such micro-atolls can have a diameter of several meters.

and biochemical processes involved in calcification are not yet fully understood and may include a decoupling of calcification from photosynthesis and possibly an intrinsic rhythm of calcification. Growth rates of Porites are significantly related to average sea surface temperature, which is one of the most important factors controlling growth rates. Lough and Barnes (2000) report that in the Indo-Pacific region, including the Hawaiian Archipelago, annual extension rates increase by 3.1 mm year1 for each 1
C increase in average annual SST, within the range 23–29
C. They further surmise that on the Great Barrier Reef, sensitivity of calcification rates to SST suggests that growth rates of Porites will – at least initially – increase as a result of global warming. There is some evidence that calcification rates did increase for much of the twentieth century, but recent findings suggest that the trend has been dampened post-1985, possibly due to a decrease in aragonite saturation levels linked to ocean acidification (De’ath et al., 2009). As mentioned above, other factors such as depth, water motion, fine particle sedimentation, and pollutants play a role in combination in controlling growth rates of Porites. However, no clear picture emerges when considering, for instance, inshore–offshore gradients. Extension rates decrease with distance from shore on the Great Barrier Reef and in Thailand, but increase with distance from shore in Jakarta Bay and in Mayotte. In the latter case, however, Priess et al. (1995) also indicate a decrease in growth rates with depth on the outer slope of the barrier reef. Through the process of calcification, corals incorporate a number of trace elements (various stable isotopes, particularly of oxygen and carbon, fulvic acids, etc.) in their skeleton throughout their lives. Their time of incorporation can be precisely determined, providing a means to characterize the environment at that time. Long-lived species such as massive Porites, in particular, therefore constitute an archive of past climatic and environmental

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conditions that can span several centuries and are a precious tool for paleoclimatic reconstructions (e.g., McGregor and Gagan, 2003).

Mortality Little is know on the natural mortality of the various species of Porites. While some massive species can live for several centuries, the situation is very different for many other massive species that never reach large sizes, and for branching species. Mortality can be the result of catastrophic events such as tropical storms or cyclones (see “Tropical Cyclone/ Hurricane”). These events can bury some colonies under thick layers of sediment and dislodge others, sometimes throwing them onto the reef top or rolling them into unviable depths down the reef slope (Massel and Done, 1993). Large Porites colonies, if over toppled in wellilluminated habitats, rarely die completely. Partial mortality may be observed on the parts of the colonies that come into contact with the substratum, but the rest of the colony may continue to live and grow normally, even colonizing the dead part of the former under surface. Tropical storms and cyclones may be associated with freshwater runoff from flooding coastal rivers. Decrease in salinity, particularly if maintained over significant periods of time, is likely to induce a bleaching of the colonies (i.e., a breakdown of the symbiotic association), which may lead to partial or total mortality (Brown, 1997). Extreme temperatures, particularly high SSTs at the end of the warm season are a frequent cause of coral bleaching. In general, Porites, particularly the large-sized species, are more resistant to extremes of temperatures than many other coral genera, e.g., Acropora, Pocillopora, Galaxea, or Pavona (Marshall and Baird, 2000). However, bleaching resulting from a temperature stress and leading to partial or total mortality has been recorded for a number of Porites species. There are also a number of other causes of bleaching which may affect Porites, in particular heavy sedimentation (or re-suspension) of fine particles, a situation commonly found in inshore reefs (see Turbid-Zone and Terrigenous Sediment-Influenced Reefs), and reefs affected by anthropogenic disturbances near the shore line (dredging, excavating, in-filling). Like many other coral genera, Porites can be affected by natural enemies which impair the integrity of the colonies to various degrees. Infestations of boring mollusks (Lithophaga), sponges (Cliona, Microciona), polychaetes (Spirobranchus), sipunculid worms, and several species of cirripeds are known to disturb natural growth and at times modify the shape of the colonies. In many instances, however, such infestations do not lead to partial or total mortality of the colonies infested but may result (particularly in the case of cirriped infestation) in the development of a pink blotch syndrome. One specialized coral predator, the crown of thorns starfish Acanthaster planci, when present in large numbers, may cause up to 90% hard coral mortality in an area. Porites, however, generally becomes

the prey of Acanthaster only after most other hard coral species have been consumed. It is possible that the massive growth form with a comparatively smooth surface of Porites makes it less easy to be climbed on by Acanthaster, and the possibility of chemical defense has also been suggested (Barnes et al., 1970). Nevertheless, recurrent A. planci outbreaks can cause major depletions of reef-wide Porites populations that cannot necessarily be fully restored between outbreaks, with the potential to cause a rapid run-down of their abundances and sizestructure (Done, 1988). Black band disease, due to an infection by the cyanophyte Phormidium corallyticum, has been reported on Porites spp. on the Great Barrier Reef, but its incidence remains very low, compared to that on other coral genera, most notably Acropora.

Summary The importance of the genus Porites on coral reefs reflects to a large extent the survivability of a handful of very environmentally tolerant species. The genus occurs in both the Atlantic and Indo-Pacific, and it is in the latter where it produces – by scleractinian coral standards – among the largest and most long-lived coral colonies. Porites is a predominant reef frame builder, and an excellent archive of recent past environmental conditions and climate. Longevity, resulting from its capacity to withstand natural and anthropogenic stress and impacts better than most other genera, has also brought the genus to its position of dominance. Massive colonies of Porites are a good example of a “long-lived species that structures coral reef systems” (Cameron and Endean, 1985). Bibliography Barnes, D. J., Brauer, R. W., and Jordan, M. R., 1970. Locomotory response of Acanthaster planci to various species of corals. Nature, London, 228, 342–344. Bernard, H. M., 1905. The family Poritidae, II, The genus Porites, Part 1. Porites of the Indo-Pacific region. Cat. Madreporarian Corals British Museum (Natural History), 5, 1–303. Brown, B. E., 1997. Coral bleaching: causes and consequences. Coral Reefs, 16(suppl), S128–S138. Buddemeier, R. A., and Kinzie, R. A., III, 1976. Coral growth. Oceanography and Marine Biology: An Annual Review, 14, 183–225. Cameron, A. M., and Endean, R., 1985. Do long lived species structure coral reef ecosystems? Proceedings of the 5th International Coral Reef Congress, Tahiti, 6, 211–215. Chornesky, E. A., and Peters, E. C., 1987. Sexual reproduction and colony growth in the scleractinian coral Porites astreoides. The Biological Bulletin, 172, 161–177. Claereboudt, M. R., and Al-Amri, I. S., 2004. Calithiscus tantillus, a new genus and new species of scleractinian coral (Scleractinia Poritidae) from the Gulf of Oman. Zootaxa, 532, 1–8. Claereboudt, M. R., 2006. Reef corals and coral reefs of the Gulf of Oman. The Historical Association of Oman, 344 p. De Blainville, H. M., 1830. Zoophytes In Dictionnaire des Sciences naturelles, Paris, 60, pp. 310–358. De’ath, G., Lough, J. M., and Fabricius K. E., 2009. Declining coral calcification on the Great Barrier Reef. Science, 323, 116–119.

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Done, T. J., and Potts, D. C., 1992. Influences of habitat and natural disturbances on contributions of massive Porites coral to reef communities. Marine Biology, 114, 479–493. Done, T. J., 1988. Simulation of recovery of pre-disturbance size structure in populations of Porites spp. damaged by the crownof-thorns starfish Acanthaster planci. Marine Biology, 100, 51–61. Goreau, N. I., Goreau, T. F., and Hayes, R. L., 1981. Settling, survivorship and spatial aggregation in planulae and juveniles of the coral Porites porites (Pallas). Bulletin of Marine Science, 31, 424–435. Gray, J. E., 1842. Pocilloporidae; In Synopsis British Museum (44th ed.). Harriott, V. J., 1983. Reproductive ecology of four scleractinian species at Lizard Island, Great Barrier Reef. Coral Reefs, 2, 9–18. Harrison, P. L., and Wallace, C. C., 1990. Reproduction, dispersal and recruitment of scleractinian corals. In Dubinsky, Z. (ed.), Coral Reefs. Elsevier, pp. 133–207. Highsmith, R. C., 1979. Coral growth rates and environmental control of density banding. Journal of Experimental Marine Biology and Ecology, 37, 105–127. Highsmith, R. C., 1982. Reproduction by fragmentation in corals. Marine Ecology Progress Series, 7, 207–226. Hunter, C. L., 1985. Assessment of clonal diversity and population structure of Porites compressa (Cnidaria, Scleractinia). Proceedings of the 5th International Coral Reef Congress, Tahiti, 6, 69–74. Hunter, C. L., and Kehoe, C. C., 1986. Patchwork patch reefs: the clonal diversity of the coral Porites compressa in Kaneohe Bay, Hawaii; In Jokiel, P. L., Richmond, R. H., and Rogers, R. A. (eds.), Coral Reef Population Biology, Hawaii Inst. Mar. Biol. Tech. Rep. 37, 124–132. Link, H. F., 1807. Beischreibung der Naturalien Sammlungen der Universität zu Rostock. 3, 161–165. Lough, J. M., and Barnes, D. J., 2000. Environmental controls on growth of the massive coral Porites. Journal of Experimental Marine Biology and Ecology, 245, 225–243. Marshall, P. A., and Baird, A. H., 2000. Bleaching of corals on the Great Barrier Reef: differential susceptibilities among taxa. Coral Reefs, 19, 155–163. Massel, S. R., and Done, T. J., 1993. Effects of cyclone waves on massive coral assemblages on the Great Barrier Reef: meteorology, hydrodynamics and demography. Coral Reefs, 12, 153–166. McGregor, H. V., and Gagan, M. K., 2003. Diagenesis and geochemistry of Porites corals from Papua New Guinea: implications for paleoclimate reconstruction. Geochimica et Cosmochimica Acta, 67, 2147–2156. Nemésio, A., 2005. Machadoporites (Coelenterata Scleractinia), a new generic name for Calithiscus Claereboudt & Al-Amri. Lundiana, 6(suppl.), 59. Pallas, P. S., 1766. Elenchus Zoophytorum. Hagæ Comitum, Apud Oetrus van Cleef, 274–336. Potts, D. C., Done, T. J., Isdale, P. J., and Fisk, D. A., 1985. Dominance of a coral community by the genus Porites (Scleractinia). Marine Ecology Progress Series, 23, 79–84. Priess, K., Thomassin, B. A., Heiss, G. A., and Dullo, W. C., 1995. Variabilité de la croissance de Porites massifs dans les récifs coralliens de Mayotte. Comptes Rendus de l’Academie des Sciences de Paris. Life sciences/Ecology 318, 1147–1154. Quelch, J. J., 1886. Report on the reef corals collected by H. M. S. Challenger during the years 1873–76. Scientific Results Voyage Challenger, London. Zoology, 16, 1–203. Resing, J. M., and Ayre D. J., 1985. The usefulness of the tissue grafting bioassay as an indicator of clonal identity in scleractinian corals (Great Barrier Reef-Australia). Proceedings of the 5th International Coral Reef Congress, Tahiti, 6, 75–81.

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Soong, K., and Lang, J. C., 1992. Reproductive integration in reef corals. The Biological Bulletin, 183, 418–431. Szmant, A. M., 1986. Reproductive ecology of Caribbean reef corals. Coral Reefs, 5, 43–54. Vaughan, T. W., 1912. The Madreporaria and marine bottom deposits of southern Florida. Carnegie Institution Washington Yearbook, 10, 147–165. Veron, J. E. N., 2002. New species described in Corals of the World. Australian Institute of Marine Science, Monograph series, 11, 1–206. Veron, J. E. N., and Pichon, M., 1982. Scleractinia of Eastern Australia. Pt IV Family Poritidae. Australian Institute of Marine Science, Monograph series, 5, 1–159. Verrill, A. E., 1864. List of the polyps and corals sent by the Museum of Comparative Zoology to other institutions in exchange, with annotations. Bulletin of the Museum of Comparative Zoology Harv., 1, 29–60. Wells J. W., 1956. Scleractinia. In Moore, R. C. (ed.), Treatise on Invertebrate Paleontology. Part F Coelenterata. Geol. Soc. America & Univ. Kansas Press, pp. F328–F444.

Cross-references Acanthaster Planci Ecomorphology Forereef/Reef Front Geomorphic Zonation Lagoons Microatoll Reef Flats Sclerochronology Temperature Change: Bleaching Tropical Cyclone/Hurricane

POROSITY VARIABILITY IN LIMESTONE SEQUENCES Barbara H. Lidz U.S. Geological Survey, St. Petersburg, FL, USA

Synonyms Aquifer system; Aquifer vs. aquiclude vs. aquifuge (confining bed); Fenestrate vs. non-fenestrate (non-perforate); Hydrostratigraphic unit; Karst in Florida Boulder Zone; Large voids in fossil coral reef; Porosity affects permeability (fluid flow); Small grain-supported interstitial space in rock or soil Definition Porosity is the state of being porous, as measured by the percentage of bulk volume of a rock or soil that is occupied by space, whether isolated or connected. In hydrocarbon-bearing limestone settings, subsurface porous strata containing the oil or gas usually underlie non-porous caprock through which hydrocarbons cannot pass. In karst-limestone settings, subsurface freshwater aquifers beneath caprock can become contaminated by saltwater intrusion during periods of drought. Islands of the Florida Keys consist of two types of emergent 125-ka limestone, a highly porous fossil coral reef with

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Porosity Variability In Limestone Sequences, Figure 1 (a) Core slab from the Saddlebunch Keys (lower Florida Keys) shows thick calcrete unconformity on top of Miami Limestone oolite. (b) Cut rock sample from the Key Largo Waterway on Key Largo (upper Florida Keys) shows the same calcrete unconformity as it developed on top of Key Largo Limestone coral reef (reprinted from Lidz et al., 2007, with permission). This figure (a and b) is also cited in the Definition and figure caption for the entry entitled, Calcrete/Caliche. See Figure 2a in entry entitled, Florida Keys, for locations of the Saddlebunch Keys and the Key Largo Waterway.

large voids and a less porous oolite with small grains and interstices. Both limestones are capped by impervious laminated Holocene calcrete whose dimensions differ greatly (Figure 1a and b). Porosity variability in the limestones is thought to be the cause. The less permeable oolite retained rainfall moisture longer, allowing longer periods of calcrete buildup. Reddish and brownish layers in both illustrated calcrete samples represent periods of influx of non-carbonate minerals on African dust. The hiatus or gap in these rock records represents an interval of >115 kyr during which no marine or terrestrial deposition is recorded.

Emerged Reefs Florida Keys Oil and Gas Reservoirs and Coral Reefs Ooids Reef Drilling

Bibliography

Synonyms (Last) Deglacial sea-level rise; Post glacial sea-level rise

Lidz, B. H., Reich, C. D., and Shinn, E. A., 2007. Systematic mapping of bedrock and habitats along the Florida reef tract: Central Key Largo to Halfmoon Shoal (Gulf of Mexico). U.S. Geological Survey Professional Paper 1751, http://pubs.usgs.gov/pp/ 2007/1751.

Cross-references Airborne Dust Impacts Calcrete/Caliche Climate Change: Impact of Sea Level Rise on Reef Flat Zonation and Productivity Density and Porosity: Influence on Reef Accretion Rates

POSTGLACIAL TRANGRESSION Guy Cabioch IRD (Institut de Recherche pour le Développement), Bondy CEDEX, France

Definition The post-glacial transgression is the rise in sea level from – 120/130 m to present position in the period 19–6/5 kyears, and caused by global glacial melting. Reef growth mode, tectonic and reconstruction of sea-level variations The corals are reliable markers of sea level (see entry Sea-level Indicators) and several curves of the deglacial

POSTGLACIAL TRANGRESSION

The sea-level curves from the last deglacial rise Corals can be used to reconstruct entirely or partly the deglacial sea-level rise curves. Several areas were investigated in the uplifted sites (island arcs) from Barbados (from 22.1 to 7.4 ka, 1 ka = 1,000 years (Fairbanks, 1989; Bard et al., 1990; Peltier and Fairbanks, 2006) and Papua New Guinea (Cutler et al., 2003 and from 13.1 to 8.3 ka, see Edwards et al., 1993) and in the subsiding site of Tahiti in French Polynesia (from 13.8 to 2.8 ka, Bard et al., 1996, 2010, Figure 1). Additional data are available in South-West Pacific in Vanuatu (from 22.6 to 6.6 ka, Cabioch et al., 2003) and in North-West Pacific in the Ryukyus (Sasaki et al., 2006). In the first curve using coral dating presented by Fairbanks (1989), two episodes of Meltwater Pulses (MWP) were identified, the MWP 1A at around 14 ka and the MWP 1B at 11.3 ka. Although these events will be presented in the entry “Meltwater Pulses”, a short description is also given below. Episodes of rapid sea-level rise During the last deglacial sea-level rise, several episodes of rapid rise occur. The existence, the timing, the magnitude, and the cause of some of these sharp rises are however still

0 20 MWP-1B ?

40 Depth (m)

sea level were proposed using their dating (see entry Uranium Series Dating). The reconstruction of these curves must take into account several parameters including the reef-growth patterns and the tectonic parameters (rate of vertical motions). The mode of reef growth must be examined with accuracy because the organisms respond differentially to more or less rapid sea-level rise (see entries Corals: Environmental Controls on Growth and Sea Level Change and Its Effect on Reef Growth). Moreover, the corals can have a broad living bathymetric range. Three growth mode strategies were defined by Neumann and Macintyre (1985) and Davies and Montaggioni (1985): (1) the “keep-up” growth mode strategy characterizes the reefs able to maintain pace with sea-level rise, composed of relatively shallow-water and high-wave energy organisms; (2) the “catch-up” growth mode strategy characterizes the reefs, which were not initially able to keep pace with sea-level rise, but caught up as or after stabilizations of the sea level. The communities are made of relatively deeper, less wave-resistant and less light-intensity organisms, progressively replaced upwards by relatively shallow-water and high-wave energy resistant organisms; (3) the “give-up” mode strategy, the organisms are not able to compensate the sea-level rise and gradually drowned. The corrections due to the tectonic parameters must also be taken into account. The uplift rate in the island arc sites and the subsidence in the intra-plate sites or in the continental margins, must be used to correct the sealevel curves. Moreover, the correction due to the loading of the ocean floor by the meltwater (hydro-isostatic effects) should also be taken into account (see Lambeck et al., 2002).

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60 80

19 ka

100 Barbados (Bard et al., 1990) Tahiti (Bard et al., 1996)

120 0

5

10

15 Years (cal ka BP, 20 1 ka = 1,000 years)

Postglacial Trangression, Figure 1 Sea level curves deduced from the coral dating in Barbados and Tahiti (Modified after Bard et al., 1996).

debated. Three episodes have been proposed: the 19 ka event, the MWP 1A, and the MWP 1B. Some of these episodes were attributed to episodes of Meltwater Pulses (MWP, see Chapter Meltwater Pulses), but the source needs to be clearly identified. According to some authors, an additional event should also be considered at around 7.5 ka (Blanchon et al., 2002; Yu et al., 2007; Siddall et al., 2010).

The 19 ka event The 19 ka event defined in the Bonaparte Gulf (Australia) by Yokoyama et al. (2000), at the end of the Last Glacial Maximum (LGM) is characterized by a rapid sea-level rise of 10–15 m. Clark et al. (2004) confirm the occurrence at 19 ka of a rapid 10-m-sea-level rise deduced from the sediment analyses of the Irish Sea. Drowning of reefs in the Marquesas (Cabioch et al., 2008) at this time adds further support to such an event. The cause of this 19 ka event is attributed by Clark et al. (2004) to the melting of at least one Northern Hemisphere ice sheet but this needs to be constrained by further geochronological data (Clark, 2009). The MWP 1A The second event at 14.5 ka is marked by a rapid sea-level rise of 20 m in less than 500 years. It is reported at various sites but the timing and the source are also still actively debated. Fairbanks (1989) first reported this event in the Barbados, but it is also observed in numerous sites from the Pacific and Atlantic Oceans. In the Sunda shelf, the sea level increased rapidly by 16 m from 14.6 to 14.3 ka. An apparent discrepancy in the timing of the MWP 1A between Barbados (Fairbanks, 1989; Bard et al., 1990; Peltier and Fairbanks, 2006), Papua New Guinea (Cutler et al., 2003), and Bonaparte Sea (Yokoyama et al., 2000) reported in Stanford et al. (2006) is probably due to the fact that the living bathymetric range of dated corals was

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not taken into account to correct the sea-level curve (Clark, 2009). The reef growth mode needs always to be considered. Nevertheless, Peltier and Fairbanks (2006) highlight the fit between the record from Barbados and Sunda shelf. The possible sources of the MWP 1A are also highly debated. The ice melting from Northern Hemisphere is probably accompanied by a contribution of the Antartica (Clark, 2009), although Peltier and Fairbanks (2006) consider that Antartica is not a significant contributor of the source of the MWP 1A. All these discussions emphasize the need to acquire new data to improve the interpretations and the models. In addition to the study of living reefs, studies of drowned reefs also sheds light on the MWP-1A event. In Hawaii, the drowning of the 150-m-coral reefs, constrained by its datation from 15.2 to 14.7 ka, is interpreted by Webster et al. (2004) as caused by the rapid sea-level rise of the MWP 1A. Similar observations can be made in the Marquesas archipelago, the drowning of the coral reefs ranging from 110 to 95 m and dated from 16.2 to 14.5 ka (Cabioch et al., 2008) can be also a consequence of the rapid sea-level rise of the MWP 1A. Similar sites of drowned reefs can be observed in various regions but at this day, no dates were performed and we cannot conclude on the causes of the drowning.

The MW 1B The MWP 1B was defined in Barbados by Fairbanks (1989). Nevertheless, the existence of such an event is still debated and discussed. Recently, Peltier and Fairbanks (2006) indicated that the source from Antartica of the MWP 1B is assumed by the model ICE-nG. Although the timing and the magnitude of the MWP 1B cannot be confirmed by the observations of the deglacial sea-level curves from Papua New Guinea (West Pacific) and Tahiti (South Pacific), this event seems to be recognizable in the West Pacific in Vanuatu (Cabioch et al., 2003) and in the Indian Ocean in Mayotte island (Zinke et al., 2003). Although the magnitude of the MWP 1B, if real, seems to be smaller than in the MWP 1A, significant changes can be observed in reef growth patterns in Vanuatu at this period (Cabioch et al., 2003): it appears that around 11.5 ka, relatively shallow-water communities were replaced by relatively deeper organisms, that is, the reef was not able to maintain pace with sea-level rise and the water depth strongly increased over the reef. Similarly, in the Marquesas archipelago, reefs at depths between 80 and 68 m and dated at around 12.4 ka seem to have been drowned (Cabioch et al., 2008). However, the existence of this event is still discussed and debated because most of coral reefs continued to grow during this period as for example in Tahiti (Bard et al., 1996). New uraniumthorium dates were recently obtained on corals from Tahiti providing a new Tahiti sea-level record (Bard et al., 2010). This new record points out no significant change in the sea-level rise during the period attributed to the MW 1B, but a slower rate of sea-level rise during the Younger Dryas (11.6–12.9 ka).

The 7.5 ka event In 2002, Blanchon et al. reported the occurrence of a 6-mrapid jump of the sea level at around 7.5 ka. Such an event at this period is rarely reported (Yu et al., 2007) and needs to be validated but many modern coral reefs settled after this date, especially in the Pacific (Montaggioni, 2005). Additional data must be collected regarding the occurrence of this event. Prospects In 2006, in the framework of the international program IODP (Integrated Ocean Drilling Program), the Tahiti Sea Level Expedition 310 was conducted by G. Camoin, Y. Iryu, and D. McInroy in Tahiti to improve our knowledge on the last deglacial sea-level rise, to analyze the SST changes during this period, and to define the reef development controlled by these two last factors (Camoin et al., 2007). Several cores were collected in 22 sites from three areas. One of the objectives of this expedition was to define the sea-level variations over the period from 20 to 10 ka. At this day, the results are not yet published but the data probably will improve our knowledge on the timing and the sources of the MWP. Another IODP expedition is currently in progress in the Australian Great Barrier Reef (Expedition 325) focused on similar topics, that is, to analyze the pattern of the last deglacial sea-level rise far from the polar sites and to provide data to solve the timing and the magnitude of the brief and rapid increases of sea-level rise (Webster et al., 2009). Summary The postglacial transgression is of broad interest, interacting with the reef growth patterns and the history of the reef development. The last sea-level rise is marked by episodes of rapid increase. Some of them are clearly attributed to melt water pulses but the timing, the magnitude, and the melting source are still debated and discussed. Additional data in sites tectonically stable and far from the polar sites are required to solve these questions. The corals and the coral reefs are reliable contributors and the key to reconstruct with accuracy the last deglacial sea-level curves, the timing and magnitude of the brief and rapid episodes of sea-level rise on condition that the bathymetric range of coral dated can be estimated and the reef growth mode defined. Bibliography Bard, E., Hamelin, B., and Fairbanks, R. G., 1990. U/Th ages obtained by mass spectometry in corals from Barbados. Sea level during the past 130 000 years. Nature, 346, 456–458. Bard, E., Hamelin, B., Arnold, M., Montaggioni, L. F., Cabioch, G., Faure, G., and Rougerie, F., 1996. Deglacial sea level record from Tahiti corals and the timing of global meltwater discharge. Nature, 382, 241–244. Bard, E., Hamelin, B., and Delanghe-Sabatier, D., 2010. Deglacial meltwater pulse 1B and Younger Dryas sea levels revisited with boreholes at Tahiti. Science, 327, 1235–1237.

POSTGLACIAL TRANGRESSION

Blanchon, P., Jones, B., and Ford, D. C., 2002. Discovery of a submerged relic reef and shoreline off Grand Cayman: further support for an early Holocene jump in sea level. Sedimentary Geology, 147, 253–270. Cabioch, G., Banks-Cutler, K., Beck, W. J., Burr, G. S., Corrège, T., Edwards, R. L., and Taylor, F. W., 2003. Continuous reef growth during the last 23 ka in a tectonically active zone (Vanuatu, SouthWest Pacific). Quaternary Science Reviews, 22, 1771–1786. Cabioch, G., Montaggioni, L. F., Frank, N., Seard, C., Sallé, E., Payri, C. E., Pelletier, B., and Paterne, M., 2008. Successive reef depositional events along the Marquesas foreslopes (French Polynesia) since 26 ka. Marine Geology, 254, 18–34. Camoin, G. F., Iryu, Y., McInroy, D. B., and the Expedition 310 Scientists, 2007. Proc. IODP, 310: Tahiti Sea Level (IODP, Integrated Ocean Drilling Program Management International, Inc.). In Proceedings IODP “Expedition Reports Tahiti Sea Level”, IODP Management International, Inc., Washington, DC, Vol. 310. Clark, P. U., 2009. Ice sheet retreat and sea level rise during the last deglaciation. PAGES News, 17(2), 64–66. Clark, P. U., McCabe, A. M., Mix, A. C., and Weaver, A. J., 2004. Rapid rise of sea level 19,000 years ago and its global implications. Science, 304, 1141–1144. Cutler, K. B., Edwards, R. L., Taylor, F. W., Cheng, H., Adkins, J., Gallup, C. D., Cutler, P. M., Burr, G. S., and Bloom, A. L., 2003. Rapid sea-level fall and deep-ocean temperature change since the last interglacial period. Earth and Planetary Science Letters, 206, 253–271. Davies, P. J., and Montaggioni, L. F., 1985. Reef growth and sea level change: the environmental signature. Proceedings of the 5th Internatioal Coral Reef Congress, Tahiti, Vol. 3, pp. 477–515. Edwards, R. L., Beck, W. J., Burr, G. S., Donahue, D. J., Chappell, J. M. A., Bloom, A. L., Druffel, E. R. M., and Taylor, F. W., 1993. A large drop in atmospheric 14C/12C reduced melting in the Younger Dryas, documented with 230Th ages of corals. Science, 260, 962–968. Fairbanks, R. G., 1989. A 17,000-year glacio-eustatic sea level record: influence of glacial melting rates on the Younger Dryas event and deep-ocean circulation. Nature, 342, 637–642. Lambeck, K., Yokoyama, Y., and Purcell, A., 2002. Into and out of the Last Glacial Maximum: sea level change during Oxygen Isotope Stages 3 and 2. Quaternary Science Reviews, 21, 343–360. Montaggioni, L. F., 2005. History of Indo-Pacific coral reef systems since the last glaciation: development patterns and controlling factors. Earth Science Reviews, 71, 1–75. Neumann, A. C., and Macintyre, I., 1985. Reef response to sea level rise: keep-up, catch-up or give-up. In Proceedings of the 5th Internatioal Coral Reef Congress, Tahiti, Vol. 3, pp. 105–110.

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Peltier, W. R., and Fairbanks R. G., 2006. Global glacial ice volume and Last Glacial Maximum duration from an extended Barbados sea level record. Quaternary Science Reviews, 25, 3322–3337. Sasaki, K., Omura, A., Miwa, T., Tsuji, Y., Matsuda, H., Nakamori, T., Iryu, Y., and Yamada, T., 2006. Th-230/U-234 and C-14 dating of a lowstand coral reef beneath the insular shelf off Irabu Island, Ryukyus, southwestern Japan. Island Arc, 15, 455–467. Siddall, M., Abe-Ouchi, A., Andersen, M., Antonioli, F., Bamber, J., Bard, E., Clark, C. J., Clark, P., Deschamps, P., Dutton, A., Elliot, M., Gallup, C., Gomez, N., Gregory, J., Huybers, P., Kawarnura, K., Kelly, M., Lambeck, K., Lowell, T., Milrovica, J., Otto-Bliesner, B., Richards, D., Stanford, J., Stirling, C., Stocker, T., Thomas, A., Thompson, B., Tornqvist, T., Riveiros, N. V., Waelbroeck, C., and Yokoyama, Y., 2010. The sea-level conundrum: case studies from palaeo-archives. Journal of Quaternary Science, 25, 19–25. Stanford, J. D., Rohling, E. J., Hunter, S. E., Roberts, A. P., Rasmussen, S. O., Bard, E., McManus, J., and Fairbanks, R. G., 2006. Timing of meltwater pluse 1 A and climate responses to melwater injections. Paleoceanography, 21, Article no PA4103. Webster, J. M., Clague, D. A., Coleman-Riker, K., Gallup, C., Braga, J. C., Potts, D., Moore, G. J., Winterer, E., and Paull, C. K., 2004. Drowning of the 150 m reef off Hawaii: a casualty of global meltwater pulse 1A? Geology, 3, 49–252. Webster, J. M., Yokoyama, Y., and Cotterill, C., 2009. Great Barrier Reef environmental changes: the last deglacial sea level rise in the South Pacific: offshore drilling northeast Australia. IODP Scientific Prospectus, 325, doi:10.2204/iodp.sp.325.2009. Yokoyama, Y., Lambeck, K., De Deckker, P., Johnston, P., and Fifield, L. K., 2000. Timing of the Last Glacial Maximum from observed sea-level minima. Nature, 406, 713–716. Yu, S.-Y., Berglund, B. E., Sandgren, P., and Lambeck, K., 2007. Evidence for a rapid sea-level rise 7600 yr ago. Geology, 35, 891–894. Zinke, J., Reijmer, J. J. G., Thomassin, B. A., Dullo, W.-Chr., Grootes, P. M., and Erlenkeuser, H., 2003. Postglacial £ooding history of Mayotte Lagoon (Comoro Archipelago, southwest Indian Ocean). Marine Geology, 194, 181–196.

Cross-references Glacio-Hydro Isostasy Meltwater Pulses Radiocarbon (14C): Dating and Corals Sea Level Change and Its Effect on Reef Growth Sea-level Indicators Submerged Reefs Uranium Series Dating

Q

QUOY, JEAN RENE (1790–1869) AND GAIMARD, JOSEPH PAUL (1796–1858) James Bowen Southern Cross University, Lismore, NSW, Australia Following the Napoleonic Wars, further scientific investigation of the southern oceans by the competitive powers of England and France resumed, although the French expedition 1800–1804 under Baudin in command of Le Géographe and Naturaliste was a fiasco marred by scurvy of his crew and conflict with the naturalist François Auguste Péron. Since the British were making important gains in the south Pacific the French renewed their interests and in September 1817 L’ Uranie under the command of Louis Claude Desaules de Freycinet sailed from Toulon, carrying aboard surgeon and zoologist Dr. Jean Réne Constand Quoy and assistant surgeon and zoologist Joseph Paul Gaimard who were to make significant contributions to helping solve the puzzle of atoll and coral reef formation in the otherwise great open spaces of the tropic oceans. By the early nineteenth century, reef atoll formation was known to have come from small animal polyps, but many other questions, however, remained unanswered. Given the enormous depth of the surrounding waters, unable to be sounded by the technology of the times, how had polyps established themselves in the first place? Upon what foundations had they erected their limestone structures? What was the nature of that “instinct” proposed by Forster by which “the animalcules forming these reefs. . . shelter their habitation from the impetuosity of the winds?” Most importantly, why were the atolls circular in shape – many being miles in diameter – readily confirmed from the crow’s-nest of the exploring ships? In that period of extensive reef investigation from 1817 to 1820, chiefly to the Mariana, Hawaiian and Dutch East

Indies island groups, the two zoologists offered an explanation from their findings in 1824 entitled Mémoire sur l’Accroissement des polypes lithophytes considéré géologiquement (“Geological aspects of coral formation”). In that joint paper they argued from their extensive examination of reefs in Pacific tropical waters that it would be a mistake to ascribe all atoll formation to polyps alone, growing up from the ocean floor. Rather, they believed that coral reefs are surface features that have as a base the same element, the same minerals which concur to form all the known islands and continents. . . that [in effect] they build their dwellings on the submarine rocks, enveloping them entirely, or in part, but properly speaking they do not form them. Thus, all these reefs, they concluded, “are, in our opinion, platforms arising from the conformation of the primitive surface.” Tragically, on the return voyage approaching Patagonia L’ Uranie was wrecked on an uncharted reef, and although the ship was beached, their supplies and a large part of their natural history specimens were lost as the violent seas rapidly destroyed the foundering vessel. Eventually, however, they were rescued by a passing American vessel which carried them and their scientific specimens to Montevideo from where, when they arrived in France their specimens – mammals, birds, fish and insects – were received by the Academy of Sciences with considerable admiration. What created intense interest were the jars of preserved marine creatures, particularly polyps and invertebrates, and the drawings made by the artists and naturalists of the expedition. In 1826 Quoy and Gaimard were to sail again on a voyage of exploration to the south Pacific aboard the frigate L’ Astrolabe under the command of Jules Sébastien César Dumont d’Urville where they were able to continue their observations on coral formations and the taxonomic distinction between those described by Peyssonnel as “supple and pliant” and “stony, hard, inflexible”.

David Hopley (ed.), Encyclopedia of Modern Coral Reefs, DOI 10.1007/978-90-481-2639-2, # Springer Science+Business Media B.V. 2011

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Bibliography Quoy, J. R. C., and Gaimard, J. P., 1824. Zoologie. In Freycinet, L. C. D. de (ed.), Voyage autour du monde . . . par M. Louis de Freycinet. Paris: Pillet Aîné. Quoy, J. R. C., and Gaimard, J. P., 1824. Mémoire sur l’accroissement des Polypes lithophtes considéré geologiquement. In Freycinet, L. C. D. de (ed.), Voyage autour

du monde . . . par M. Louis de Freycinet. Paris: Pillet Aîné, Ch XV, pp. 658–671; reprinted, Annales des Sciences Naturelles, VI. 1825, pp. 273–290. Quoy, J. R. C., and Gaimard, J. P., 1828. Remarques sur les polypes à polypiers pirreux et flexibles. Annales des Sciences Naturelles, XIV, 236–249.

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RADIOCARBON (14C): DATING AND CORALS Stewart Fallon Australian National University, Canberra, ACT, Australia

Definition Radiocarbon or 14C is the radioactive isotope of carbon. It is the basis for radiocarbon dating and is useful for dating materials that contain carbon back in time to around 50, 000 years ago. Introduction There are three isotopes of carbon found in nature. They are carbon-12, carbon-13, and carbon-14. Carbon-12 accounts for
99.8% of all carbon atoms, carbon-13 accounts for
1% of carbon atoms while
1 in every 1 billion carbon atoms is carbon-14. Hereafter, these isotopes will be referred to as 12C, 13C, and 14C. 14C is radioactive and has a half-life of 5,730 years. The half-life is the time taken for an amount of a radioactive isotope to decay to half its original value. Because this decay is constant it can be used as a “clock” to measure elapsed time assuming the starting amount is known. A unique characteristic of 14C is that it is constantly formed in the atmosphere (Figure 1). Production and decay 14 C atoms are produced in the upper atmosphere where neutrons from cosmic rays knock a proton from nitrogen-14 atoms. These newly formed 14C atoms rapidly oxidize to form 14CO2. They are chemically indistinguishable from 12CO2 and 13CO2. They then join the earth’s carbon cycle. Plants incorporate 14 C during photosynthesis and organisms that eat plants take up this 14C. 14C gets into the dissolved inorganic carbon pool in the oceans, lakes, and rivers. From there it gets into shell,

corals, and other marine organisms. When a plant or an animal dies, it no longer exchanges CO2 with the atmosphere (ceases to take 14C into its being). This starts the radioactive decay “clock.” 14C decays by emitting an electron, which converts a neutron to a proton, converting it back to its original 14N form.

History of radiocarbon dating Willard Libby invented radiocarbon dating in the late 1940s. This invention revolutionized science as a means to provide ages to events over the past
50,000 years. Libby was awarded the Nobel Prize for chemistry for this contribution. His first publication showed the comparisons between known age samples and radiocarbon age (Libby et al., 1949; Libby, 1952). Measuring 14C To obtain the radiocarbon age of a sample one must determine the proportion of 14C it contains. Originally, this was done by what is known as “conventional” methods, either proportional gas counters or liquid scintillation counters. The gas counter detects the decaying beta particles from a carbon sample that has been converted to a gas (CO2, methane, and acetylene). A liquid scintillation measurement needs the carbon to be converted into benzene, the instrument then measures the flashes of light (scintillations) as the beta particles interact with a phosphor in the benzene. The main limitation of these techniques is sample size, and hundreds of grams of carbon are needed to count enough decaying beta particles. This is especially true for old samples with low beta activity. In the late 1970s and early 1980s, the dating of small samples became possible using accelerator mass spectrometry (AMS; Muller, 1977; Nelson et al., 1977). This method needs less than 1 mg of carbon and directly measures the abundance of the individual ions of carbon (14C, 12C, and 13C).

David Hopley (ed.), Encyclopedia of Modern Coral Reefs, DOI 10.1007/978-90-481-2639-2, # Springer Science+Business Media B.V. 2011

RADIOCARBON (14C): DATING AND CORALS

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To obtain a radiocarbon age the sample activity or the C/12C ratio must be compared to a standard material of known age. All radiocarbon laboratories either standardize to the US National Bureau of Standards Oxalic Acid I (OX-I), which is derived from Sugar Beets in grown in 1955 or a secondary standard NBS OX-II (grown in 1977) or Australian National University Sucrose (ANU), which is sugar from the 1974 sugar cane growing season in Australia. Both the OX-II and ANU have been extensively cross calibrated to OX-I and can be used to normalize a sample for radiocarbon dating. The absolute radiocarbon standard is 1890 wood, the OX-I standard has an activity of 0.95 of this wood. The definition of year “0,” “modern” or “present” is 1950, there is no real reason for this other than to commemorate the publication of the first radiocarbon dates. The radiocarbon age is determined by the equation

14

t ¼ 8; 033lnðAsn=AonÞ where –8,033 represents the mean lifetime of 14C (Stuiver and Polach, 1977), Asn is the activity in counts per minute of the sample, and Aon is the counts per minute of the modern standard. A variant of this equation is also used when

the samples are analyzed by AMS. All radiocarbon ages are normalized to a d13C of –25% relative to Pee Dee Belemnite (PDB).

Calibration to time In the 1950s, it was observed that the radiocarbon timescale was not perfect. The age of known artefacts from Egypt were too young when measured by radiocarbon dating. A scientist from the Netherlands (Hessel de Vries) tested this by radiocarbon dating tree rings of known ages (de Vries, 1958). He noted some discrepancies indicating that radiocarbon results would need to be “calibrated” to convert them to calendar ages. de Vries also postulated that the fluctuations were due to the production of 14C and how it changed during variations in cosmic ray production. This brings us to two reasons why a radiocarbon date is not a true calendar age. The true half-life of 14C is 5,730 years and not the originally measured 5,568 years used in the radiocarbon age calculation, and the proportion of 14C in the atmosphere is not consistent through time. The latter is due in part to fluctuations in the cosmic ray flux into our atmosphere (e.g., sunspot activity). Since then there have been many studies examining the

Cosmic radiation Carbon 14 Nitrogen 14

Neutron capture

All three isotopes of carbon, (common C-12, rare C-13 and radioactive C-14) are absorbed by living organisms

Soil Following death and burial, wood and bones lose C-14 as it changes to N-14 by beta decay. Nitrogen 14 Carbon 14

Beta decay

Beta particle Proton

Neutron

Radiocarbon (14C): Dating and Corals, Figure 1 Schematic of 14C production and decay in the atmosphere. 14C is produced in the atmosphere by cosmic neutrons colliding with Nitrogen atoms. The newly formed 14C is oxidized to 14CO2 where it then enters the biosphere. Following an organism’s death, radioactive decay occurs converting the 14C back to 14N.

RADIOCARBON (14C): DATING AND CORALS

variations in the 14C production and its effects on the radiocarbon age to calendar age calibration (e.g., Stuiver, 1971; Edwards et al., 1993; Kitagawa and Van de Plicht, 1998; Stuiver et al., 1998; Fairbanks et al., 2005). The proportional amount of 14C to total carbon has also changed during the industrial revolution (
1890). Since fossil fuel is derived from millions of year old organic carbon, it contains no 14C. The burning of fossil fuels has caused a dilution of 14C in the atmosphere, this is the so-called Suess effect named after Hans Suess (Suess, 1965, 1980). It is essential to have radiocarbon ages calibrated to calendar ages so as to have an accurate measure of time. It is also important to be able to compare ages with samples dated by other means, for example, uranium-series dating. It therefore became necessary to create a calibration between radiocarbon dates and calendar age. The ideal calibration material must have a precise calendar age and sample the atmosphere (carbon reservoir of interest).

Tree-ring calibration Fortunately, there is a suitable calibration material available in nature and this is called annual tree rings. Since those first measurements in the 1950s a detailed, precise calibration between radiocarbon and calendar age has been developed using many long-lived tree species. Dendrochronology provides the accurate calendar age for each ring in the tree, and then a radiocarbon age can be assigned to each calendar age. Several tree-ring chronologies have been constructed including the Belfast Irish Oak chronology (Baillie et al., 1983; Brown et al., 1986) back to
7,200 years and the Stuttgart–Hohenheim oak and pine chronology (e.g., Friedrich et al., 2004; Schaub et al., 2008; Hua et al., 2009) back to
12, 594 years using thousands of accurately dated tree rings. However this is as far back in time as the continuous tree-ring radiocarbon calibration can be extended at present. More old trees are being discovered every year and this may eventually increase this calibration dataset at a later date. There are also a number of “floating” tree-ring chronologies that are being developed. They are called floating because they do not have a direct calendar age and must use the radiocarbon to match their ages. Many sections of old subfossil New Zealand Kauri trees have been found that span time from 25 to 60,000 years old (Hogg et al., 2006; Turney et al., 2007). In order to extend the calibration curve back to 50,000 years other archives have been targeted, these include Foraminifera preserved in continuous annually laminated (varved) sediments (e.g., Kitagawa and van der Plicht, 2000; Hughen et al., 2004), tropical surface corals (e.g., Bard et al., 1998; Burr et al., 1996, 2009; Cutler et al., 2004; Fairbanks et al., 2005; Chiu et al., 2007), and speleothems (e.g., Beck et al., 2001; Genty et al., 1999). These additional archives extract their carbon from the dissolved inorganic carbon surrounding them and not directly from the atmosphere. But the carbon is closely linked to the atmosphere and can be corrected for this

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offset. Both fossil corals and speleothems can be precisely dated to a calendar age by uranium-series dating and therefore can be used in extending the radiocarbon calibration curve. Beyond the tree-ring calibration, fossil corals and Foraminifera from Cariaco Basin and the Iberian Margin make up the calendar age to radiocarbon age calibration (Figure 2).

Calibration curves Over the last 20 or so years there have been several calibration “curves” ratified by the Radiocarbon International Community (IntCal98, IntCal04, IntCal09; Stuiver et al., 1998; Reimer et al., 2004, McCormac et al., 2004; Reimer et al., 2009), there have also been other calibration curves proposed by individual research groups (e.g., Fairbanks et al., 2005). Figure 2 shows the most recent IntCal09 calibration curve superimposed over many of the coral and foram varve archives. These calibration curves form the basis of several online calibration programs that take the radiocarbon age and output a calibrated age, the four major online calibration programs are the following: Calib – http://intcal.qub.ac.uk/calib/ CalPal – http://www.calpal.de/ OxCal – http://c14.arch.ox.ac.uk/embed.php?File=oxcal. html Fairbanks – http://radiocarbon.ldeo.columbia.edu/research/ radcarbcal.htm Radiocarbon in the ocean Marine organisms have a further complication when it comes to radiocarbon dating. The exchange between the ocean and atmospheric 14CO2 takes on average 10 years to come into equilibrium (Broecker et al., 1985), which never completely happens. Because the reservoir of carbon in the ocean is so vast and the mixing between the surface and deep ocean is sufficiently long, radioactive decay of carbon in the ocean occurs. The deep ocean can have an apparent age of several thousand years. This old carbon mixes upward by a process called “upwelling.” Amounts of upwelling vary throughout the oceans of the world. This results in the surface ocean having an average apparent age of
400 years although there is considerable variability. This is called the reservoir age or reservoir effect (e.g., Druffel et al., 2008; Eiriksson et al., 2004; Franke et al., 2008). Marine shells of known age collected prior to 1955 and independently dated corals have been used to measure this reservoir variability (e.g., Bourke and Hua, 2009; Culleton et al., 2006; Petchy et al., 2009). Online databases are available to estimate the reservoir age of a marine sample (Reimer and Reimer, 2009 http:// calib.org/marine; Butzin et al., 2005; http://radiocarbon. ldeo.columbia.edu/research/resage.htm). This is then used to adjust the radiocarbon age and calibrate to a calendar age. A full marine calibration curve is also available (Marine04, Marine09) to calibrate a marine radiocarbon age, it was calculated using an ocean-atmosphere box

RADIOCARBON (14C): DATING AND CORALS

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Conventional radiocarbon age (yr BP)

50⫻103

1:1 line

40

30

20 Barbados coral (Fairbanks et al., 2005) Kirimati coral (Fairbanks et al., 2005) Araki coral (Fairbanks et al., 2005) Cariaco basin (Hughen et al., 2006) INTCAL09 (Reimer et al., 2009)

10

0

0

10

30 20 Calendar age (yr BP)

40

50⫻103

Radiocarbon (14C): Dating and Corals, Figure 2 Radiocarbon calibration figure, conventional radiocarbon age on the y-axis vs. Calendar age on the x-axis. The IntCal09 dataset is used back to 50,000 years BP comprises tree-ring data, (Reimer et al., 2009, red line), corals from various locations can also be used to constrain this calibration. Shown here is corals data from Barbados (blue square), Kirimati (green triangle), and Araki (purple diamond) from Fairbanks et al. (2005). Also shown are Foraminifera from Cariaco Basin in the thin line (Hughen et al., 2006).

diffusion model for the time period 0–10,500 years (Oeschger et al., 1975; Stuiver and Braziunas, 1993) and foram and coral data from 10,500 to 12,500 years as described in Hughen et al. (2004). Beyond 12,500 years, the atmospheric calibration curve is used with a constant reservoir age of 405 years (Reimer et al., 2009).

“Bomb” radiocarbon During the 1950s and 1960s, nuclear weapons testing generated excess neutrons in the atmosphere thereby creating manmade 14C. This production ceased in 1963 with the signing of the nuclear test ban treaty, however not before the 14C/C ratio in the atmosphere nearly doubled. This “bomb” radiocarbon has been used to help understand the uptake of CO2 by the ocean and by the terrestrial biosphere. The subsequent invasion of this “Bomb” 14C into the surface ocean has increased the radiocarbon difference between the surface and the deep ocean (e.g., Broecker, et al., 1985). The use of 14C as a global ocean circulation tracer was a primary objective of the study of the distribution of natural and bomb-produced 14C in the Geochemical Ocean Sections Study (GEOSECS) of the early 1970s (Ostlund and Stuiver, 1980; Broecker et al., 1985) and of the present-day World Ocean Circulation Experiment (WOCE; Key et al., 1996). The GEOSECS data identified a surface water gradient of post-bomb 14C from the equator toward the temperate latitudes. Broecker and Peng (1982) interpreted this distribution as representing

upwelling of low 14C water from the lower thermocline in equatorial regions, with migration of the 14C rich surface water toward higher latitudes. Radiocarbon measurements of coral skeletal material have been used to study how the radiocarbon content of the tropical surface ocean has varied through time (e.g., Druffel 1981; Guilderson et al., 2000, 2009). Many coral genera construct massive colonies often 200–400 years old, which in shallow reef environments have growth rates on the order of 1 cm year1. Because the radiocarbon in the coral aragonite skeleton reflects seawater radiocarbon content at the time of deposition, radiocarbon measurements across annual skeletal density bands in such corals make it possible to reconstruct the annual mean radiocarbon content of the surface ocean back to pre-bomb and preindustrial values

Summary Radiocarbon is a useful means for obtaining the age of death of a carbon-bearing organism. With the help of international scientists, a robust calibration has been developed back to 50,000 years ago. Annual tree rings provide the calibration back to
12,594 year BP and corals and forams helped refine this calibration back to 50,000 years ago using uranium-series dating in conjunction with radiocarbon dating. Corals have also played a role in trying to understand the oceanic uptake of CO2 and for tracking ocean currents and circulation.

RADIOCARBON (14C): DATING AND CORALS

Bibliography Baillie, M. G. L., Pilcher, J. R., and Pearson, G. W., 1983. Dendrochronology at Belfast as a Background to High-Precision Calibration. Radiocarbon, 25, 171–178. Bard, E., Arnold, M., Hamelin, B., Tisnerat-Laborde, N., and Cabioch, G., 1998. Radiocarbon calibration by means of mass spectrometric 230Th/234U and 14C ages of corals: an updated database including samples from Barbados, Mururoa and Tahiti. Radiocarbon, 40(3),1085–1092. Beck, J. W., Richards, D. A., Edwards, R. L., Silberman, B. W., Smart, P. L., Donahue, D. J., Herrera-Osterheld, S., Burr, G. S., Calsoyas, L., Jull, A. J. T., and Biddulph, D., 2001. Extremely large variations of atmospheric 14C concentrations during he last glacital period. Science, 292(5526), 2453–2458. Bourke, P., and Hua Q., 2009. Examining late Holocene marine reservoir effect in archaeological fauna at Hope Inlet, Beagle Gulf, north Australia. In Fairbairn, A., O’Connor, S., and Marwick, B. (eds.), New Directions in Archaeological Science. Terra Australis, Canberra: ANU E, Vol. 28, pp. 175–185. Broecker, W. S., Peng, T.-S., Ostlund, G., and Stuiver, M., 1985. The distribution of bomb radiocarbon in the ocean. Journal of Geophysical Research, 90, 6953–6970. Broecker, W. S., and Peng, T.-S., 1982. Tracers in the sea, LamontDoherty Geological Observatory, 690 pp. Bronk Ramsey, C., 1995. Radiocarbon calibration and analysis of stratigraphy: OxCal program. Radiocarbon, 37, 425–430. Brown, D. M., Munro, M. A. R., Baillie, M. G. L., and Pilcher, J. R., 1986. Dendrochronology – the Absolute Irish Standard. Radiocarbon, 28, 279–283. Burr, G. S., Gray, S. C., Edwards, R. I., Taylor, F. W., Donahue, D. J., Recy, J., Cabioch, G., and Beck, J. W., 1996. A radiocarbon calibration during the last deglaciation based on TIMS 230Th ages of AMS 14C dated corals from Vanuatu, New Hebridies. Radiocarbon, 38, 11–12. Burr, G. S., Beck, J. W., Correge, T., Cabioch, G., Taylor, F. W., and Donahue, D. J., 2009. Modern and Pleistocene reservoir ages inferred from South Pacific corals. Radiocarbon, 51(1), 319–335. Butzin, M., Prange, M., and Lohmann, G., 2005. Radiocarbon simulations for the glacial ocean: the effects of wind stress, Southern Ocean sea ice and Heinrich events. Earth and Planetary Science Letters, 235, 45–61. Chiu, T-C, Fairbanks, R. G., Cao, L., and Mortlock, R. A., 2007. Analysis of the atmospheric 14C record spanning the past 50,000 years derived from high-precision 230Th/234U/238U and 231 Pa/235U and 14C dates on fossil corals. Quaternary Science Reviews, 26, 18–36. Culleton, B. J., Kennett, D. J., Ingram, B. L., Erlandson, J. M., and Southon, J. R., 2006. Intrashell radiocarbon variability in marine molluscs. Radiocarbon, 48(3), 387–400. Cutler, K. B., Gray, S. C., Burr, G. S., Edwards, R. L., Taylor, F. W., Cabioch, G., Beck, J. W., Cheng, H., and Moore, J., 2004. Radiocarbon calibration to 50 kyr BP with paired 14C and 230 Th dating of corals from Vanuatu and Papua New Guinea. Radiocarbon, 46(3), 1127–1160. deVries, H. L., 1958. Variation in concentration of radiocarbon with time and location on Earth. Proceedings of Koninkl Ned Akademie van Wetenschappen, 61, 94–102. Druffel, E. R. M., 1981. Radiocarbon in annual coral rings from the eastern tropical Pacific Ocean., Geophysical Research Letters, 8, 59–62. Druffel, E. R. M., Robinson, L. F., Griffin, S., Halley, R. B., Southon, J. R., and Adkins, J. F., 2008. Low reservoir ages for the surface ocean from mid-Holocene Florida corals. Paleoceanography, 23, PA2209, DOI:10.1029 2007PA001527. Edwards, R. L., Beck, J. W., Burr, G. S., Donahue, D. J., Chappell, J. M. A., Bloom, A. L., Druffel, E. R. M., and

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Taylor, F. W., 1993. A large drop in atmospheric 14C/12C and reduced melting in the Younger Dryas, documented with 230 Th ages of corals. Science, 260(5110), 962–968. Eiriksson, J., Larsen, G., Knudsen, K. L., Heinemeier, J., and Simonarson, L. A., 2004. Marine reservoir age variability and water mass distribution in the Iceland Sea. Quaternary Science Reviews, 23(20–22), 2247–2268. Fairbanks, R. G., Mortlock, R. A., Chiu, T.-C., Cao, L., Kaplan, A., Guilderson, T. P., Fairbanks, T. W., and Bloom, A. L., 2005. Marine Radiocarbon Calibration Curve Spanning 10,000 to 50,000 Years B.P. Based on Paired 230Th/234U/238U and 14C Dates on Pristine Corals. Quaternary Science Reviews, 24, 1781–1796. Franke J., Paul, A., and Schultz, M., 2008. Modeling variations of marine reservoir ages during the last 45000 years. Climate of the Past, 4, 125–136. Friedrich, M., Remmele, S., Kromer, B., Hofmann, J., Spurk, M., Kaiser, K. F., Orcel, C., and Küppers, M., 2004. The 12,460-year Hohenheim oak and pine tree-ring chronology from Central Europe- a unique annual record for radiocarbon calibration and paleoenvironment reconstructions. Radiocarbon, 46(3), 1111–1122. Genty, D., Massault, M., Gilmour, M., Baker, A., Verheyden, S., and Kepens, E., 1999. Calculation of past dead carbon proportion and variability by the comparison of AMS 14C and TIMS U/Th ages on two Holocene stalagmites. Radiocarbon 41(3), 252–270. Guilderson, T. P., Schrag, D. P., Goddard, E., Kashgarian, M., Wellington, G. M., and Linsley, B. K., 2000. Southwest subtropical Pacific surface water radiocarbon in a high resolution coral record. Radiocarbon, 42(2), 249–256. Guilderson, T. P., Fallon, S., Moore, M. D., Schrag, D. P., and Charles, C. D., 2009. Seasonally resolved surface water D14C variability in the Lombok Strait: a coralline perspective. Journal of Geophysical Research, 114, C7, doi:10.1029/2008JC004876. Hogg, A. G., Turney, C. S. M., Palmer, J. G., Fifield, L. K., and Baillie, M. G. L., 2006. The potential for extending IntCal04 using OIS-3 New Zealand sub-fossil Kauri. PAGES News, 14(3), 11–12. Hughen, K. A., Baillie, M. G. L., Bard, E., Beck, J. W., Bertrand, C. J. H., Blackwell, P. G., Buck, C. E., Burr, G. S., Cutler, K. B., Damon, P. E., Edwards, R. L., Fairbanks, R. G., Friedrich, M., Guilderson, T. P., Kromer, B., McCormac, G., Manning, S., Ramsey, C. B., Reimer, P. J., Reimer, R. W., Remmele, S., Southon, J. R., Stuiver, M., Talamo, S., Taylor, F. W., van der Plicht, J., and Weyhenmeyer, C. E., 2004. Marine04 marine radiocarbon age calibration, 0–26 cal kyr BP. Radiocarbon, 46, 1059–1086. Hua, Q., Barbetti, M., Fink, D., Kaiser, K. F., Friedrich, M., Kromer, B., Levchenko, V. A., Zoppi, U., Smith, A. M., and Bertuch, F., 2009. Atmospheric 14C variations derived from tree rings during the early Younger Dryas. Quaternary Science Reviews 28(25–26), 2982–2990. Key, R. M., Quay, P. D., Jones, G. A., McNichol, A. P., von Reden, K. F., and Schneider, R. J., 1996. WOCE AMS Radiocarbon I: Pacific Ocean results; P6, P16 and P17. Radiocarbon, 38(3), 425–518. Kitagawa, H., 1995. Extension of radiocarbon calibration curve. Quaternary Research, 34, 185–190. Kitagawa, H., and van der Plicht, J., 1998. A 40,000-year varve chronology from Lake Suigetsu, Japan: extension of the C-14 calibration curve. Radiocarbon, 40(1), 505–515. Kitagawa, H., and van der Plicht, J., 2000. Atmospheric radiocarbon calibration beyond 11,900 cap BP from Lake Suigetsu laminated sediments. Radiocarbon, 42(3), 369–380. Libby, W. F., 1952. Radiocarbon Dating. Chicago: Chicago University Press.

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Libby, W. F., Anderson, E. C., and Arnold, J. R., 1949. Age determination by radiocarbon content: worldwide assay of natural radiocarbon. Science, 109, 227–228. McCormac, F. G., Hogg, A. G., Blackwell, P. G., Buck, C. E., Higham, T. F. G., and Reimer, P. J., 2004. SHCal04 Southern Hemisphere calibration, 0–11cal kyr BP. Radiocarbon 46(3), 1087–1092. Muller, R. A., 1977. Radioisotope dating with a cyclotron. Science, 196, 489–494. Nelson, D. E., Korteling, R. G., and Stott, W. R., 1977. Carbon14: direct detectin at natural concentrations. Science, 196, 507–508. Oeschger, H., Siengenthaler, U., Schotterer, U., and Gugelmann, A., 1975. A box diffusion model to study the carbon dioxide exchange in nature. Tellus, 27, 168–192. Östlund, G., and Stuiver, M., 1980. GEOSECS Pacific radiocarbon. Radiocarbon, 22, 25–53. Petchey, F., 2009. Dating marine shell in Oceania: issues and prospects. In Fairbairn, A., O’Connor, S., and Marwick, B. (eds.), New Directions in Archaeological Science. Terra Australis, Canberra: ANU E, Vol. 28, pp. 157–172. Reimer, P. J., Baillie, M. G. L., Bard, E., Bayliss, A., Beck, J. W., Bertrand, C. J. H., Blackwell, P. G., Buck, C. E., Burr, G. S., Cutler, K. B., Damon, P. E., Edwards, R. L., Fairbanks, R. G., Friedrich, M., Guilderson, T. P., Hogg, A. G., Hughen, K. A., Kromer, B., McCormac, F. G., Manning, S. W., Ramsey, C. B., Reimer, R. W., Remmele, S., Southon, J. R., Stuiver, M., Talamo, S., Taylor, F. W., van der Plicht, J., and Weyhenmeyer, C. E., 2004. IntCal04 Terrestrial radiocarbon age calibration, 26–0 ka BP. Radiocarbon, 46, 1029–1058. Reimer, P. J., Baillie, M. G. L., Bard, E., Bayliss, A., Beck, J. W., Blackwell, P. G., Bronk Ramsey, C., Buck, C. E., Burr, G. S., Edwards, R. L., Friedrich, M., Groots, P. M., Guilderson, Hajdas, I., Heaton, T. J., Hogg, A. G., Hughen, K. A., Kaiser, K. F., Kromer, B., McCormac, F. G., Manning, S. W., Reimer, R. W., Richers, D. A., Southon, J. R., Talamo, S., Turney, C. S. M., van der Plicht, J., and Weyhenmeyer, C. E., 2009. IntCal09 and MARINE09 radiocarbon age calibration curves, 0–50,000 years CAL BP. Radiocarbon, 51(4), 1111–1150. Reimer, P. J., and Reimer, R. W., 2009. Marine Reservoir Correction Database. http://calib.org/marine Schaub, M., Buntgen, U., Kaiser, K. F., Kromer, B., Talamo, S., Andersen K. K., and Rasmussen, S. O., 2008. Late glacial environmental variability from Swiss tree rings. Quaternary Science Reviews, 27(1–2), 29–41. Suess, H. E., 1965. Secular variation of the cosmic-ray produced carbon-14 in the atmosphere and their interaction. Journal of Geophysical Research, 70, 5937–5952. Suess, H. E., 1980. The radiocarbon record in tree rings of the last 8000 years. Radiocarbon, 22, 200–209. Stuiver, M., 1971. Evidnce for the variatin of atmospheric 14C content in the Late Quaternary. In Turekian, K. K. (ed.), The Late Cenozoic Glacial Ages. New Haven: Yale University Press. Stuiver M., and Polach, H. A., 1977. Discussion: Reporting of 14C Data. Radiocarbon, 19(3), 355–363. Stuiver, M., Reimer, P. J., Bard, E., Beck, J. W., Burr, G. S., Hughen, K. A., Kromer, B., McCormac, G., van der Plicht, J., and Spurk, M., 1998. INTCAL.98 radiocarbon age calibration, 24,000–0 cal BP. Radiocarbon, 40, 1041–1083. Stuiver, M., Reimer, P. J., and Reimer, R., 2005. Radiocarbon calibration program revision 5.0.1. Copyright 1986–2005. Stuiver M., 1970. Long-term 14C variations. In Olsson, I. U. (ed.), Radiocarbon Variations and Absolute Chronology, 12th Nobel Symposium New York: Wiley, pp. 197–213. Stuiver, M., and Braziunas, T. F., 1993. Modeling atmospheric 14C influences and 14C ages of marine samples to 10,000 BC. Radiocarbon, 35(1), 35–65.

Stuiver, M., Reimer, P. J., and Braziunas, T. F., 1998. High-precision radiocarbon age calibration for terrestrial and marine samples. Radiocarbon, 40, 1127–1151. Turney, C. S. M., Fifield, L. K., Palmer, J. G., Hogg, A. G., Baillie, M. G. L., Galbraith, R., Ogden, J., Lorrey, A., and Tims, S. G., 2007. Towards a radiocarbon calibration for oxygen isotope stage 3 using New Zealand Kauri (Agathis australis). Radiocarbon, 49(2), 447–457.

Cross-references Uranium Series Dating

RECENT SEA LEVEL TRENDS Philip L. Woodworth National Oceanography Centre Liverpool, Liverpool, UK

Synonyms Sea level; Sea surface height; Still water level Definitions Sea level. Height of the sea relative to the height of a benchmark on the nearby land (in tide gauge data analysis). Height of the sea relative to a reference surface such as an ellipsoid or geoid (in altimeter data analysis). Mean sea level. Sea level averaged over a period of time, such as a month or year, long enough that fluctuations caused by waves and tides are largely removed (in tide gauge data analysis). Still water level. The sea level that results from ocean processes including tides, surges, and mean sea level changes but not including waves. Sea surface height. A term synonymous with “sea level” but used mostly in altimeter data analysis. Sea level trends. Rates of change of sea level over a stated period either for a given location or averaged over a region or the global ocean. Introduction Sea levels have been measured with what are usually called “tide gauges” for many hundreds of years, both for practical purposes, such as efficient operation of a harbor, or for scientific research. For example, the first extended set of sea level measurements in the UK comprised heights and times of high water at Liverpool during 1764–1793 recorded by William Hutchinson (Woodworth, 1999). By the 1830s, the automatic tide gauge had been developed, consisting of a float in a stilling well and with the float’s motion recorded on a paper chart. The credit for this development is often given to Palmer following his installation of a gauge at Sheerness, UK in 1832 (Pugh, 1987). In the 1970s, the first satellite radar altimeters were launched. Altimeters record the propagation time of radar pulses transmitted

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Global sea level rise 150 Church and White (2009) 100

Holgate and Woodworth (2004) Jevrejeva et al. (2006)

50 Sea level (mm)

from the satellite, reflected off the ocean, and received back at the satellite. Sea levels can be measured by this means with an accuracy of several centimeters and on a near-global basis (Fu and Cazenave, 2001). This technique attained optimal reliability and accuracy with the deployment of the TOPEX/POSEIDON (T/P) mission in 1992. Since that time, the community has had nearcontinuous availability of precise altimeter data from T/P and JASON-1 and -2, complemented by data sets from the lower-flying ERS-1 and -2 and ENVISAT satellites. Tide gauges and altimeters have together demonstrated that sea levels change over many different temporal and spatial scales.

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Recent sea level trends Since the last ice age around 20,000 years ago, mean sea level (MSL) has risen worldwide by more than 120 m. Over the last few millennia, the rate of rise has been no more than a few tenths of mm/year. However, an acceleration appears to have taken place between the eighteenth and nineteenth centuries and into the twentieth century, based on the small number of available long European tide gauge records (Woodworth, 1999; Jevrejeva et al., 2008) and on complementary data from salt marshes (e.g., Gehrels et al., 2006). A consensus has emerged that the twentieth century rise in global sea level was closer to 2 than 1 mm/year, with values around 1.7 mm/year having been obtained recently for the past century (Cazenave and Nerem, 2004; Church and White, 2006) or past halfcentury (Church et al., 2004; Holgate and Woodworth, 2004). However, the rate of change was far from constant, with an acceleration around 1920–1930, a deceleration after 1960, and a relatively recent acceleration in the late 1980s or early 1990s (e.g., Woodworth et al., 2009; Douglas, 2008; Merrifield et al., 2009). The high rate in the latter period of over 3 mm/year was observed not only by tide gauges but also by satellite altimetry (Beckley et al., 2007). It has been suggested this high rate may be connected partly with a sea level recovery that started in the mid-1980s following a sea level fall after the El Chichón volcanic eruption (Domingues et al., 2008). Figure 1 presents findings on global sea level rise by different authors, and there is general agreement between them. However, it must always be kept in mind that our knowledge of recent change is based largely on a sparse data set, that of the Permanent Service for Mean Sea Level (PSMSL) (Woodworth and Player, 2003), and that there can be considerable spatial and temporal variability in sea level change (cf. Figure 1 of Wunsch et al., 2007 for a demonstration of spatial variability in trends since the early 1990s). Sea level time series at particular locations, including coral islands, can be found in individual reports (e.g., Woodworth, 2005) and from the PSMSL web site www.psmsl.org. Vertical land movements (VLMs) due to glacial isostatic adjustment (GIA), tectonic or anthropogenic processes can play important roles in the sea level changes observed by tide gauges and by other in situ

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Recent Sea Level Trends, Figure 1 Global sea levels from three recent studies, all based on MSL data from the PSMSL. Note the higher rates of rise after 1930 and in recent years. Values denoted Church and White (2009) are updated from those reported in Church and White (2006).

techniques. VLMs due to GIA can be estimated with the use of geodynamic models (e.g., Peltier, 2001), while advanced geodetic techniques (notably Global Positioning System and Absolute Gravity) are being developed to monitor VLMs, of whatever origin, at tide gauge sites (Woodworth, 2006).

Understanding the observed sea level rise MSL is often described as an “integral parameter”, providing an integration of sea level changes resulting from many climate-change related processes over many years, of which oceanic thermal expansion, melting of glaciers, ice caps and ice sheets, and modifications to hydrological exchanges between land and ocean are the most important. Consequently, if MSL is changing, it points to major changes in one or more of the climate-related drivers of that change. Ice sheets have the potential to raise sea level by far more than the other processes (see Table 11.3 of Church et al., 2001), although their contribution during the past century appears to have been small. The composition of the various contributions to sea level change (the “sea level budget”) forms a major research topic which is addressed regularly in the scientific assessments of the Intergovernmental Panel on Climate Change (IPCC), the most recent being the Third Assessment Report (TAR) (Church et al., 2001) and Fourth Assessment Report (AR4) (Bindoff et al., 2007). Figure 2 summarizes the budget calculated by the AR4. The inability of these assessments to account for most of the observed rise (the “sea level enigma” of

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Budget of global sea level rise Thermal expansion

Glaciers and ice caps

Greenland

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Recent Sea Level Trends, Figure 2 Contributions to sea level rise (upper four entries), their sum, and the observed rates of sea level rise for 1961–2003 (red) and 1993–2003 (blue). The bars represent the uncertainty range. The difference (Obs – Sum) represents the “sea level enigma.” For details, see Bindoff et al. (2007).

Munk, 2002) has to some extent been resolved, at least for recent decades, by a reworking of hydrographic data sets and through the availability of more complete space and in situ instrumentation for the determination of thermal expansion (Domingues et al., 2008). At the time of writing, “budget estimates” for the last decade, during which the community has had available information from satellite gravity [i.e., the Gravity Recovery And Climate Experiment (GRACE) mission] and Argo hydrography in addition to altimeter and tide gauge data, have been more consistent, although discrepancies remain (Willis et al., 2008; Cazenave et al., 2009). One is confident that consensus will eventually be achieved as global observation systems become more complete. Reviews of the field have been undertaken by individual scientists (e.g., Cazenave and Nerem, 2004; Woodworth, 2006) and national and international study groups in between the IPCC assessments. One such major review of the status of research was undertaken by the World Climate Research Program (WCRP) in 2006 (Church et al., 2007) from which full reports on each sector of the field will be published in book form in 2010 (Church et al., 2010). One chapter is concerned with the use of sea level information derived from coral reefs, in addition to salt marsh and other coastal geological data, as a complement to and extension back in time of the instrumental record (Lambeck et al., 2010).

Changes in extreme sea levels As extreme sea level events often result in flooding and loss of life, an important question is whether their amplitudes and frequencies are changing, and if the levels of extreme high waters are changing in a significantly different way to MSLs. The only study which has attempted a quasi-global investigation of this topic is that of Woodworth and Blackman (2004) who studied data from 141 stations, and concluded that there is indeed evidence for an increase in extreme high water levels worldwide since 1975, as reported frequently in the press. However, in most cases, the secular changes and the interannual variability in the extremes were found to be similar to those in MSL. A number of other studies of sea level extremes at particular locations are available, although as they are for different epochs and use different methods, it is difficult to arrive at general conclusions. A review of these studies and of the meteorological forcing factors which result in extreme levels has been provided by Lowe et al. (2010). Predictions of future sea level trends The IPCC assessments also provide regular updates on future sea level rise predictions based on climate modeling. The TAR projected a global averaged sea level rise of between 20 and 70 cm between 1990 and 2100 using

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the full range of IPCC greenhouse gas scenarios and a range of climate models. When an additional uncertainty for land-ice changes was included, the full range of projected sea level rise was 9–88 cm. For the IPCC AR4 (Meehl et al., 2007), the range of sea level projections using a larger range of models was 18–59 cm (90% confidence limits) over the period from 1980–1999 to 2090–2099. The largest contribution was from ocean thermal expansion with the next largest contribution from glaciers and ice caps. However, there is increasing concern about the stability of ice sheets, and recognizing this deficiency, the AR4 increased the upper limit of the projected sea level rise by 10–20 cm, implying an overall range of projected sea level rise of 18–79 cm. It is unclear what confidence intervals to assign to this range given the ice sheet uncertainties. Note that they also stated that “larger values cannot be excluded, but understanding of these effects is too limited to assess their likelihood or provide a best estimate or an upper bound for sea level rise.” While the 2001 and 2007 IPCC projections are somewhat different in how they treat ice sheet uncertainties and the confidence limits quoted, a comparison of the projections (Figure 6 of Church et al., 2008) shows the end results are similar, except that the lower limit of the 2001 projections has been raised from 9 cm in the TAR to 18 cm in the AR4. Despite the additional allowance for ice sheet uncertainties, a number of scientists remain concerned that the ice-sheet contributions in the AR4 may have been underestimated, and they adopt a more phenomenonological approach to estimating future sea level rise. For example, Rahmstorf (2007) developed a simple statistical model that related twentieth century surface temperature change to twentieth century sea level change. Using this relationship and projected surface temperature increases, he estimated that twenty-first century sea level rise might exceed the IPCC projections and be as much as 1.4 m. Holgate et al. (2007) raised concerns that Rahmstorf’s model is too simplistic and may not adequately represent future change. Similar conclusions, that the AR4 sea level rise predictions may have been underestimated have been based on analysis of longer term temperature and sea level information (Grinsted et al., 2009). One remains concerned at the low physics content of some of these parameterizations. Nevertheless, the concern that the IPCC sea level projections may be biased low has been reinforced by a comparison of observed and projected sea level rise from 1990 to the present. For this period, observed sea level has been rising at the very upper end of the IPCC TAR projections (Holgate and Woodworth, 2004; Rahmstorf et al., 2007), indicating once more that one or more of the model contributions to sea level rise may be underestimated. Future IPCC assessments are likely to pay greater attention to regional changes in MSL, as opposed to globalaverage ones, and include in impact studies the role of regional land movements. In addition, they will study in more detail changes in climate extremes (extreme sea

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levels leading to coastal flooding in this case) as well as in climate means (MSL in this case). For example, the decimetric twentieth century sea level rise has already doubled the risk of flooding since 1901 at many locations around the UK coastline. If sea level rises faster in the twenty-first century, as suggested by the AR4, there will be greater risk to the coastal environment and infrastructure. In addition, climate-related changes in regional wind fields will result in changes in storm surge frequency and magnitude which will modify the changes in risk due to MSL rise alone (Woodworth et al., 2007; Lowe et al., 2010).

Need for ongoing monitoring of sea level Descriptions of global and national sea level networks and their data sets can be found at the web sites of the PSMSL and the Global Sea Level Observing System (GLOSS, www.gloss-sealevel.org), while methods for monitoring sea level are explained in the PSMSL training web pages. Real-time data provision, in addition to the delayed-mode data needed for scientific research, is increasingly encouraged for two reasons: the data are then available to a wide range of new users in “operational oceanography” including coastal protection; and faults can be identified faster, leading to better delayed mode data sets in the long term. Some sea level stations are now “multi-hazard” sites, with sensors specifically designed for the high rate recording needed for tsunami monitoring. It is clear that the world needs a coordinated and complete sea level monitoring network at global, regional, and local scales, including information both from in situ sources and from a range of space-based instrumentation (notably altimetry and space gravity, Wilson et al., 2010). Bibliography Beckley, B. D., Lemoine, F. G., Lutchke, S. B., Ray, R. D., and Zelensky, N. P., 2007. A reassessment of global and regional mean sea level trends from TOPEX and Jason-1 altimtery based on revised reference frame and orbits. Geophysical Research Letters, 34, L14608, doi:10.1029/2007GL030002. Bindoff, N., Willebrand, J., Artale, V., Cazenave, A., Gregory, J., Gulev, S., Hanawa, K., Le Quéré, C., Levitus, S., Nojiri, Y., Shum, C., Talley, L., and Unnikrishnan, A., 2007. Observations: oceanic climate change and sea level. In Solomon, S., Qin, D., and Manning, M. (eds.), Climate Change 2007: The Physical Science Basis. Contribution of Working Group 1 to the Fourth Assessment Report of the Intergovernmental Panel on Climate Change. United Kingdom and New York: Cambridge University Press. Cazenave, A., Dominh, K., Guinehut, S., Berthier, E., Llovel, W., Ramillien, G., Ablain, M., and Larnicol, G., 2009. Sea level budget over 2003–2008: a reevaluation from GRACE space gravimetry, satellite altimetry and Argo. Global and Planetary Change, 65, 83–88, doi:10.1016/j.gloplacha.2008.10.004. Cazenave, A., and Nerem, R. S., 2004. Present-day sea level change: observations and causes. Reviews of Geophysics, 42, RG3001, doi:10.1029/2003RG000139. Church, J. A., Gregory, J. M., Huybrechts, P., Kuhn, M., Lambeck, K., Nhuan, M. T., Qin, D., and Woodworth, P. L., 2001. Changes in sea level. In Houghton, J. T., Ding, Y., Griggs, D. J., Noguer, M., van der Linden, P., Dai, X., Maskell, K., and Johnson, C. I. (eds.), Climate Change 2001:

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Cross-references Climate Change: Impact of Sea Level Rise on Reef Flat Zonation and Productivity Sea Level Change and Its Effect on Reef Growth Sea-Level Indicators

RED SEA AND GULF OF AQABA

RED SEA AND GULF OF AQABA Yonathan Shaked1, Amatzia Genin1,2 1 The Interuniversity Institute of Marine Sciences, Eilat, Israel 2 The Hebrew University of Jerusalem, Jerusalem, Israel

Definition The Red Sea is a narrow elongated water body extending some 2,000 km SE–NW, between latitudes 16 N and 28 N, from the Gulf of Aden through which it connects from the Gulf of Aden (through which it connects to the Indian Ocean) to the Gulf of Aqaba (Figure 1). It is flanked by east Africa (Egypt, Sudan, Eritrea, and Djibuti) to the west and Arabia (Saudi Arabia and Yemen) to the east. The Red Sea hosts spectacular coral reefs that pertain to the Indo-Pacific domain. Its northern tributaries, the Gulfs of Suez and Aqaba (reaching latitude 29.5 N) are home to some of the northernmost coral reefs in the world.

Red Sea And Gulf Of Aqaba, Figure 1 Satellite image of the Red Sea and its northern tributaries, the Gulfs of Suez and Aqaba.

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Red Sea coral reefs – geology and ecology Setting An incipient oceanic basin separating the African Plate from the Arabian Plate, the Red Sea was born as an extensional basin during the late Oligocene – early Miocene, with actual sea floor formation occurring since about 5 million years ago at its southern part. The Red Sea is connected to the Indian Ocean by way of the shallow and narrow sill of Bab-el-Mandeb (137 m deep and 29 km wide) and the Gulf of Aden. From its southern end at Bab-el-Mandeb, the Red Sea extends nearly 2,000 km until it diverges into two smaller gulfs, the Gulf of Suez and the Gulf of Aqaba. This junction connects the mostly extensional Red Sea with the abandoned rift of the Gulf of Suez and the mostly leftlateral Dead Sea transform of which the Gulf of Aqaba is an active part. Rifting processes resulted in block tilting and an asymmetrical bathymetry across the Red Sea. The Egyptian–Sudanese (western) margin is steep with a narrow shelf and deep pull-apart basins floored by oceanic crust, whereas the Arabian (eastern) margin is less steep and floored by attenuated continental crust. The Red Sea proper is up to 350 km wide and has a maximum depth of 2,800 m, but its average depth is only
450 m, since around 40% of the Red Sea is shallower than 100 m. The Gulf of Suez is about 320 km long and 70 km wide, with a maximum depth of less than 100 m, while the Gulf of Aqaba is 180 km long and up to 25 km wide with an average depth of 900 m and a maximum depth of nearly 1,850 m. The margins of the sea on land form elevated shoulders reaching up to 3,000–4,000 m, a few tens of kilometers away from the shore. The coasts are generally narrow, with little or no coastal plain, and sandy beaches are found only in the vicinity of creek valley outlets. The Red Sea is surrounded by deserts and evaporation greatly exceeds precipitation. Freshwater input from the continent is limited to occasional flash floods. Reefs – morphology, distribution, and description Warm temperatures and little runoff support reef growth, and active tectonics create steep margins that are often coral-covered to a depth of several dozen meters. On shallower slopes and wider margins, shallow reef shoals and a few coral islands develop. Fringing reefs are the most common form of coral terraces along the coasts of the Red Sea (Head, 1987), but some patch reefs and coral islands also occur, especially off the Sudanese coast and the southern parts of both the western and eastern coasts. In general, reefs are better developed along the northern part of the Red Sea down to about latitude 18–20 N (Behairy et al., 1992), although coral growth rates were found to be higher in the south, decreasing northwards (Schuhmacher et al., 1995). To the south, sandy beaches become increasingly abundant and mangroves occupy parts of the coasts. At the southern parts of the Red Sea, reefs of coralline algae are also found

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(Behairy et al., 1992). Fringing reefs on the northern parts of the Red Sea are generally narrow, being limited by the steep coastal margins, and form deep “walls”. To the south, the coastal plain is wider and the margins are less steep, allowing the development of wider lagoons and reefs further from shore. Patch reefs somewhat removed from the shores may form chains resembling barrier reefs. Along the coast from Al-Wadj south to Jeddah (eastern shore), a long chain of patch reefs was called the “Little Barrier Reef ” by Sheppard (1985). Further south, groups of coral-supporting islands are found on both sides of the sea – the Dahlak and Farasan archipelagoes off the coast of Eritrea and Saudi Arabia, respectively. Some of the reefs growing on wide shallow carbonate platforms in the central and southern parts of the Red Sea have a circular shape enclosing a lagoon, thus resembling Pacific atolls. The best studied of these is perhaps Sanganeb Atoll, off Port Sudan (Schuhmacher and Mergner, 1985; Mergner and Schuhmacher, 1985; Reinicke et al., 2003). Many of the modern coral reefs along the Red Sea grow over fossil Pleistocene reefs. Fringing coral reefs are also the overwhelmingly dominant form along the shores of the Gulfs of Suez and Aqaba, where the former are smaller and less developed than the latter. The difference may be due to the low winter temperatures at the shallow Gulf of Suez. At greater depths, down to at least 65–70 m, coral carpets covering marginal slopes are abundant. Growth of these deep reefs is facilitated by the clarity of the water that allows light penetration to a depth of
100 m.

Geological history of coral reefs at the Red Sea Coral reefs seem to have been intermittently present along the shores of the Red Sea since Miocene times, persisting during periods in which good connection with the open ocean was maintained. Connection of the Red Sea with the Mediterranean finally ceased during the Pliocene, and from that period Indo-Pacific flora and fauna were introduced (Braithwaite, 1987). During glacial periods, the combination of low temperatures, high evaporation, and restricted water passage over the shallow sill at Bab-el-Mandeb due to low sea level possibly prevented the development of coral reefs. Conditions at the northern Red Sea and Gulf of Aqaba during the last glacial maximum are estimated at 4–5 C lower than present-day temperatures and salinity in the range of 50–53% (Arz et al., 2003; Siddall et al., 2003; Almogi-Labin et al., 2008). Fossil Pleistocene coral reefs are found along the Red Sea coasts on both sides, off of Sudan and Saudi Arabia. These reach an elevation of at least 16 m above current sea level. Raised fossil coral reefs are also present at the Gulfs of Suez and Aqaba, with terraces at the Gulf of Aqaba, dating as far back as 300 ka, reaching some 35 m above sea level (Figures 2, 3; Gvirtzman et al., 1992). Fossil sequences of 3–4 reef complexes are found along the southern coast of Sinai and at the northeastern end of the Gulf of Aqaba, just south of the city Aqaba (Friedman, 1968; Gvirtzman et al., 1992; Al-Rifaiy and Cherif, 1988). The elevated Pleistocene sequences seem to represent Pleistocene highstands, reaching their present elevation as a result of uplifting tectonic movements along

Red Sea And Gulf Of Aqaba, Figure 2 Uplifted fossil Pleistocene coral reefs covered by beach rock some 20 m above the sea. South of Aqaba, on the northeastern side of the Gulf of Aqaba. (Photo courtesy by Y. Shaked.)

RED SEA AND GULF OF AQABA

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Red Sea And Gulf Of Aqaba, Figure 3 Fossil corals comprising uplifted Pleistocene reefs south of Aqaba. (Photo courtesy by Y. Shaked.)

those margins of the Gulf. On Tiran island at the entrance to the Gulf of Aqaba, where uplifting tectonics have particularly high rates, marine coastal terraces are found up to 500 m above the present-day sea level. Fossilized late Pleistocene fringing reefs are recognized to an elevation of 60 m above sea level, where an age of 146 ka was determined from a terrace at 40 m (Goldberg and Beyth, 1991). Submerged terraces, such as those found at
120 and 65–70 m depth along the northwestern end of the Gulf of Aqaba (Makovsky et al., 2008), likely represent lower sea level stand-stills. A reliable age control of these submerged features is not yet available. Since conditions at the Red Sea and its subsidiary gulfs were not favorable to reef formation during glacial periods, it seems that most modern reefs were established during the Holocene (Braithwaite, 1987). A coral from a fossil reef retrieved from 11 m below sea level yielded an age of 7,000 years (Shaked et al., 2004). Slightly exposed reefs along both margins of the Gulf of Aqaba indicate a sustained (ca. 7–4 ka) Holocene sea level highstand,
1 m above present sea level. During this interval Holocene fringing reefs developed morphological reef flats that are now found at or just above the high-tide level, commonly partly eroded and covered with beach rock (Shaked et al., 2004, 2005). The modern reefs are found further offshore and are limited by the present sea level.

Present conditions The Red Sea’s deep, elongated morphology coupled with its location within a hyper arid region create a distinct antiestuarine circulation at Bab-el-Mandeb, where relatively warm water (25.5–30 C) with normal salinity (36.5%) enters the Red Sea from the Gulf of Aden, forming the

surface water layer that cools and evaporates along the route northward. The Red Sea bottom water (21.7 C, 40.6%) is formed mostly in the Gulf of Aqaba (Plähn et al., 2002), flowing into the Red Sea over the shallow (250 m) and narrow (16 km) sill in the Straits of Tiran. The presence of a shallow sill at Bab-el-Mandeb greatly affects the extent and diversity of the coral reefs in the Red Sea, especially those at higher latitudes in the Gulfs of Aqaba and Suez. Unlike “normal” open oceans, where the deep and intermediate water is cold, the shallow sill at Bab-el-Mandeb effectively separates the deep waters in the semi-enclosed Red Sea from those of the Indian Ocean. For example, the water at 1,500 m depth in the Gulf of Aden (the water body connecting the Red Sea to the Indian Ocean) is n discrete habitat patches and dispersion described by n x n (Mij) containing the probability of larval dispersal from each of the patches. Each row is a discrete version of the DK for the patch i

A B C D A B C D Source population

(Modified from Botsford et al., 2009, Coral Reef)

Reef Interconnectivity/Larval Dispersal, Figure 1 Relationship between dispersal and connectivity.

Influence of hydrodynamic and biological process In idealized coastal models described by Siegel et al. (2003), the dispersal kernel only accounts for advection by stationary mean currents and turbulent diffusion, determined by the temporal scales of current fluctuations. While diffusion tends to increase local retention over multiple generations, advection increases the spread of the species. The relative importance of the two processes can be defined by the Peclet number (Pe). Recent studies indicate that despite their small size, coral reef fish larvae are not passive (Leis, 2007 for review). They have a diversity of traits and considerable behavioral capabilities that can lead to successful completion of the early pelagic life phase (Paris et al., 2007). In particular, vertical migration during ontogeny increases retention near natal reefs and decreases dispersion losses, likely enhancing survival (Paris and Cowen, 2004). Larvae can come back to their native reefs or can be exchanged among breeding subpopulations. However, the dynamics of these interactions at both the individual and population levels are not fully understood (deYoung et al., 2004). For those larvae that do not return home, the extent to which larval behavior influences their arrival pattern among adjacent and distant reefs (or larval connectivity network) is not known. More importantly, interactions of the small-scale larval movements with transport

processes due to larger-scale currents need to be quantified with regard to the spatial patterns of recruitment. Since actively moving larvae may be diluted and to some extent dispersed by currents, an in situ study of them is very difficult. Larval behavior is associated with the perceptual range of each individual, representing its informational window onto the environment. For coral reef fish larvae in particular, a suite of sensory systems operating at different scales can play an important role in orientation: acoustic (Simpson et al., 2005) and olfactory (Atema et al., 2002) signals in the shorter range, and possibly visual and magnetic signals (Kingsford et al., 2002) for longer-range navigation. Understanding the orientation behavior of larvae is a prerequisite to understanding, measuring, and predicting demographic (ecological) connectivity and may also be relevant to genetic (evolutionary) connectivity (Paris et al., 2009). Understanding how larvae orientate and the sensory cues they use for orientation constitutes one of the largest gaps in our knowledge of larval behavior.

Reef interconnectivity and habitat Population connectivity not only depends on life history but also on the seascape defined by both the local currents and habitat configuration. The interactions between physical and biological factors and their role in shaping

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populations have been previously discussed in landscape ecology (Levins, 1969), but this discussion is a relatively recent development for marine ecosystems (Barber et al., 2002; Baums et al., 2006; Cowen et al., 2006). The spatial arrangements and connectivities of marine populations are poorly understood, yet they are assumed to enhance resilience to disturbance and be of critical importance for population persistence (Kinlan et al., 2005; Hasting and Botsford, 2006). Baums et al. (2006) demonstrated that patterns of connectivity are indeed linked to the fragmentation of the coral reef habitat. Connectedness (i.e., number of connected subpopulations) and connectivity networks are changed with extended larval competency period only when distances between habitat patches are significant. This implies that the response from reef populations to climate change through the growth rate and survival of the early life history stages is a nonlinear process and may be counterintuitive. Indeed, dispersal kernels are not necessarily shrinking with increased temperature (Paris et al., 2008; Munday et al., 2009). Maintaining the reef system as pristine as possible is important since it affects the production of larvae and their settlement survival, and subsequently, their recruitment into the adult population (i.e., “realized connectivity” sensu Pineda et al., 2007) and resilience in response to disturbances. Reef damage also impacts reef interconnectivity through the loss of habitat patches that may serve as stepping stone linkages for the population network. Reefs depend on pristine areas for larval production, which is necessary for long-term persistence. If pristine areas are not maintained, eventually the entire reef will likely be lost. Coral reef conservation would benefit from greater understanding of the degree of connectivity between coral reefs, or between land and reefs, that may exist through the movement of water masses (Paris and Cherubin, 2008). Land use, such as clearing of native vegetation and replacement with intensive agriculture, has increased pollution transport to the coastal ocean to a level many times the natural rate (McKergow et al., 2005). Degraded water quality of streams has yielded degradation of receiving bodies, such as lagoons and coastal waters, where most coral reefs reside. Sediments and nutrients, in excess, are almost universally recognized as having inhibitory or negative effects on reef communities.

Estimating reef interconnectivity Larval dispersal is described by the number of larvae N dispersed as a function of the distance y from the birth place. Another way to represent dispersion is to describe a dispersal kernel representing the probability density function (PDF) to find a larva at a distance y from its birth location. In other words, it is the probability that a larva settles at a given distance from its source population. Measures of natal dispersal (sensu Clobert et al., 2001) are typically determined by the dispersal kernel K(x,y), defined as the probability of a larva settling at a distance y, given that

it was released at a spawning location x. Dispersal kernels are thus the spatial distributions of dispersed larvae and can be represented in two or three dimensions. The modal dispersal distance from the dispersal kernel has demographic relevance (e.g., population spatial pattern, persistence), while the tail, representing long-distance dispersal, is relevant on an evolutionary level (e.g., genetic mixing, species persistence; Hanski and Gaggiotti, 2004; Steneck et al., 2006). Currently, the spatial scales of the dispersal of reef fish larvae are estimated using indirect and empirical techniques (Thorrold et al., 2002; Jones et al., 2005; Planes et al., 2009) or modeling approaches (Cowen et al., 2000, 2006; James et al., 2002).

Empirical approaches Reef interconnectivity can be empirically estimated by indirect (i.e., population genetics and paternity analysis) or by direct approaches (i.e., tagging/natural markers and larval/settlement observations). Although they differ in the characteristics of the dispersal that they can quantify, empirical methods can provide evidence of (1) connectivity, or the strength of the paths between subpopulations (or whether a subpopulation is a source or a sink), (2) local retention, or larvae returning to their natal population (sensu Paris and Cowen, 2004), and (3) self-recruitment, or strength of replacement paths which are closed loops, typically calculated as the proportion of settlers at a location that were spawned locally (Jones et al., 1999; Swearer et al., 1999; Almany et al., 2007). The latter is a measure of isolation of a subpopulation. If almost all populations within the metapopulation exhibit high self-recruitment, then they will also be characterized by high local retention and narrow dispersal kernels (Botfords et al., 2009). Among the indirect methods, the most commonly used is population genetics that describes connectivity by comparing allele frequencies among spatially discrete subpopulations. High levels of genetic similarity between populations suggests high gene flow over time. This is a valuable approach in assessing patterns and degrees of connectivity when methods to directly track larvae are not possible. Indeed, adult populations represent an accumulation of genetic signature from larval sources over time, influenced not only by larval exchange but also by ecological and evolutionary forces. Thus, such estimates of gene flow are an adequate approximation of contemporary levels of genetic connectivity only for species with relatively short life span and strictly sexual reproduction. For reef organisms with overlap across generations, population genetics may not represent present-day connectivity patterns, even for high evolving genes (Hughes et al., 2003). It is also important to note that traditional population genetic F-statistics that have been viewed as a proxy for dispersal over evolutionary time scales are not sensitive to recent changes in gene flow and genetic structure of long-lived organisms that retain the signature of past

REEF INTERCONNECTIVITY/LARVAL DISPERSAL

events. Other population genetics statistics involving assignment methods can evaluate contemporary connectivity rates without the unrealistic assumptions required by traditional methods (reviewed in Manel et al., 2005). While genetic assignment techniques can link a settler to its natal reef, parentage analysis can identify its actual parent (Jones et al. (2005) provide a promising first example). Parentage analysis yields information on the critical features of dispersal for multiple spatial scales, (Figure 2). The main limitation is performing adequate sampling and collecting a significant proportion of the source adult population and new recruits over the entire potential range of dispersal (although less critical, Planes et al., 2009). Besides genetic methods, direct estimates are possible using geochemical tags recorded from the environment in calcified structures of marine larvae (i.e., otoliths, statoliths, shells). This technique can be used to determine the origin of individuals settling into juvenile habitats, or to examine natal homing of spawning adults that migrate significant distances after settlement (Becker et al., 2007; Thorrold et al., 2001). This approach provides sufficient information to estimate the larval dispersal matrix but is not appropriate to estimate dispersal kernels due to the

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relatively large spatial scales of variability of water properties in coral reef environments (Thorrold et al., 1998, 2001). Finally, larval settlement tracking combined with observations of the chemical and physical environments provide accurate estimates of local retention, while allowing quantification of the role of larval behavior in shaping dispersal kernels (critical to calibrate conceptdriven biophysical connectivity models). However, this approach requires extensive high-frequency sampling relevant to larval time and space scales (Paris and Cowen, 2004). In addition, local retention is difficult to measure as the reproductive output and the fate of all offspring from a particular population must be known.

Numerical modeling approaches Spatially explicit individual based modeling (IBM) has emerged as a key tool for understanding organism–environment interactions (Werner et al., 2001) and is particularly relevant to investigate larval fish fluxes in the complex coral reef ecosystem. Connectivity models aim at predicting the rate of exchange of individuals (i.e., larval fluxes) between the populations forming a metapopulation. Spatially explicit IBMs have become

Kimbe Island

D

C

a Heusner Cape Heusner

Kimbe Bay

6%

South Bay Reef 100%

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Kimbe Island

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10% Restort Island

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b Kimbe Bay

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Reef Interconnectivity/Larval Dispersal, Figure 2 Empirical estimates of multi-scale larval linkages using DNA parentage analysis of (a) the clownfish (Amphiprion percula), (b) between 5 lagoons of Kimbe Is. and (c) from Kimbe Is. to marine reserves (dotted red boxes) proposed in western Kimbe Bay, (d) northeast of Papua New Guinea. Modified from Planes et al. (2009, PNAS).

Source population (i)

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Sink population ( j) Dispersal probability matrix (P)

0 (km)

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Reef Interconnectivity/Larval Dispersal, Figure 3 Several types of matrices.

B

Distance matrix (D)

Sink loaction ( j)

(Modified from Cowen et al., 2006, Science)

•

•

•

The distance matrix D = dij represents the distances between habitat patches ij The dispersal probability matrix P = pi represents the probability of migration between population patches ij The adjacency matrix A = aij represents significant larval linkages between nodes, where aij = I if node i and j are connected, otherwise aij = 0.

Connections between nodes (population patches) are represented by several types of matrices:

Adjacency matrix (A)

Prop. Surv. > 0.1 0.075 < Prop. Surv. < 0.1 0.05 < Prop. Surv. < 0.075 0.025 < Prop. Surv. < 0.05

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REEF INTERCONNECTIVITY/LARVAL DISPERSAL

the most efficient tools in connectivity studies (Werner et al., 2001). The typical output for n populations is an n  n matrix in which element (i, j ) is the probability for an individual to transit from i to j. These square matrices are called connectivity matrices or transition probability matrices, each of whose rows (i) contain numbers summing to 1. In order to describe a system at ecological scales, the proportion of successful recruits must reflect the recruitment rates (i.e., number of recruits per generation) required to replenish the local population to a minimum of zero growth (Cowen et al., 2006). Such recruitment rates can be estimated a posteriori to match adult mortality rates using simple population growth models (e.g., Nt = Nt–1 ert). Similarly, demographic connectivity models can be a posteriori scaled by production (e.g., relative spawning biomass per unit population or proportion of adult habitat in each population). Because connectivity models are by nature spatially explicit, it is strongly recommended to couple the Lagrangian tracking algorithm with a geographic information system (GIS). The GIS serves to delineate the source populations as well as the recruitment habitat along an individual particle’s path. It is also important to incorporate the uncertainties into the connectivity model (e.g., stochastic Lagrangian model (LSM), stochastic mortality); otherwise, the analytical value of the transition matrices is limited. The connectivity matrix describes the probability that an individual moves during its pelagic larval stage from the birthplace (or source population) to its settlement location (or sink population) as a settling larva, all in a threedimensional dynamic system. Such transition probability matrices are of considerable value for metapopulation and genetic studies (Hedrick, 2000) as well as for spatial management and conservation issues (Urban and Keitt, 2001). We show that they also provide a method to quantify the relative influence of biological and physical factors on realized larval dispersal and on levels and spatial patterns of recruits (Paris et al., 2007). The likelihood of larval exchange from one population to another is represented in a transition probability matrix (Paris et al., 2007). The content of a given matrix element describes the probability of an individual larva, making the transition from its source population and successfully reaching the settlement stage in the destination population. Elements along the diagonal of the matrix (where source = sink) represent self-recruitment within a population. Connections between populations may be represented by several types of matrices (Figure 3): (1) the distance matrix dij represents the distances between reefs i and j; (2) the transition probability matrix Pij represents the probability that an individual larva in node i at time t will disperse to node j at time t þ k, where k is the pelagic larval duration; (3) the adjacency matrix (or edge) A = aij is a binary matrix in which each element is defined as aij = 1, if nodes i and j are connected, otherwise aij = 0. This matrix is mostly used to analyze connectivity networks (Urban and Keitt, 2001). The expected flux F from node i to node j is: Fij = Si/Stot  Pij, where Si is the size of the population

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in node i and Stot equals SSi. We set Si to be constant, corresponding with uniform particle release at all locations. Models of population connectivity emphasize where and how frequently larval linkages occur, how these observed patterns are created, and output spatially explicit transition probability matrices. These connectivity matrices describe the probability that an individual moves during its pelagic larval stage from the birthplace (or source population) to its settlement location (or sink population) and are of considerable value for metapopulation and genetic studies (Hedrick, 2000), as well as for spatial management and conservation issues (Urban and Keitt, 2001). Lagrangian stochastic models (LSMs) are being increasingly applied to track the dispersal of larvae, in which motions at small scales that are not resolved by ocean general circulation models (OGCMs) are usually parameterized (Paris et al., 2005). These include motions due to small-scale currents and random or oriented motions of individual, simulated larvae (Paris et al., 2002; Siegel et al., 2003; Codling et al., 2004).

Summary Coral reefs are naturally fragmented and the complex interactions of the reef structure and their larval behavior with the circulation at a series of spatial scales (e.g., lagoon, fringing reef, and global circulation) determine their degree of interconnectivity. Typical connectivity in marine populations ranges broadly from demographic to genetic time scales depending on the strength and frequency of larval migrations (Cowen et al., 2000, 2006). However, the spectrum of connectivity among corals is still uncertain given their longevity and complex reproductive strategies. Indeed, on one hand, a few migrants could suffice to cause demographic connectivity. Conversely, migration could play only a small role on the genetic structure of corals that are highly clonal (i.e., rely mostly on asexual propagation). An integration of numerical and empirical approaches across space and timescales offers the greatest potential for advances in understanding reef interconnectivity, requiring more interdisciplinary interaction (Levin, 2006; Werner et al., 2008). Mapping of reef interconnectivity and identifying important coral larval pathways and corridors are keys for their conservation (Treml et al., 2008). Indeed, larval connectivity helps determine the spacing and size of marine protected areas (Pelc et al., 2010), prioritize the protection of critical stepping stone reefs and nodes of populations’ networks, and maintain key linkages, enhancing reef resilience to climate change–induced stress (Mumby et al., accepted). Bibliography Almany, G. R., Berumen, M. L., Thorrold, S. R., Planes, S., and Jones, G. P., 2007. Local replenishment of coral reef fish populations in a marine reserve. Science, 316, 742–744. Atema, J., Kingsford, M. J., and Gerlach, G., 2002. Larval reef fish could use odour for detection, retention and orientation to reefs. Marine Ecology Progress Series. 241: 151–160.

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Baums, I., Paris, C. B., and Cherubin, L. M., 2006. A biooceanographic filter to larval dispersal in a reef-building coral. Limnology and Oceanography, 51(5), 1969–1981. Becker, B. J., Levin, L. A., Fodrie, F. J., and McMillan, P. A., 2007. Complex larval connectivity patterns among marine invertebrate populations. Proceedings of the National Academy of Sciences, 104(9), 3267–3272. Bonhomme, F., and Planes, S., 2000. Some evolutionary arguments about what maintains the pelagic interval in reef fishes. Environmental Biology of Fishes, 59, 365–383. Botsford, L. W., White, J. W., CoVroth, M.-A., Paris, C. B., Planes, S., Shearer, T. L., Thorrold, S. R., and Jones, G. P., 2009. Measuring connectivity and estimating resilience of coral reef metapopulations in MPAs: matching empirical efforts to modelling needs. Theme Section: Larval connectivity, resilience and the future of coral reefs. Coral Reefs, doi: 10.1007/s00338009-0466-z. Butler, M. J. IV, Paris, C. B., Goldstein, J. S., Matsuda H., and Cowen, R. K. Review. Behavior constrains the dispersal of long-lived spiny lobster larvae. Limnology and Oceanography. Clobert, J., Danchin, E., Dhondt, A. A., and Nichols, J. D., 2001. Dispersal. Oxford: Oxford University Press. Cowen, R. K., Gawarkiewicz, G., Pineda, J., Thorrold, S. R., and Werner, F. E., 2007. Population connectivity in marine systems: an overview. Oceanography, 20, 14–21. Cowen, R. K., Lwiza, K. M. M., Sponaugle, S., Paris, C. B., and Olson, D. B., 2000. Connectivity of marine populations: open or closed ? Science, 287, 857–860. Cowen, R. K., Paris, C. B., and Srinivan, A., 2006. Scaling of connectivity in marine populations. Science, 311, 522–527. Gaines, S., and Roughgarden, J., 1985. Larval settlement rate: A leading determinant of structure in an ecological community of the marine intertidal zone. Proceedings of the National Academy of Sciences USA, 82, 3707–3711. Jones, G. P., Milicich, M. J., Emslie, M. J., and Lunow, C., 1999. Self-recruitment in a coral reef fish population. Nature, 402, 802–804. Jones, G. P., Planes, S., and Thorrold, S. R., 2005. Coral reef fish larvae settle close to home. Current Biology, 15, 1314–1318. Hughes, T. P., Baird, A. H., Bellwood, D. R., Card, M., Connolly, S. R., Folke, C., Grosberg, R., Hoegh-Guldberg, O., Jackson, B.C., Kleypas, J., Lough, J. M., Marshall, P., Nystro, M., Palumbi, S. R., Pandolfi, J. M., Rosen, B., and Roughgarden, J., 2003. climate change, human impacts, and the resilience of coral reefs. Science, 301, 929–933. Kool, J. T., Paris, C. B., Andrefouet, S., and Cowen, R. K., 2009. Complex migration and the development of genetic structure in subdivided populations: an example from Caribbean coral reef ecosystems. Ecography, 32, 1–10. Levin, L. A., 2006. Recent progress in understanding larval dispersal: new directions and digressions. Integrative and Comparative Biology, 46, 282–297. Llopiz, J. K., and Cowen, R. K., 2009. Variability in the trophic role of coral reef fish larvae in the oceanic plankton. Marine Ecology Progress Series, 381, 259–272. McKergow, L. A., Prosser, I. P., Hughes, A. O., and Brodie, J., 2005. Regional scale nutrient modeling: exports to the Great Barrier Reef world heritage area. Marine Pollution Bulletin, 51, 186–199. Manel, S., Gaggiotti, O. E., and Waple, R. S., 2005. Assignment methods: matching biological questions with appropriate techniques. Trends in Ecology and Evolution, 20(3), 136–142. Mumby, P., Elliott, I., Eakin, C., Skirving, W., Paris C. B., Edwards, H., Enriquez, S., Iglesias-Prieto, R., Cherubin, L. M., and Stevens, J., Accepted. Reserve design for uncertain responses of coral reefs to climate change. PloS Biology.

Munday, P. L., Leis, J. M., Lough, J. M., Paris, C. B., Kingsford, M. J., Berumen, M. L., and Lambrechts, J., 2009. Climate change and coral reef connectivity. Theme Section: Larval connectivity, resilience and the future of coral reefs. Coral Reefs, 28(2), 379–395. Nathan, R., 2006. Long distance dispersal in plants. Science, 313(5788), 786–788. Paris, C. B., and Chérubin, L. M., 2008. River-reef connectivity in the Meso-American region. Coral Reefs, 27, 773–781. Paris, C. B., and Cowen, R. K., 2004. Direct evidence of a biophysical retention mechanism for coral reef fish larvae. Limnology and Oceanography, 49, 1964–1979. Paris, C. B., Cowen, R. K., Lwiza, K. M. M., Wang, D. P., and Olson, D. B., 2002. Multivariate objective analysis of the coastal circulation of Barbados, West Indies: implication for larval transport. Deep Sea Research Part I: Oceanographic Research Papers, 49, 1363–1386. Paris, C. B., Cowen, R. K., Claro, R., and Lindeman, K. C., 2005. Larval transport pathways from Cuban spawning aggregations (Snappers; Lutjanidae) based on biophysical modeling. Marine Ecology Progress Series, 296, 93–106. Paris, C. B., Cherubin, L. M., and Cowen, R. K., 2007. Surfing, spinning, or diving from reef to reef: effects on population connectivity. Marine Ecology Progress Series, 347, 285–300. Pelc, R. A., Warner, R. R., Gaines, S. D., and Paris, C. B., 2010. Detecting larval export form marine reserves. PNAS, doi: 0.1073/pnas.0907368107. Pineda, J., Hare, J. A., and Sponaugle, S., 2007. Larval transport and dispersal in the ocean and consequences for population connectivity. Oceanography, 20(3), 22–39. Planes, S., Jones, G. P., and Thorrold, S. A., 2009. Larval dispersal connects fish populations in a network of marine protected areas. PNAS, 106(14), 5693–5697. Sadovy de Mitcheson, Y., Cornish, A., Domeier, M., Colin, P., Russell, M., and Lindeman, K., 2008. Fish Spawning Aggregations; a global baseline. Conservation Biology, 22(5),1233–1244. Schwarz, J. A., Weis, Æ. V. M., and Potts, D. C., 2002. Feeding behavior and acquisition of zooxanthellae by planula larvae of the sea anemone Anthopleura elegantissima. Marine Biology, 140, 471–478. Shanks, A. L., 2009. Pelagic larval duration and dispersal distance revisited. Biological Bulletin, 216, 373–385. Siegel, D. A., Kinlan, B. P., Gaylord, B., and Gaines, S. D., 2003. Lagrangian descriptions of marine larval dispersion. Marine Ecology–Progress Series, 260, 83–96. Simpson, S. D., Meekan, M., Montgomery, J., McCauley, R., and Jeffs, A., 2005. Homeward sound. Science, 308, 221. Steneck, R. S., Paris, C. B., Arnold, S. N., Butler, M. J., Ablan Lagman, M. C., Alcala, A. C., McCook, L. J., Russ, G. R., and Sale, P. F., 2009. Thinking and managing outside the box: coalescing connectivity networks to build region-wide resilience in coral reef ecosystems. Theme Section: Larval connectivity, resilience and the future of coral reefs. Coral Reefs, doi: 10.1007/ s00338-009-0470-3. Swearer, S. E., Shima, J. S., Hellberg, M. E., Thorrold, S. R., Jones, G. P., Robertson, D. R., Morgan, S. G., Selkoe, K. A., Ruiz, G. M., and Warner, R. R., 2002. Evidence of self-recruitment in demersal marine populations. Bulletin of Marine Science, 70, 251–271. Thorrold, S. R., Latkoczy, C., Swart, P. K., and Jones, C. M., 2001. Natal homing in a marine fish metapopulation. Science, 291, 297–299. Thorson, G., 1946. Reproduction and larval development of Danish marine bottom invertebrates; with special reference to the planktonic larvae in the Sound (Øresund). Meddelelser fra Kommissionen for Danmarks Fiskeri-og Havundersøgelser. Serie Plankton, 4(1), 1–523.

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Treml, E. A., Halpin, P. N., Urban, D. L., and Pratson, L. F., 2008. Modeling population connectivity by ocean currents, a graphtheoretic approach for marine conservation. Landscape Ecology, 23, 19–36. Urban, D., and Keitt, T., 2001. Landscape connectivity: a graphtheoretic perspective. Ecology, 82(5), 1205–1218. Werner, F. E., Cowen, R. K., and Paris, C. B., 2007. Coupled biophysical models: present capabilities and necessary developments for future studies of population connectivity. Oceanography, 20(3), 54–69.

of reef-restoration practices in other areas of the world is beyond the scope of this paper. For a global perspective on these and related issues, the following reports are recommended: Hatcher (1984), Salvat (1987), Guzman (1991), Clark and Edwards (1995), Harriott and Fisk (1995), Rinkevich (1995), Heeger and Soto (2000), Jaap (2000), Riegl (2001), Clark (2002), Omori and Fujiwara (2004), Challenger (2006), Kaufman (2006), Jokiel et al. (2006), Precht (2006), and Edwards and Gomez (2007).

Cross-references

Legal mandate for reef restoration in the FKNMS Section 312 of the National Marine Sanctuaries Act is a liability provision that authorizes the National Oceanic and Atmospheric Administration (NOAA) to seek damages from those responsible for injuring sanctuary resources (Davidson, 2006). The Act further mandates that NOAA “restore, replace or acquire the equivalent” of injured resources. Thus, coral reef restoration in the NOAA National Marine Sanctuaries is not a management option, but a legal requirement. Decision making for restoring injured reefs in the FKNMS is guided by findings and recommendations of the damage assessment and restoration program (DARP) and is performed by sanctuary personnel following injury assessment of the grounding site.

Adaptation Climate Change and Coral Reefs Conservation and Marine Protection Areas Coral Reef, Definition Fringing Reef Circulation Global Ocean Circulation and Coral Reefs Hydrodynamics of Coral Reef Systems Lagoon Circulation Reef Structure

REEF RESTORATION J. Harold Hudson1, William B. Goodwin2 1 Reef Tech Inc., Miami, Fl, USA 2 Florida Keys National Marine Sanctuary, Key Largo, Fl, USA

Synonyms Addressing vessel-grounding impacts; Coral habitat restoration; Coral resource rehabilitation Definitions Resource restoration. Resource restoration consists of an attempt to overcome, through manipulation, the factors that impede the natural recovery of an impaired resource. Contship Houston. Contship is an abbreviation for container ship. Coral reef restoration. Reef restoration is the process of returning damaged coral reefs to a state that is functionally equivalent to their uninjured counterparts. Emergency restoration. Emergency restoration involves salvage and reattachment of at-risk coral colonies immediately following a vessel-grounding incident. Introduction Reef restoration is a relatively new and rapidly expanding discipline that attempts to return degraded coral reefs to some degree of their original physical, ecological, economic, and aesthetic functionality on a much-reduced time scale, with the ultimate goal being the full, natural recovery of the impaired resource. The focus of this report is a brief summary of ship-grounding restoration practices in the Florida Keys National Marine Sanctuary (FKNMS), with emphasis on major restoration projects and lessons learned from their implementation. To include examples

Vessel-grounding injuries to coral reefs in the FKNMS The size and nature of injuries to coral reefs from vessel groundings are influenced mainly by three factors: hull length, weight, and vessel construction material. A general rule of thumb is: the larger the vessel, the greater the injury. For purposes of assigning levels of injury, vessel-grounding incidents in this report are grouped by hull length into three categories: small (30 m). It is reasonable to assume that, when a 122 m (400 ft) steel-hulled cargo ship goes aground on a shallow coral reef, there will be massive destruction of reef resources. Such was the case of the M/V Wellwood that grounded on Molasses Reef, off Key Largo in the Florida Keys in August 1984 (Figure 1). Massive coral colonies were split in half, overturned, and reduced to piles of rubble (Figure 2). The area beneath the ship hull was completely scarified and flattened (Figure 3), leaving the calcium carbonatebased reef foundation (reef framework) fractured and devoid of any living corals (Hudson and Diaz, 1988). If a large vessel remains aground for a period of days or weeks before its removal, waves and high seas can cause its hull to act like a pestle in a mortar, grinding away additional reef-framework material (Hudson and Franklin, 2005a). Prolonged shading of uninjured portions of the reef by the hull, can, after several weeks, cause shaded corals to bleach by expelling their zooxanthellae (Gittings et al., 1993). The hull of a large vessel can act as a giant plow, pushing up fragmented and intact coral colonies, coral rubble, reef framework, and sand into massive rubble berms on the periphery of the grounding site

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Reef Restoration, Figure 1 The M/V Wellwood, hard aground on Molasses Reef in the (then) Key Largo National Marine Sanctuary, August 1984. (Photo courtesy of FKNMS.)

Reef Restoration, Figure 2 Living Montastrea faveolata dislodged and split in half by impact of M/V Wellwood hull. (Photo courtesy of FKNMS.)

(Schmahl et al., 2006). A large, deep-draft vessel will often strike and shear off corals and sponges and topographic highs of the reef before finally becoming hard aground. In some instances, this inbound path of destruction may stretch for hundreds of meters and may impact several thousand square meters of coral substrate (Schmahl et al., 2006). The kinds of injuries inflicted upon a coral reef when a medium-size (10–30 m in length)

vessel goes aground, while similar to many of those seen in large-vessel incidents, usually differ by an order of magnitude in size and severity. Total reef injury sustained from intermediate-size vessel groundings rarely exceeds the 100 m2 threshold. Vessels belonging to the smallest-size category (6,000 years it is important to remember that mature coral cays were almost certainly nonexistent during the Holocene transgression, significant for the birds and turtles which now use them as nesting sites. Vegetated cays also serve as roosting sites on bird migration routes and as sources of nutrients leached from guano and from rich organic soils, thereby enhancing algal growth on the reef flat (Chen and Krol, 1997). Vegetated coral cays are an integral part of reef-flat ecology but have a dubious future in the face of rising sea levels and modified sediment budgets produced by global climatic change.

Bibliography Aston, J. P. 1995. The Relative Mobilities of Coral Cays on the Great Barrier Reef Can Be Modelled. Unpubl. M.Sc. Thesis, Sir George Fisher Centre, James Cook University. p. 67. Chen, D., and Krol, A. 1997. Hydrology of Heron Island, Great Barrier Reef. In Vacher, H. L., and Quinn, T. M. (eds.), Geology and Hydrogeology of Carbonate Islands Developments in Sedimentology. Amsterdam: Elsevier, Vol. 54, pp. 867–884.

Flood, P. G. 1986. Sensitivity of coral cays to climatic variations, southern Great Barrier Reef, Australia. Coral Reefs, 5, 13–18. Hopley, D. 1997. Geology of reef islands of the Great Barrier Reef, Australia. In Vacher, H. L., and Quinn, T. M. (eds.), Geology and Hydrogeology of Carbonate Islands, Developments in Sedimentology. Amsterdam: Elsevier, Vol. 54, pp. 835–866. Hopley, D., Smithers, S. G., and Parnell, K. E. 2007. The Geomorphology of the Great Barrier Reef: Development, Diversity and Change. Cambridge: Cambridge University Press. p. 532.

Cross-references Atoll Islands (Motu) Beach Rock Cay Formation Coral Cay Classification and Evolution Coral Cays-Geohydrology Coral Cays, Vegetational Succession Infrastructure and Reef Islands Phosphatic Cay Sandstone Soils of Low Elevation Coral Structures Unvegetated Cays

VOLCANIC DISTURBANCES AND CORAL REEFS Peter Houk Pacific Marine Resources Institute, Inc., Saipan, MP, USA

Definition Volcanic eruptions refer to the release of volcanic ash and magma to the atmosphere and surrounding waters containing coral reefs. Introduction Volcanoes have been emphasized in coral reef studies since Charles Darwin first formalized his thoughts surrounding the creation of tropical islands and atolls (Figure 1, Darwin, 1842). Volcanic activity is arguably the most influential natural disturbance that impacts reef growth and modern coral assemblages. However, its frequency of occurrence is much lower than other natural disturbances such as climate-induced bleaching and tropical cyclones. Physically, eruptions deliver hot magma that forms igneous rock of varying composition (Le Bas and Streckeisen, 1991), as well as volcanic ash, to the nearshore waters. Both can smother vast expanses of living coral reefs rapidly (Pandolfi et al., 2006), but longer-lasting, secondary impacts also exist. Volcanic ash particles contain numerous minerals in high concentrations that influence surrounding waters; most notably iron, magnesium, and silica (Flaathen and Gislason, 2007). Minerals from volcanic ash provide a pulse of key limiting ingredients that facilitate the rapid growth of phytoplankton, translating to enriched productivity of surface waters, often at the scale of entire oceanographic regions (Uematsu et al., 2004). These plankton blooms eventually die and decay, releasing nutrients and bacteria as they sink. Corals cannot uptake all of these available nutrients as fast as macroalgae, so following major volcanic ash input,

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Volcanic Disturbances and Coral Reefs, Figure 1 Anatahan, Commonwealth of the Northern Mariana Islands, following a 2003 eruption that provided no lava flow but lots of volcanic ash to surrounding waters.

patches of macroalgae and cyanobacteria growth often emerge on the surrounding coral reefs. Even when the volcanic activity has ceased, re-suspension of volcanic ash during large-wave events can continue to facilitate this cycle, as well as increase turbidity levels and reduce the penetration of sunlight for substantial periods of time.

Reef growth It has long been recognized that igneous rock created from magma differs in porosity and permeability (Davis, 1969), and thus will provide for differential connectivity with island aquifer systems. Houk and Starmer (2010) show that the remote volcanic Mariana Islands, Western Pacific Ocean, have strikingly different (up to 1 part per thousand) ambient salinity levels, attributed to differential igneous rock porosity and permeability. In turn, this creates selective environmental conditions favorable for algae and sponge growth where low salinity exists, and increased coral growth where salinity is higher (Figure 2). The former yield low-rugosity reef structures with little three-dimensionality through time, while the latter provides for high-relief primary coral framework development. Clearly, volcanic bedrock is not all the same and differing characteristics have consequences for reef growth over time. Modern assemblages Due to its rarity and unpredictable timing, relatively few studies have examined the impacts of volcanic activity upon reefs and subsequent recovery. Studies have shown that while lava flows are detrimental to living reef communities, regrowth of vibrant coral assemblages can occur within 5 years in Indonesia (Tomascik et al., 1996). In Hawaii, where wave exposure and seasonal temperature variation are much higher, studies suggest a much longer time period may be required for modern coral growth to return, up to 50 years (Grigg and Maragos, 1974). A 2003 eruption of Anatahan, Mariana Islands, yielded no lava flow but large quantities of volcanic ash that smothered most of the surrounding reef (Figure 1, Houk and Starmer, 2010).

Volcanic Disturbances and Coral Reefs, Figure 2 Growth of modern assemblages on coral reefs surrounding volcanic islands strongly differs depending upon the characteristics of igneous rock. High-porosity bedrock has high connectivity with the islands aquifer and provides for continuous algae growth over time and low three-dimensionality of reef structure (a), compared with low-porosity systems (b).

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Here, environmental conditions are similar to Hawaii and wind-generated waves often re-suspend volcanic ash, thus selecting against rapid coral settlement and growth. Predictions using other islands that have been erupting at varying times throughout the past 100 years revealed that 90 years is estimated for coral abundance and species richness recovery to expected levels at the entire island scale.

Bibliography Davis, S. N., 1969. Porosity and permeability of natural materials. In Wiest, R. J. M. D. (ed.), Flow Through Porous Media, New York: Academic Press, pp. 53–89. Darwin, C. R., 1842. The Structure and Distribution of Coral Reefs, London: Smith, Elder and Company. Flaathen, T. K., and Gislason, S. R., 2007. The effect of volcanic eruptions on the chemistry of surface waters: the 1991 and 2000 eruptions of Mt. Hekla, Iceland. Journal of Volcanology and Geothermal Research, 164, 293–316. Grigg, R. W., and Maragos, R. E., 1974. Recolonization of hermatypic corals on submerged Hawaiin lava flows. Ecology, 55, 387–395. Houk, P., and Starmer, J., 2010. Constraints on the diversity and distribution of coral-reef assemblages in the volcanic Northern Mariana Islands. Coral Reefs, 29, 59–70, doi:10.1007/s00338009-0545-1. Le Bas, M. J., and Streckeisen, A. L., 1991. The IUGS systematics of igneous rocks. Journal of the Geological Society, 148, 825–833. Pandolfi, J. M., Tudhope, A. W., Burr, G., Chappell, J., Edinger, E., Frey, M., Steneck, R., Sharma, C., Yeates, A., and Jennions, M., 2006. Mass mortality following disturbance in Holocene coral reefs from Papua New Guinea. Geological Society of America, 34, 949–952. Tomascik, T., van Woesik, R., and Mah, A. J., 1996. Rapid coral colonization of a recent lava flow following a volcanic eruption, Banda Islands, Indonesia. Coral Reefs, 15, 169–175. Uematsu, M., Toratani, M., Kajino, M., Narita, Y., Senga, Y., and Kimoto, T., 2004. Enhancement of primary productivity in the western North Pacific caused by the eruption of the Miyake-jima Volcano. Geophysical Research Letters, 31, L06106, doi:10.1029/2003GL018790.

Cross-references Antecedent Platforms Corals: Environmental Controls on Growth Darwin, Charles (1809–1882) Oceanic Hotspots Plate Tectonics

VOLCANIC LOADING AND ISOSTASY Kurt Lambeck Australian National University, Canberra, ACT, Australia The Earth’s mechanical response to changes in surface loading is usually described in terms of local or regional isostasy in which the load is supported to varying degrees by the strength of the crust or lithosphere and by the reaction of the underlying mantle to the deformation of this layer.

For a general formulation for this response see Chapter Glacio-Hydro Isostasy and the response to volcanic loading can be seen as a special case of this. The essential differences between the two cases are (1) the timescale of loading is usually much longer than that of the glacial cycles, (2) the length scale of the loads are shorter than that of the large ice sheets, (3) the load-stress magnitude of the volcanic load may be much larger than that of ice loads and (4) the volumes of the volcanic loads are comparatively small such that when the volcanoes form in an ocean environment the water displaced and the concomitant hydro-isostasy can usually be ignored. The formation of large volcanic complexes usually occurs over periods of a million years or longer. This is greater than the typical relaxation time of the mantle inferred from the glacio-hydro isostatic analyses such that the mantle can usually be considered as a fluid of zero viscosity. However, the loading time is likely to be comparable to the relaxation times of the lithosphere, particularly for the lower lithosphere, such that the viscosity structure of the lithosphere may be important. The smaller length scale of the volcanic loads and the unimportance of the hydro-isostatic loading means that the isostatic models need not be global and flat-earth representations are usually adequate. But, combined with the load-stress magnitudes being generally larger, it also means that the structure of the crust and the lithospheric part of the mantle may also be important in determining the isostatic response. The larger load stresses also mean that elastic failure of the crust may occur. Another difference between the two loading problems is that the process of volcanic loading itself may have a modifying effect on lithospheric properties because of heat transport into the lithosphere during the active phase of loading. The isostatic response to volcanic (and sediment) loads can be qualitatively represented by the regional isostatic model in which the stress-bearing layer overlies a fluid mantle of zero viscosity. Usually, this layer will have depth-dependent physical properties but for computational convenience any lateral variability is usually ignored. The time history of loading is also mostly ignored. Figure 1 illustrates the schematic elastic response for a disk-shaped load for which there exist convenient analytical solutions (such solutions also exist for other axially symmetries such as parabolic profile loads). The deformation beneath the load is characterized by a “flexural parameter” or “radius of relative stiffness,” l, defined as l4 ¼ D=ðrm
rl Þg D is the flexural rigidity and for a depth (z) dependence of elastic parameters (shear modulus m and bulk modulus Κ with l = Κ
(2m/3) where the integration is across the elastic layer from its upper surface z = þH/2 to its lower surface z =
H/2. For typical elastic parameter values, l  50 km and much of the spatial variation in the deformation occurs over a distance of about 4l.

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A = 60 km A = 120 A = 180 A = 240 A = 300 A = 360 A = 420 r (km)

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Volcanic Loading and Isostasy, Figure 1 Deflection of an elastic plate of flexural rigidity D overlying a zero viscosity fluid, subject to a disk load of variable radius A, density 2.8 gm/cm3 and height 5 km. The moat forming at the edge of the load is assumed to be filled in with material of the same density up to the level of the undeformed surface consistent with the definition of D used here. From Lambeck and Nakiboglu (1980).

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Height (km)

Uplift (m)

Once the load radius exceeds about 3–4l, the displacement beneath the center of the load approaches that of local isostatic compensation (see Chapter GlacioHydro Isostasy) with a maximum displacement of zmax  –(rl/rm)h where h and rl are the height and density of the load (h is measured with respect to the undeformed surface). Solutions (1, 2) in Figure 1, with a load radius of up to about 100 km, are representative of single volcanic complexes, whereas the larger diameter loads are more representative of basaltic plateau or large sedimentary loads. The solutions for large radius begin to resemble the surface deflection across continental margins from melt-water loading in the glacio-hydro isostatic problem. Normally, a broad trough develops around the load down to a maximum depth of z  0.5zmax that can be expected to be partly filled as part of the volcanic complex and partly by water if the formation is in an ocean environment. This is seen around some of the large volcanic complexes such as the Hawaiian chain. Thus, in more realistic models the question of how this moat is filled must be considered. Beyond the moat, a small uplifted bulge occurs on the order of some tens of meters in peak amplitude and typically 100–150 km in width (Figure 2), with the actual dimensions being a function of the lithospheric parameters, load dimensions and the degree to which the moat has been filled. If a viscous component is introduced to describe stress relaxation in the lithosphere, particularly in the lower part of this layer, then the sea floor subsides until the local isostatic limit is approached and the outer bulge migrates inwards. Observational evidence that supports these models of regional compensation or lithospheric flexure comes from gravity (including geoid), seismology and sea-level

40 20 200

250 300 r (km)

4

60 r (km) 350

400

Volcanic Loading and Isostasy, Figure 2 Uplift of the peripheral bulge for the disk load of 60 km radius. From Lambeck (1981).

investigations. The seismological evidence includes mapping of the slope of seismic reflectors within the crust beneath the load and moat and variations in the depth of the Moho. The seismic data can also define some of the internal crustal structure beneath the volcano and establish whether faulting has occurred as a result of loading. Gravimetric and geoid surveys indicate positive anomalies over seamounts and volcanoes that are consistent with the mass distribution within the regional compensation models and permit the flexural parameters to be estimated. Primary outcome of such studies include estimates of the effective elastic thickness of the lithosphere, the dependence of D on the age of the lithosphere at the time of volcanic seamount formation, and evidence for relaxation of the load stress for the larger volcanic

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VOLCANIC LOADING AND ISOSTASY

loads. These estimates of the effective elastic thickness are generally significantly less than the estimates from hydro-isostatic analyses, indicating that stress relaxation occurs in the lower lithosphere on timescales between the glacial cycles and volcanic loading and longer. The sealevel data point to the uplift of atolls and other islands positioned on the peripheral bulge. This appears, for example, to have been the case for Henderson Island that was uplifted as a result of the formation of the Pitcairn volcanic complex at 0.8 million years ago. Another example is the island of Makatea which was probably uplifted during the formation of the Tahiti volcanic complex.

Bibliography Watts, A. B., 2001. Isostasy and Flexure of the Lithosphere. Cambridge: Cambridge University Press, p. 458. Lambeck, K., and Nakiboglu, S.M., 1980. Seamount loading and stress in the ocean lithosphere. Journal of Geophysical Research, 85, 6403–6418. Lambeck, K., 1981. Flexure of the ocean lithosphere from island uplift, bathymetry and geoid height observations: the Society Islands. Geophys Journal of the Royal Astronomical Society, 67, 91–114.

Cross-references Glacio-Hydro Isostasy

W

WACKESTONE Peter Flood University of New England, Armidale, NSW, Australia Wackestone is a type of limestone recognized in the Dunham (1962) Classification Scheme in which grains make up more than 10% of the rock but the grains are mud supported, i.e., the grains float in a mud matrix.

Bibliography Dunham, R. L., 1962. Classification of carbonate rocks according to depositional texture. Memoir American Association Petroleum Geologists, 1, 108–121.

Cross-references Porosity Variability In Limestone Sequences

WALTHER, JOHANNES (1860–1937) Eberhard Gischler Institut fuer Geowissenschaften, Frankfurt am Main, Germany Johannes Walther was one of the founders of sedimentology and paleoecology (Middleton, 1973; Seibold, 1992; Ginsburg et al., 1994), and conducted four fundamental studies on modern coral reefs (Walther, 1885, 1888, 1891, 1910). Walther was born on 20 July 1860 in Thuringia, Germany. He studied geology and biology at the University of Jena, where his mentor was the famous biologist Ernst Haeckel. Walther completed his Ph.D. in 1886. Subsequently, he studied modern algal reefs in the

Mediterranean and fossil reefs in the Tertiary of Sicily and in the Triassic of the Austrian Calcareous Alps (Walther, 1885). He was among the first who recognized the significance of calcareous algae as reef builders. Based on his studies in the modern reefs, and using thin-section and chemical analyses, he attempted to explain the textures he found in the fossil reef limestones. In his work on the coral reefs of the Sinai Peninsula, Walther (1888) was able to study modern and elevated subfossil reefs nearby. The latter offered three-dimensional sections, which led Walther to the conclusions that only 40% of a reef is formed by coral framework and that 60% is made of detritus filling the interstices. Furthermore, he realized the importance of antecedent topography for reef development and distribution. Walther (1891) established his findings when studying modern and subfossil reefs between India and Ceylon. Again, he stressed the importance of detritus in reefs and the fact that a great part of this material is produced by the destructive activity of organisms. Also, he considered changes in relative sea-level, when discussing subfossil reef terraces. In 1890, Walther became a Professor at the University in Jena. From 1906 to 1928, he held the chair of geology and paleontology at the University of Halle. He returned to the algal reefs in the Mediterranean in 1910 and found significant facies changes that had occurred over the past 25 years, geologically a very short time period. In 1914, Walther was able to visit the Great Barrier Reef of Australia, but he had to return home early due to the outbreak of WW I. Unlike other reef researchers during that time period, Walther did not contribute to the coral reef problem, i.e., the interpretation of shallow atolls and barrier reefs in oceanic settings in the light of Darwin’s (1842) subsidence theory. Even so, he was among the first to recognize the importance of calcareous algae for reefbuilding, described what is today called bioerosion, in great detail, hinted to the significance of loose sediment

David Hopley (ed.), Encyclopedia of Modern Coral Reefs, DOI 10.1007/978-90-481-2639-2, # Springer Science+Business Media B.V. 2011

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in modern reefs, and elaborated the meaning of antecedent topography for reef development. Walther died on 4 May 1937 in Hofgastein, Austria.

See also Chapter Waves and Wave-Driven Currents for definitions of other terms concerning waves. Mathematical symbols are defined where they first appear in the text, or on a relevant figure. The subscript o refers to the deep water values of the various quantities.

Bibliography Darwin, C. R., 1842. The Structure and Distribution of Coral Reefs. London: Smith Elder, 214 p. Ginsburg, R. N., Gischler, E., and Schlager, W., 1994. Johannes Wather on reefs: pioneering concepts of Biogeology 1885–1910. Geological Milestones, 2, 141 p. Middleton, G. V., 1973. Johannes Walther’s law of the correlation of facies. Geological Society of America Bulletin, 84, 979–988. Seibold, I., 1992. Der Weg zur Biogeologie: Johannes Walther 1860–1937. Ein Forscherleben im Wandel der deutschen Universität. Berlin: Springer, 196 p. Walther, J., 1885. Die gesteinsbildenden Kalkalgen des Golfes von Neapel und die Entstehung structurloser Kalke. Z. dt. Geol. Ges., 37, 329–357. Walther, J., 1888. Die Korallenriffe der Sinaihalbinsel. Geologische und biologische Beobachtungen. Abh. Königl. Sächs. Ges. Wiss., 24, 439–505. Walther, J., 1891. Die Adamsbrücke und die Korallenriffe der Palkstrasse. Petermanns Mitt. Erg.-Heft, 102, 1–40. Walther, J., 1910. Die Sedimente der Taubenbank im Golfe von Neapel. Abh. Königl. Preuss. Akad. Wiss. 1910/3, pp. 1–40.

Cross-references Antecedent Platforms Binding Organisms Bioerosion Fringing Reefs Darwin, Charles (1809–1882) Holocene Reefs: Thickness and Characteristics Red Sea And Gulf Of Aqaba Rhodoliths

WAVE SET-UP Michael R. Gourlay The University of Queensland, Brisbane, QLD, Australia

Definitions Wave run-up. The vertical distance between the ocean (tide) level and the maximum height reached by the uprush of waves breaking on a beach or structure. Wave set-up and set-down: Positive and negative changes in mean water level produced as waves shoal and break on beaches and reefs. Wave thrust: A vertically integrated force per unit wave crest length produced by changes in wave momentum as waves propagate in shoaling and shallow water. It has two components; an isotropic pressure and a unidirectional force in the direction of wave propagation. Also known as radiation stress. Shoaling. The process by which the wave height, length, celerity and other properties of the waves change as they travel from deep(er) water into shallow(er) water.

Introduction In the late 1940s, it was observed at Bikini Atoll that the predominant 12–15 s swell waves breaking on the reef caused a significant inflow of ocean water into the atoll lagoon. The accompanying increase in water level in the lagoon, known as wave set-up, was 0.45–0.6 m (Munk and Sargent, 1954; Von Arx, 1954). As the significance of waves as a dominating agent in determining reef-top morphology and ecology was increasingly recognized in the late 1970s and the 1980s, it became apparent that wave set-up plays a significant role in defining wave-driven current systems on coral reefs (e.g., Roberts, 1981; Hearn et al., 1986). (see Chapter Waves and Wave-Driven Currents). Wave set-up is also an important factor, particularly in microtidal environments, in determining groundwater levels in beaches on coral cays or other reef-protected shorelines. These wave-induced groundwater levels influence the formation of Beach Rock, as well as the success of turtle nesting behind the beach. Wave set-up and associated wave run-up also have been identified as important factors in causing flooding of low-lying reef-protected coasts and reef islands (see Chapter Infrastructure and Reef Islands). Wave set-up on beaches A wave train propagating onto a beach or reef produces changes in mean pressure within the water and a unidirectional force in the direction of wave propagation (Longuet-Higgins and Stewart, 1964). This wave thrust increases as the water depth decreases. On a beach offshore of the breaker zone, if it is assumed that there is no frictional dissipation at the bottom, the increasing wave thrust is balanced by a change in mean hydrostatic pressure such that the mean water level decreases as the waves propagate shoreward, that is, wave set-down occurs. Within the surf zone where energy is being dissipated by wave breaking, the wave thrust decreases as the breaking surge travels shoreward and consequently the mean water level rises , that is, wave set-up occurs (Figure 1a). Figure 1a shows the occurrence of wave set-down and set-up on a relatively steep beach. The still water level (s.w.l.) is the water level if there were no waves present and is the reference level for measuring set-down and set-up. In an oceanic situation the s.w.l. will vary with the tides and meteorological influences such as storm surge. The changes in mean water level (m.w.l.) occur when waves approach, shoal, break and run up the beach. The m.w.l. can not be seen but it can be measured either as the time-averaged mean position of the water level at any specific location approaching and within the surf zone or

WAVE SET-UP

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Wave Set-Up, Figure 1 Wave set-up on beaches. (a) Waves breaking on a steep, plane beach. (b) Waves breaking on a beach with offshore bar.

as the mean hydrostatic pressure on the bottom at that location. Maximum set-down
b occurs in the vicinity of the break point; the maximum set-up
m occurs where the m.w.l. intersects the beach face. However, measurements of set-up in the swash zone are not always reliable and there are differences of opinion as to their meaning (Gourlay, 1992). Some researchers prefer to refer to the set-up measured at the location where the s.w.l. intersects the beach profile (Dean and Walton, 2009). If it is assumed that the height of the breaking surge is a constant fraction g of the actual water depth d (including set-up
, i.e., d ¼
 þ h) in the surf zone, the theory of Bowen et al. (1968) predicts that the ratio of the m.w.l slope in the surf zone tan b to the beach slope tan a is a simple function of g, that is,
 8 tan b= tan a ¼ K ¼ 1= 1 þ 2 3g Laboratory experiments (Bowen et al., 1968) confirmed the existence of set-down and set-up on plane beaches. Subsequent analysis of a wide range of laboratory data for impermeable plane beaches suggested that the maximum set-up on fairly steep beaches (tan a  0.1) is about 30% of the initial breaker height Hb, whereas for flat beaches (tan a  0.04) it is only 15% of it (Gourlay, 1992). For comparison, wave set-down probably would be no more than 5% of the breaker height.

Nielsen (1988) measured wave set-up on a beach at Dee Why, NSW, which has a relatively steep face and an offshore bar. There were different set-up profiles for small and large waves. Small waves broke directly on the beach face creating a steep m.w.l gradient shoreward of the initial breakpoint. Large waves broke on the offshore bar or flatter portions of the profile. In this case, the m.w.l had two distinct gradients, a flat one in the outer surf zone and a steeper one following a second break point inshore on the beach face (Figure 1b). These experiments also revealed that, even if there were no rainfall, the groundwater level behind the beach is elevated above the ocean tide level. This groundwater level varies with changes both in the height of the waves and in tide levels. Laboratory experiments on wave-formed beaches (Gourlay, 1992) indicated wave set-up profiles for large and small waves similar in general form to those observed by Nielsen. They also showed that, for steep beaches, the wave set-up and hence the wave-induced groundwater level were larger for beaches formed of fine sand in comparison with those formed of medium sand or fine pebbles (see discussion below). Recent analyses of various field data indicate that the set-up at the still water line is about 20% of the significant wave height Hos with a variation of about 10% (Dean and Walton, 2009). Set-up
m should not be confused with run-up Ru. The latter term refers to the maximum vertical height on the beach reached by the swash/uprush from the breaking

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WAVE SET-UP

waves. Run-up is commonly measured relative to s.w.l. on beaches and engineering structures but this may not be appropriate for reef-protected beaches (see discussion below). For a plane beach or structure, the run-up is related to the wave conditions and beach slope by the following equation: Ru pffiffiffiffiffiffi ffi ¼ C tan a T gH where the parameter C (1) varies with the roughness and permeability of the beach (see Nielsen, 2009).

Wave set-up on coral reefs Wave set-up on reefs is governed by the same theoretical principles as wave set-up on beaches. However, there are some differences in approach because the morphology of reefs is different from that of beaches. In general, the profile of a reef will have a reef face of variable slope, a reef crest/ rim, and a reef flat and/or lagoon. In the case of a fringing reef or cay on a platform reef, normally there will be a relatively steep beach at the far side of the reef flat (Figure 2). With a platform reef or atoll lagoon, there will be a leeward reef rim or reef. When the reef face is not too steep (tan a < 0.1) and the waves break on it, the theory of Bowen et al. (1968) for set-up on a plane beach has been applied by Tait (1972)

to give an estimate of wave set-up on a horizontal reef top. As reef faces are generally much rougher than sandy beach profiles, bottom friction may be significant and probably will reduce the wave height and hence the wave thrust sufficiently to eliminate any wave set-down on the reef face prior to breaking. With this assumption, the wave set-up
r on a horizontal reef is given by
r ¼ Kðhb  hr Þ (Figure 2a). The set-up is maximum when hr = 0, i.e., when the tide level is equal to the reef top level, and it reduces to zero when hr  hb, i.e., when waves pass over the reef without breaking. During the period between 1980 and 1996, several researchers measured wave set-up on laboratory models of coral reefs with widely varying profiles (Gourlay, 1996a,b). Reef-face slopes varied from greater than 1 in 50 to vertical. One model reproduced a platform reef with flow across it, whereas all the others were fringing reefs with horizontal or very flat reef tops; one of these latter had a shallow lagoon between the reef crest and the beach behind the reef. Two models involved experiments with irregular waves as well as regular ones. In only one case was any attempt made to measure the wave set-up on the actual reef in the field. In general, these laboratory experiments indicate that wave set-up on a quasi horizontal reef increases with both increasing ocean wave height and wave period and decreasing water depth over the reef top. No wave set-up

Wave Set-Up, Figure 2 Wave set-up on coral reefs. (a) Waves breaking on reef face. (b) Waves breaking at reef edge.

WAVE SET-UP

occurs if the water depth is large enough for the waves to pass over the reef without breaking. This situation occurs when the deep water ocean wave height Ho is less than 0.4 hr. For steep-faced reefs, waves which just pass over the reef edge may break on the reef top, whereas larger waves will break at the reef edge. For waves breaking on the top of a steep-faced reef (Figure 2b), a theoretical analysis using wave thrust theory predicted that the set-up on a horizontal reef is about 14% of the ocean wave height (Gourlay, 1996a). This is close to the maximum set-up on a flat beach. For flatter reef faces (Figure 2a), larger waves break on the reef face seaward of the reef crest and in some cases dissipate significant energy before the broken waves reach the reef crest. In these latter situations, wave set-up on the reef top approaches a limiting value with increasing wave height. Furthermore, the magnitude of the set-up on the reef top decreases as the slope of the reef face decreases. Analysis of experimental data for both steep and flatter reef faces shows that wave set-up on the reef top increases with increasing Ho2T, which quantity is proportional to the ocean wave power or energy flux (Gourlay, 1996b) (see Chapter Waves and Wave-Driven Currents). Munk and Sargent (1954) first suggested a possible relationship between wave set-up on coral reefs and ocean wave power. Gourlay (1996b) applied this principle to waves breaking at the edge of a steep-faced horizontal reef. This analysis produced a relationship for wave setup on a steep-faced horizontal reef which reduces to the following form: "
2 # 0:015Kp g1=2 Ho2 T dr 1  0:16
r ¼ 3=2 H o dr where the reef profile factor Kp (1) varies with the reef face slope tan a and dr is the water depth, including wave set-up, on the reef top. Ocean waves are variable in height and period and are commonly represented by the significant wave height Hos. In calculating wave set-up on a reef, Ho in the above formula is the height of a regular sinusoidal wave, which is equivalent to the root-mean-square wave height Horms ð¼ Hposffiffi2Þ of irregular ocean waves (see Chapter Waves and Wave-Driven Currents). Numerical models which reproduce wave transformation and breaking on flatter but still relatively steep reef faces have been developed by other researchers. Both Skotner and Apelt (1999) and Massel and Gourlay (2000) have computed wave set-up values which agree reasonably well with those measured on laboratory models with different reef face slopes and profiles. Irregular ocean waves tend to occur in groups of larger waves followed by groups of smaller waves. These groups of larger waves pump water onto the reef top or lagoon, increasing the set-up above its mean value but the following groups of smaller waves are unable to sustain this higher water level and water flows back seawards over the reef rim reducing the set-up below the mean value.

1147

Hence, the wave groups create a dynamic component of wave set-up which is superimposed upon the steady wave set-up calculated using Horms. The magnitude of the dynamic set-up or surf beat depends upon the width of the reef. The wider the reef the greater its capacity to absorb and smooth out the fluctuations in set-up resulting from the wave groups. On the other hand, the dynamic set-up fluctuations decrease in magnitude as the water depth over the reef crest decreases, even though the magnitude of the steady wave set-up increases with decreasing water depth. Hence, surf beat will be more significant at high tide than low tide and at high spring tides than neap tides (Seelig, 1983; Gourlay, 1996b). In certain situations, resonant surf beat oscillations may be instigated by the coincidence of the wave group period with the natural period of oscillation of the water body on the reef flat. Substantial amplification of the surf beat at the shoreline and consequent severe erosion or damage to shore facilities is possible (Nakasa and Hino, 1990). Recent laboratory and numerical modeling studies on the same reef profile as used in Seelig’s experiments show clearly the influence of surf beat (infragravity waves) in producing resonant set-up oscillations on a narrow coral reef. Such oscillations dominate the run-up on a reef-protected beach (Nwogu and Demirbilek, 2009). Field observations following Typhoon Rus at Guam in 1991 found an inverse relationship between reef flat width and overwash height (run-up). That is, the highest overwash occurred along steep coasts adjacent to narrow reef flats where the contribution of waves to overwash was greatest (Jaffe and Richmond, 1993). Various theoretical and empirical formulae representing the processes just described have been utilized for calculating estimates of the wave set-up on a coral reef and the conditions under which a reef-top island is likely to be flooded during an extreme event (Gourlay, 1996b; 1997). After an allowance has been made for storm surge, these calculations initially determine the set-up resulting from waves breaking on the reef rim and then consider the influence of the surf beat upon water levels in the lagoon. Calculation of the wave run-up on the beach, relative to the m. w. l. on the reef top, gives another estimate of the waves’ flooding potential for a specific situation. Such calculations clearly show both the protective action of coral reefs in reducing the size of waves breaking on reef island beaches and the potential for wave set-up to raise reef-top water levels significantly higher than high tide level and hence flood reef islands during extreme events. To determine the wave-induced groundwater level
g behind the beach (Figure 3), it is necessary to calculate the wave set-up caused by the residual waves breaking on the relatively steep plane reef island beach for the appropriate ocean tide plus set-up water level on the reef top in front of the beach. At the time of writing, there is no simple way for calculating the wave set-up in a permeable beach of a particular sand size, let alone for a permeable beach comprised of a wide range of materials, such as is frequently found on reef island beaches. However, advanced numerical modelling simulating the

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Wave Set-Up, Figure 3 Wave set-up, wave run-up and wave-induced groundwater for a reef top beach.

interaction between waves and beach groundwater has produced encouraging results when compared with experimental laboratory data for a beach of a homogeneous sand (Ang et al., 2004). The ground water level behind a reef island beach will also be affected by changing tide levels on the reef flat, as well as rainfall on the island.

Measurements of wave set-up on coral reefs There have been very few actual measurements of wave set-up on coral reefs. Apart from situations where set-up caused by swells from distant storms or hurricanes has resulted in flooding or damage to infrastructure on reef islands (Gourlay, 1996b; Amanaki et al., 2003), the most complete set of field data of wave set-up across a reef flat has been obtained by Jago et al. (2006) at Lady Elliot Island, the southernmost reef in the Great Barrier Reef region. These measurements were made with simple slow response stilling wells and measuring techniques similar to those developed by Nielsen (1999) for measuring wave set-up on beaches. Waves and mean water levels were measured over both rising and falling tides during low to moderate wave conditions (Hos < 0.5 m). Both wind waves and swell were present. The spring and neap tidal ranges at this site are 1.7 m and 0.9 m, respectively. The set-up was observed to vary both spatially and temporally across the reef with changing water depth. At low tide, both wind waves and swell broke on the reef edge generating set-up on the reef flat. At midtide, set-up occurred both at the reef edge, where wind waves broke, and on the island’s beach, where the swell broke. At high tide, no waves broke at the reef edge and there was negligible set-up there but significant set-up occurred on the beach, where both wind waves and swell broke. Setdown also was observed both on the reef face before waves broke on the reef edge and on the reef flat as waves passed over the reef rim into deeper water on the reef flat. Recent measurements of wave set-up and wave-driven currents in an atoll lagoon in the Cook Islands have shown that wave pumping drives the flushing of these lagoons

when there are no deep passages or gaps in their surrounding reef (see Chapter Waves and Wave-Driven Currents).

Summary Waves breaking on a reef or beach create an increase in m.w.l in the surf zone called wave set-up. This set-up increases with increasing height and period of the ocean waves and decreases with increasing water depth over the reef, becoming negligible when waves pass over the reef without breaking. Maximum set-up occurs when the ocean water (tide) level is about the same level as the reef top. The water level differences caused by wave set-up drive reef-top current systems and the flushing on many atoll lagoons. Wave set-up may contribute significantly to the groundwater elevation in cays and the sedimentary margins of continental islands. Wave set-up caused by both local storm waves and swells from distant storms may cause significant flooding, erosion and damage on reef islands, particularly where resonant oscillations occur on a reef top. Bibliography Amanaki, D., Imrie, J., Colleter, G., Foster, M., and Cummings, P., 2003. Wave setup induced damage to the Nafanua Harbour breakwaters, Eua, Kingdom of Tonga. In Coastal Engineering 2002: Proceedings 28th International Conference. Cardiff, Wales: World Scientific, Vol. 2, pp. 1633–1637. Ang, L. S., Sum, C. H. -Y., Baldock, T. E., Li, L., and Nielsen, P., 2004. Measurement and modelling of controlled beach groundwater levels under wave action. In Proceedings 15th Australasian Fluid Mechanics Conference. University of Sydney, CD-ROM. Bowen, A. J., Inman, D. L., and Simmons, V. P., 1968. Wave “setdown” and set-up. Journal Geophysical Research, 73, 2569– 2577. Dean, R. G., and Walton, T. L., 2009. Wave set-up. In Kim, Y. C. (ed.), Handbook of Coastal and Ocean Engineering. Singapore: World Scientific, pp. 1–23. Gourlay, M. R., 1992. Wave set-up and beach water table: Interaction between surf zone hydraulics and groundwater hydraulics. Coastal Engineering, 17, 93–144. Gourlay, M. R., 1996a. Wave set-up on coral reefs. 1. Set-up and wave-generated flow on an idealised two dimensional horizontal reef. Coastal Engineering, 27, 161–193. Gourlay, M. R., 1996b. Wave set-up on coral reefs. 2. Set-up on reefs with various profiles. Coastal Engineering, 28, 17–55.

WAVE SHOALING AND REFRACTION

Gourlay, M. R., 1997. Wave set-up on coral reefs: Some practical applications. In Proceedings 13th Australasian Coastal and Ocean Engineering Conference and 6th Australasian Port and Harbour Conference. University of Canterbury, N. Z., Centre for Advanced Engineering, Vol. 2, pp. 959–964. Hearn, C. J., Hatcher, B. G., Masini, R. J., and Simpson, C. J., 1986. Oceanographic processes on the Ningaloo Coral Reef, Western Australia. University of Western Australia, Centre for Water Research, Report Number: ED-86–171. Jaffe, B. E., and Richmond, B. M., 1993. Overwash variability on the shoreline of Guam during typhoon Rus. In Proceedings Seventh International Coral Reef Symposium. Guam, 1992: University of Guam Press, Vol. 1, pp. 257–264. Jago, O. K., Kench, P. S., and Brander, R. W., 2006. Field observations of wave-driven water-level gradients across a coral reef flat. Journal of Geophysical Research, 112, C06027. Longuet-Higgins, M. S., and Stewart, R. W., 1964. Radiation stress in water waves, a physical discussion with applications. DeepSea Research, 11, 529–562. Massel, S. R., and Gourlay, M. R., 2000. On the modelling of wave breaking and set-up on coral reefs. Coastal Engineering, 39, 1–27. Munk, W. H., and Sargent, M. C., 1954. Adjustment of Bikini Atoll to ocean waves. U.S. Geological Survey Professional Paper, 260-C, 275–280. Nakasa, E., and Hino, M., 1990. Reef-zone disaster caused by borelike surf beat. Coastal Engineering in Japan, 33, 49–61. Nielsen, P., 2009. Coastal and Estuarine Processes. Advanced Series on Ocean Engineering. Singapore: World Scientific, Vol. 29. Nielsen, P., 1999. Simple equipment for coastal engineering research and teaching. In Proceedings 5th International Conference on Coastal and Port Engineering in Developing Countries. Cape Town, South Africa, pp. 1029–1037. Nielsen, P., 2009. Coastal and Estuarine Processes. World Scientific, Advanced Series on Ocean Engineering – Vol. 29. Nwogu, O., and Demirbilek, Z., 2009. Nonlinear wave transformation and runup over fringing coral reefs. In Coastal Engineering 2008: Proceedings 31st International Conference. Hamburg, Germany: World Scientific, Vol. 1, pp. 242–254. Roberts H. H., 1981. Physical processes and sediment flux through reef-lagoon systems. In Proceedings 17th International Coastal Engineering Conference. Sydney, 1980. American Society of Civil Engineers, Vol. 1, pp. 946–962. Seelig, W. N., 1983. Laboratory study of reef-lagoon system hydraulics. Journal Waterways, Port, Coastal and Ocean Engineering, 109, 380–391. Skotner, C., and Apelt, C. J., 1999. Application of a Boussinesq model for the computation of breaking waves Part 2: Waveinduced setdown and setup on a submerged coral reef. Ocean Engineering, 26, 927–947. Tait, R. J., 1972. Wave set-up on coral reefs. Journal Geophysical Research, 77, 2207–2211. Von Arx, W. S., 1954. Circulation systems in Bikini and Rongelap Lagoons. U.S. Geological Survey Professional Paper, 260-B, 265–273.

Cross-references Cay Formation Fringing Reef Circulation Hydrodynamics of Coral Reef Systems Infrastructure and Reef Islands Lagoons Reef Front Wave Energy Waves and Wave-Driven Currents Wave Shoaling and Refraction

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WAVE SHOALING AND REFRACTION Michael R. Gourlay The University of Queensland, Brisbane, QLD, Australia

Definitions Diffraction. The process by which wave energy is transmitted laterally along a wave crest into a sheltered zone. Most commonly observed behind breakwaters, small islands and exposed reefs which create a sheltered zone behind them. Orthogonal. A line at right angles to the wave crest indicating the direction of travel of the wave. Also known as wave ray. Refraction. (1) The process by which the direction of a wave moving in shallow water at an angle to the bottom contours is changed; the part of the wave crest advancing in shallower water moves more slowly than the part of the crest still advancing in deeper water. (2) The bending of wave crests by currents. Shoaling. The process by which the wave height, length, celerity and other properties of the waves change as they travel from deep(er) water to shallow(er) water. See also Waves and Wave-Driven Currents for definitions of other terms concerning waves. Mathematical symbols are defined where they first appear in the text, or in a relevant figure. The subscript o refers to the deep water values of the various quantities. Introduction When ocean waves propagate around or onto coral reefs, their characteristics change as they interact with the changing seabed morphology. The resulting changes in wave height and direction determine when and where waves break and the patterns of wave crests and wavedriven currents that develop on a given reef. Shoaling of waves The shoaling process occurs as waves approach a straight beach or reef face with their crests parallel to the bottom contours. As waves propagate from deep water into water of decreasing depth h, both their celerity C and length L decrease (see Waves and Wave-Driven Currents). The wave height H also changes, eventually becoming increasingly larger until the waves break. In the simple case of swell propagating from deep water into shallow water with no reflection and no dissipation of wave energy, energy conservation requires that the wave power remains constant (see Waves and Wave-Driven Currents). Using these assumptions with small amplitude (sinusoidal) wave theory gives the following relationship for the variation of wave height as a wave train propagates shoreward toward a beach or over a gently sloping reef face. H Co ¼ ¼ Ks Ho 2nC

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where Ks is the shoaling coefficient and n = Cg/C, where Cg is the group velocity. The parameter n varies with the relative depth h/L or h/Lo. In shallow water, where C = √(gh),
1 8ph 4 Ks ¼ Lo Finite amplitude (cnoidal) wave theory indicates that the shoaling in reality is accentuated as the ocean wave steepness Ho/Lo increases and is much more pronounced than predicted by small amplitude theory (Figure 1). Hence, the latter does not adequately predict wave heights prior to the occurrence of wave breaking, which will occur in shallow water when H/h approaches 0.8.

Refraction of waves In shoaling water when the wave crests cross the depth contours at an angle y, that part of the wave crest which first reaches a given water depth is slowed down relative to the rest of the crest and the wave crest becomes curved. The wave direction is changed and the crests tend to become parallel with the depth contours (Figure 2). For example, when long-crested waves of constant period cross a simple straight step from depth h1 into shallower depth h2 (Figure 3) simple geometry leads to the result sin y2/sin y1 = C2/C1 or sin y/C is a constant. This is the well-known Snell’s Law, which also governs the refraction of light at the boundary between two materials with different refractive indices. Refraction not only causes the wave crests to bend as the wave direction changes but the spacing b between the orthogonals also changes (Figure 2). When waves approach a straight plane beach at an angle, the spacing between orthogonals increases and hence the wave height reduces. Assuming that the wave energy between orthogonals remains constant, this change in wave height is

given by the refraction coefficient KR = √(bo/b) and is additional to that caused by shoaling. Hence, the wave height in shallow water is given by H/Ho = KR.Ks. Refraction thus results in an additional change in wave height over and above that which occurs due to the direct shoaling effect. In general, a decrease in spacing of the orthogonals represents an increase in wave height while an increase in orthogonal spacing represents a decrease in wave height. This results in wave energy being either concentrated behind submarine ridges and on headlands or dispersed behind submarine canyons and in bays. Submerged shoals and reefs usually focus wave energy behind them (Figure 4). Where orthogonals intersect, as in Figure 4b, the simple theory breaks down. Two new processes come into play. Firstly, when there is a large change of wave height along a short length of wave crest, energy will be transferred laterally along the crest by the process known as diffraction which will tend to reduce the concentrating effect of refraction. Furthermore, as the concentrated waves travel into increasingly shallow water they will deform and eventually break, dissipating a significant amount of their energy. Where waves break at an angle to the shoreline, alongshore currents will be generated (see Waves and Wave-Driven Currents). Formerly, wave refraction diagrams, showing the effect of natural bottom topography in changing wave directions and heights within a particular area, were produced manually using graphical procedures. These diagrams are now produced with digital computers using numerical modeling techniques, based upon the differential equation relating the bending of an orthogonal to the change in wave celerity (see Dean and Dalrymple, 1991).

Refraction of waves by currents The previous discussion of wave refraction assumes that the only movement of the water is that caused by the

Wave Shoaling and Refraction, Figure 1 Comparison between shoaling coefficients according to small amplitude (sinusoidal) and finite amplitude (cnoidal) wave theories (from Svendsen and Brink-Kjaer, 1973).

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Wave Shoaling and Refraction, Figure 2 Wave refraction when waves approach a plane beach travelling from deep water into shallow water.

waves and that refraction is caused by changes in wave celerity as waves cross the bottom contours at an angle. However, if there are tidal currents, wind- or wave-driven currents or flow from a river mouth or a stormwater outlet, then the waves will be affected by those currents. If the waves are travelling in the same direction as the current, then their celerity will increase and they will become longer and flatter. If the waves are travelling in the direction opposite to the current, their celerity will decrease and they will become shorter and steeper and breaking may occur earlier than when no current is present. If waves cross a region where the current velocity varies in the horizontal plane, refraction will occur as the wave crests bend under the influence of the different current velocities (Nielsen, 2009).

Wave refraction around reefs Wave refraction around small platform reefs is an essential process in the formation of many reef-top islands. In the simplified examples shown (Figure 5), deep water ocean waves are refracted at the vertical reef face where the wave direction changes abruptly in accordance with Snell’s law. They then propagate on the reef top as shallow water waves. If the waves are large enough to break on the reef rim, they will generate an alongshore current flowing along the reef rim toward the leeward end of the reef platform. Essentially, the reef acts as a lens focussing the wave energy. If the energy is focussed off the reef, i.e., similarly to Figure 4b, then it is not possible for a cay to form on that reef. However, if the shape and orientation of a reef are favorable, the energy will be focussed on the reef platform toward its leeward side (Figure 5a). Stable cays form when the long axis of the reef lens is aligned with the dominant wave direction and the refraction pattern also does not change much with changing ocean wave direction or period. However, it is virtually impossible for a stable cay to form on a circular reef (Figure 5b), since changes

Wave Shoaling and Refraction, Figure 3 Refraction of waves at a step in bottom elevation – derivation of Snell’s Law. dt is the time interval during which a wave crest between two orthogonals completes its crossing of the step.

in wave direction and wave period will change the location of the focal zone. Other nonelongated reefs with sharp corners may experience significant changes in the refraction patterns and hence cay locations change when the dominant waves change direction (Figure 6a and b) (see Infrastructure and Reef Islands, Figure 9, for the consequences of this cay location change). Wave interference zones (intersecting orthogonals) also may occur on small reef platforms (Figure 5a). These are generally located on the windward end of the reef platform as waves from either side of the windward end are refracted toward each other. The larger waves breaking on the windward end of the reef are able to move larger detrital material which, if present in sufficient quantities, will form a shingle bank, possibly with a T-shaped head, which in turn may develop into a windward shingle cay.

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Wave Shoaling and Refraction, Figure 4 Refraction of waves at different coastal features. (a) Concentration and dispersion of wave energy along an irregular shoreline. (b) Focussing of wave energy behind a submerged shoal.

Wave Shoaling and Refraction, Figure 5 Wave refraction on simple reef platforms – effect of reef shape on the formation of cays Wave period 8 s; reef top water depth 2 m; only every fifth wave crest is shown on reef top (from Gourlay, 1988).

All these processes can be seen in a drawing derived from an aerial photograph of Raine Reef in the outer northern Great Barrier Reef (Figure 7). The axis of this reef is aligned in the NW/WNW – ESE/SE direction which corresponds with the dominant southeasterly trade winds. Water depths around this reef are 200–300 m, so the ocean waves are deep water ones. On this occasion, 9 May 1963, the ocean waves were approaching from the east-northeast. These waves were travelling around both sides of the reef platform, intersecting in deep water behind the reef as both wave trains were diffracted into the sheltered zone there. When the wave crests reached the reef edge, they refracted in the very shallow water on the reef top, changing direction significantly particularly along the southern edge of the reef. After breaking on the reef rim they reformed with much shorter wave lengths and travelled across the reef crest onto the reef flat. Under these typical conditions, an interference zone occurs at the eastern end of the reef but no significant

shingle bank has yet formed there on this reef. However, there is a definite focal zone at the western end of the reef and the large stable cay, Raine Island, is located there. Its shoreline is generally aligned with the refracted wave crests, although there is evidence that these waves might be causing some easterly alongshore sediment movement on the northeastern shoreline of the cay and hence, if this persisted, would reshape the eastern end of the cay. In this case, local shoreline realignment is constrained by beach rock outcrops.

Summary The processes of wave shoaling and refraction change the height and direction of waves as they interact with reefs of various shapes and sizes. Significant refraction can occur along the reef edge/rim of an elongated reef aligned with the direction of the approaching ocean waves, As waves break on the reef rim, alongshore currents will transport sediments toward the leeward end of the reef. Diffraction

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Wave Shoaling and Refraction, Figure 6 Wave refraction on simplified form of Gannet Reef, southern Great Barrier Reef – effect of change of wave direction on cay location. Wave period 8 s; reef top water depth 2 m; only every fifth wave crest is shown on reef top (from Gourlay, 1988).

Wave Shoaling and Refraction, Figure 7 Wave crests on Raine Reef, outer northern Great Barrier Reef, on 9 May 1963 – based on an aerial photograph (from Gourlay and Hacker, 1991, Figure 3.8).

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behind the reef will make the sheltered zone behind the reef less sheltered than might be expected. Interference of refracted waves on the reef platform may result in the formation of a windward shingle bank or cay, while sand cays form in the focal zone created by the refracted waves toward the leeward end of the reef. The shape and location of these interference and focal zones vary with changes in wave direction and wave period with consequent modification of the shoreline alignment of cays as their beaches respond to changes in alongshore transport around them.

Bibliography Dean, R. G., and Dalrymple, R. A., 1991. Water Wave Mechanics for Engineers and Scientists. Singapore: World Scientific, Advanced Series on Ocean Engineering – Vol. 2. Gourlay, M. R., 1988. Coral cays: products of wave action and geological processes in a biogenic environment. In Proceedings 6th International Coral Reef Symposium, Townsville, Australia, Vol. 2, pp. 491–496. Gourlay, M. R., and Hacker, J. L. F., 1991. Raine Island: Coastal Processes and Sedimentology. University of Queensland, Civil Engineering Report CH40/91. Nielsen, P., 2009. Coastal and Estuarine Processes. Singapore: World Scientific, Advanced Series on Ocean Engineering – Vol. 29. Svendsen, I. A., and Brink-Kjaer, O., 1973. Shoaling of cnoidal waves. In Proceedings 13th Coastal Engineering Conference, Vancouver, Canada, 1972. American Society of Civil Engineers, New York, Vol. 1, pp. 365–383.

Cross-references Cay Formation Coral Cay Classification and Evolution Fringing Reef Circulation Hydrodynamics of Coral Reef Systems Reef Front Wave Energy Wave Set-Up Waves and Wave-Driven Currents

WAVES AND WAVE-DRIVEN CURRENTS Michael R. Gourlay The University of Queensland, Brisbane, QLD, Australia

Definitions Sea. Waves which are being acted upon by the wind that generated them. They are relatively short and steep with short crest lengths. Also called storm waves. Swell. Waves which are propagating freely no longer under the influence of the wind that generated them. They are relatively long and flat with long crest lengths. Wave celerity C. The velocity of the wave form. Also known as phase velocity. Wave height H. The vertical distance between a wave crest and the following or preceeding wave trough (see Figure 6). Also described as wave amplitude a = H/2. Wave length L. Distance between successive wave crests (see Figure 2).

Wave period T. Time interval between the passing of two successive wave crests at a fixed location. Wave frequency f is the inverse of the wave period ( f = 1/T ). Orbital velocity. Velocity of the water particles during the passing of a wave. Significant wave height Hs. Mean height of the highest 33.3% of all waves in a wave record. Significant wave period Ts. Mean period of the highest 33.3% of the waves in a wave record. Root mean square wave height Hrms. Wave height which is equivalent to the height of a sinusoidal wave with the same energy as the recorded waves. See also Chapters Wave Shoaling and Refraction; Wave Set-Up. Mathematical symbols are defined where they first appear in the text or on a relevant figure. The subscript o refers to the deep water values of the various quantities. The descriptive terms used to describe coral reef morphology vary with the needs and previous practice of the various scientific disciplines. Those used in this entry and related ones (see Chapters Wave Shoaling and Refraction and Wave Set-Up) are defined in Gourlay and Colleter (2005, pp. 355–356) and in Figure 1 below.

Introduction Observations of ocean waves have been made by mariners and engineers for many centuries with empirical relationships being developed for practical use in the nineteenth century when the mathematical study of water waves also began. Scientific study of ocean waves essentially began during World War 2 when knowledge of sea conditions was required at specific locations where amphibious landings were to be made. Coastal engineers developed these new insights about ocean waves in the post World War 2 period using laboratory models and by applying the theoretical and practical knowledge that had been developed during the war. Measurements of ocean waves were first made for engineering purposes such as the design of breakwaters and other structures and for prediction of coastal sediment movements. During the 1960s and 1970s, oceanographers and earth scientists increasingly built on these foundations with field experiments and new theoretical approaches, while, subsequently, engineers and mathematicians have developed mathematical models based upon these measurements and theories. The study of waves on coral reefs is comparatively recent, with the first significant scientific observations being reported in the mid 1970s (Roberts et al., 1975). Ocean waves and coral reefs Ocean waves are one of the most important physical phenomena shaping the morphology and influencing the ecology of modern coral reefs. Waves passing over shallow reefs agitate the water over the reef surface and modify the flow within the coral matrix. When waves break on the edge

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Waves and Wave-Driven Currents, Figure 1 Definition figure for reef morphology terms.

or surface of a reef the water is aerated; pieces of coral are broken off and transported, together with larvae and other organisms; currents are generated on the reef-top; these wave-driven currents carry organisms and detrital material to other locations. Where planar reef platforms have developed, waves create windward shingle banks and islands, including Atoll Islands (motu) and also the sandbanks that develop into leeward Vegetated Cays. There is a large spectrum of waves generated by various causes in the oceans surrounding coral reefs. The shorter period waves are generated by wind blowing over the water surface as various moving weather systems form, grow, and decay. These “wind waves” – sea and swell – have periods from a few seconds to about 20 or more seconds. The predictable tides generated by periodic changes in the gravitational attraction between the earth, moon, and sun are much longer, circa 12 or 24 h (see Chapter Tidal Effects on Coral Reefs). The moving weather systems also produce even longer waves such as the storm surge associated with tropical cyclones (Tropical Cyclone/Hurricane) and continental shelf waves associated with the passage of larger moving weather systems around large land masses such as the Australian continent. Apparently random events such as landslides and earthquakes may generate Tsunamis, long low waves that propagate rapidly over large distances of deep ocean before shortening and amplifying on shallow shelf margins or in confined bays with often disastrous results. This entry is primarily concerned with wind-generated ocean waves and their transformation as they propagate around and onto reefs and into reef lagoons. As these waves transform and break on reefs they also cause wave-driven currents that are superimposed on the tidal and wind-driven current systems. Indeed in many situations the wave-driven

currents dominate the water circulation system on specific reefs. Important aspects of the reef-top wave system are treated in separate chapters – Wave Set-Up; Wave Shoaling and Refraction. Comprehensive treatments of the mechanics of water waves are available in Dean and Dalrymple (1991); Massel (1999) and Nielsen (2009).

Wind-generated ocean waves When wind blows across the surface of the ocean, that surface becomes unstable, developing small waves that propagate in the general direction toward which the wind is blowing. These waves increase in size as the wind speed increases and reach a maximum size, which depends either upon the duration of the wind event or the distance over which the wind is blowing. The propagation of wind waves over the ocean is governed by gravity, which acts to restore the disturbed water surface to its original smooth condition. Waves that are being acted on by the generating wind are known as sea or storm waves, whereas those that have escaped from the influence of the generating wind are known as swell (see definitions of sea and swell). Wave propagation The basic mathematical theory used to represent ocean waves is the small amplitude or Airy wave theory, which represents waves as sinusoidal in form and assumes, among other things, that the wave steepness H/L is small (Figure 2). In its simplest form this is a linear theory which means that, if more than one wave train is present, the water movement of the resulting wave motion can be obtained by superimposing the water movements of the individual wave trains.

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The size of ocean waves is measured in terms of their height H and their length L or period T, where L = CT (Figure 2). C is the celerity, which is the velocity at which an individual wave travels over the water surface. Windgenerated ocean waves are oscillatory waves which transmit wave energy across the ocean surface but do not transport water particles with them. Nevertheless, as waves travel across the ocean they cause the water particles to oscillate about a mean position as each wave passes. Oscillatory ocean waves are progressive, that is, they travel in a given direction. However, if they meet an obstacle – a reef, cliff, or breakwater – any of their energy not dissipated at the obstacle will be reflected as a wave traveling in the opposite direction. The superposition of the reflected wave on the original or incident wave will produce a partial or complete standing wave where the water surface oscillates up and down at distances located at half wave length intervals from the reflective obstacle. If the incident wave train approaches the reflective structure at an angle, it will be reflected with the same angle. Gravity waves propagating in the ocean in most cases are initially deep water waves that propagate without influencing the bottom or being influenced by it. According to small amplitude theory their celerity Co increases with their period, i.e., Co = gT/2p and their wave length Lo = gT 2/2p. This relationship applies when h/Lo > 0.5, where h is the water depth. As the relative water depth h/L reduces below 0.5, both the celerity and wave length are increasingly affected,

Waves and Wave-Driven Currents, Figure 2 Definitions of wave height and wave length.

becoming smaller as h/L reduces, and when h/L  0.05 pffiffiffiffiffi the waves are shallow water waves with C ¼ gh. That is, the wave celerity depends only upon the water depth. The celerity of intermediate depth waves depends upon both the wave period and the relative depth h/L, where L also becomes shorter as the depth becomes shallower (see also Chapter Wave Shoaling and Refraction). The type of motion experienced by the water particles disturbed by waves also changes as the waves propagate from deep water into shallow water. In deep water the water particle orbits are circular in form with a diameter equal to the wave height at sea level and decreasing exponentially with increasing depth below that level (Figure 3). When h/L  0.5, the orbital motion at the bottom is negligible. As waves propagate into increasingly shallow water, the water particle orbits become elliptical in shape with the horizontal motion becoming relatively larger than the vertical motion. In shallow water the horizontal motion is essentially the same at all depths, and there is significant oscillatory water motion at the bottom. In reality individual fluid particles also drift forward in the direction of wave propagation because the positive velocities along the upper parts of their orbits are greater than the negative ones along the lower parts. This process is known as mass transport and the total flow rate per unit length of wave crest generated by it is gH 2/8C (Nielsen, 2009). The various formulae for calculating the properties of waves propagating from deep water into intermediate depths involve hyperbolic functions of the relative depth h/L, where L is one of the unknown properties. Hence, evaluation of these properties involves iterative calculations or the use of tables of previously calculated values. However, explicit approximations, where the wave properties C, L, etc. are functions of h/Lo, are now available to simplify these calculations (Nielsen, 2009). Ocean waves often appear to travel in groups of larger waves followed by groups of smaller waves – fishermen have long held the belief that every seventh or tenth or whatever wave is bigger than the others. Wave theory supports this general idea in that when two wave trains of slightly different wave length travel in the same direction, superposition of their wave forms results in the physical phenomenon known as beats. This involves a periodic variation of the amplitude of the basic wave frequency,

Waves and Wave-Driven Currents, Figure 3 Wave orbital motion: (a) Shallow water, (b) intermediate depths, (c) deep water.

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causing a series of wave groups that move in the same direction as the component waves but at a smaller velocity, the group velocity Cg. Wind-generated ocean waves are short waves in that the effect of bottom friction upon their propagation and energy transmission is confined to a thin oscillatory boundary layer immediately above the bottom surface. By contrast, tide-generated waves, storm surges, and tsunamis are long waves in that they are shallow water waves in the ocean, and the effects of bottom friction create a boundary layer extending over the full water depth. Coral reefs generally have much rougher surfaces than most other coastal or shelf bottoms, and it is possible that in some situations wind-generated waves propagating over reefs may have long wave characteristics. As an ocean wave train propagates across an initially still water surface it deforms that surface and potential energy is stored in the wave form. At the same time, the water particles acquire kinetic energy as they move in their wave-induced orbits. For small amplitude waves the potential and kinetic energies are equal to one another and the total energy per unit water surface area (energy density) E = rgH 2/8. Moreover, energy is transmitted in the direction of wave propagation at the group velocity Cg, which is less than the celerity. Hence, the wave power or energy flux of the wave train, P = ECg. In deep water Cg = Co/2 and so, Po ¼

rg 2 Ho2 T : 32p

In shallow water Cg ! C. The linear wave theory, on which the preceding description of ocean waves is based, assumes that surface gravity waves are regular sinusoidal waves of relatively small amplitude, that is, either H/L or H/h is small. In reality, ocean waves are neither regular nor have small amplitude. Furthermore, some important wave phenomena, such as wave breaking, wave set-up, and the wave-driven currents the latter generates, can only be explained when nonlinear terms associated with the water motion are considered (see Chapter Wave Set-Up).

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Finite amplitude wave theories give alternative representations of ocean waves when either of the conditions H/ L ! 0 or H/h ! 0 do not hold, for example, as waves approach the breaking condition. If H/L ! 1, Stokes theory with various orders of approximation represents deep water waves. First order Stokes theory is identical with small amplitude wave theory. The second order approximation, that is, including terms in H2, has an asymmetric wave profile with the crest amplitude larger than the trough amplitude; crests are sharper and troughs are flatter than sinusoidal small amplitude waves (Figure 4a). It also predicts the mass transport of water particles in the direction of wave propagation. Third order Stokes theory predicts that the wave celerity, which according to small amplitude theory is proportional to the wave period, is also a function of wave steepness. In shallow water, when the relative depth h/L ! 0, a second order solution of the linear shallow water long wave equations results in the cnoidal wave theory. Again these finite amplitude waves have sharp higher crests and flat shallower troughs than sinusoidal small amplitude waves (Figure 4b). Their celerity depends not only on the water depth, but also on the relative wave height H/h. Stokes theory becomes increasingly unsatisfactory as waves enter intermediate depths, that is, h/L < 0.5, while cnoidal theory also is increasingly unsatisfactory as h/L increases above 0.05. See Nielsen (2009) for further discussion of the applicability of these theories.

Wave breaking Wave breaking occurs when oscillatory waves become unstable and are transformed into translatory waves that transport water in the direction of wave propagation. This will occur when the horizontal orbital velocity of the water particles at the wave crest exceeds the wave celerity or when their maximum vertical acceleration exceeds the gravitational acceleration. In deep water, ocean waves will break when their wave steepness Ho/Lo approaches 0.142 (or 1/7). This process is a gradual one and appears in the formation of increasing numbers of “white caps” on the sea surface as the wave crests spill and develop small rollers on their downwind side. Spilling breakers of similar form

Waves and Wave-Driven Currents, Figure 4 Wave profiles for finite amplitude waves: (a) Stokes waves in deep water, (b) Cnoidal waves in shallow water.

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occur when steep waves break on mildly sloping beaches. Less steep waves breaking on steeper beaches peak up, their forward face becomes vertical, and their crest plunges forward like a jet into the preceding wave trough. On even steeper beaches the lower part of the forward face of the wave becomes vertical and collapses like a partial plunging wave. Surging breakers occur when waves of low steepness break on steep beaches. The forward face of the wave remains relatively smooth and the wave runs up the beach with little foam or turbulence. The uprush–backwash cycle of a surging breaker is completed before the next wave surges up the beach (Figure 5). The region where waves break is known as the surf zone. It may be divided into three subzones. There is an outer zone where the breaking waves rapidly change shape and develop a surf roller form. Potential energy is transferred into kinetic energy with little loss of energy. In the inner zone the surf roller becomes a bore or moving surge traveling landward over the seaward flowing undertow. Energy dissipation and Wave Set-Up occur in this zone. The swash zone occurs on the beach face where uprush and backwash of the breaking waves define the run up and run down limits (see Chapter Wave Set-Up, Figure 1).

Waves and Wave-Driven Currents, Figure 5 Beach and breaker types for plane beaches. Ib = Hb1/2/g1/2T tan a, where Hb is breaking wave height and tan a is beach slope (reprinted from Gourlay, 1992, with permission from Elsevier).

Irregular waves Mariners and others accustomed to observing and experiencing ocean waves have been aware that the periodic regular waves conceived by mathematicians are an inadequate model for wind waves. Visual observations show that the height and period of successive individual waves vary considerably. Moreover, visual observations of wave height have been recognized as giving an estimated wave height somewhere between the actual mean wave height Hav and the largest waves Hm occurring at the time of observation. The advent of wave recorders produced many short records of a hundred or more waves of different heights and periods. The concepts of significant wave height Hs, root mean square wave height Hrms, significant period Ts, and zero crossing period Tz (Figure 6 and definitions) have been developed to characterize the waves in a given wave record. Statistical theory provides a theoretical wave height distribution – the Rayleigh distribution – which relates Hs, Hrms, and Hav to the percent of waves in the record equaling or exceeding a given height. Hrms (=Hs /√2) is the wave height equivalent to a sinusoidal wave with the same energy as the recorded waves. The maximum wave height in a given record follows an extreme value distribution, increasing as the number of waves recorded increases. Using Hs to characterize the short term statistics of each wave record of say 20-min duration, the long-term statistical occurrence of wave heights can be obtained by using an exceedence plot for Hs at a given location, in which a given value of Hs is plotted against the percent of time it is equalled or exceeded. However, a more fundamental and physically meaningful long-term wave statistic would be one based upon the maximum value of Hs during various independent storm events. The sea surface is continually changing. Since the wave celerity in deep water depends upon the wave length or wave period, the various waves present are continually moving in and out of phase with each other. The resulting sea surface is quite irregular and often consists of alternating groups of high and low waves. Moreover, the highest waves do not persist but are randomly distributed in space

Waves and Wave-Driven Currents, Figure 6 Definitions of wave height and period (Tz) in a record of irregular ocean waves.

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and time. Such a sea surface can be approximated by the superposition of a large number of small amplitude wave trains from different directions. This leads to the concept of the wave spectrum in which the energy associated with a given frequency band has been determined for the full range of frequencies present in a given wave record (Figure 7). If the area under the wave spectrum diagram is designated as mo, it can be shown that Hs = 4√mo, so measurement of the wave spectrum gives an alternative estimate of the significant wave height Hs. The period Tp, corresponding to the peak of the frequency spectrum, is commonly used as the characteristic period for a given wave condition. For wind-generated storm waves, the spectrum is broad, and Tp corresponds with a celerity approaching

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the wind speed generating the waves. Swell waves have narrower spectra and larger values of Tp than storm waves. In open ocean and other situations, where locally generated wind waves are superimposed upon underlying swell, the wave spectra will have at least two peaks and hence two values of Tp (Figure 7). A useful means of showing the variation of ocean wave conditions at a given location for a period of time is an Hs–Tp scatter diagram (Figure 8). Such a diagram commonly will show an upper limiting curve where Hs / Tp2, representing the maximum wind waves (sea) that can be generated at that site. In open waters this diagram also has a long tail with relatively low Hs but increasing Tp representing swell that has come from distant storms.

Waves and Wave-Driven Currents, Figure 7 Ocean wave spectra (minor irregularities of spectral profiles have been removed) (A) Storm waves Hos = 4 m, Tp = 9.5 s; (B) wind waves Hos = 2.1 m, Tp = 7 s; (C) swell Hos = 1.7 m, Tp = 14 s; (D) sea and swell Hos = 1.5 m, Tp1 = 10 s, Tp2 = 4.4 s.

Waves and Wave-Driven Currents, Figure 8 Hos vs. Tp scatter diagram for ocean waves on southern side of Heron Reef, southern GBR: November 1996 to March 1997 (from Gourlay and Hacker, 2008a, p. 111).

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Wave transformation on reefs “The interaction of physical processes with coral reefs is complex; each reef system has its own unique set of environmental and morphologic conditions” (Roberts et al., 1975, p. 234). The characteristics of ocean waves depend upon the climatic conditions generating them, and their transformation as they propagate over reefs is significantly influenced by the morphology of the reef system and adjacent or adjoining land masses. Moreover, wave transformation and accompanying wave driven-currents occur in water depths that vary with the periodic water level variations caused by the tides. Hence, the wave driven-currents are superimposed upon and interact with those generated by the tides, as well as with wind-driven currents. Climate, reef morphology, and tides all vary in different parts of the world and so do wave interactions with local coral reefs. Reef morphology may vary from relatively wide fringing reefs with a flat seaward face fronting a beach to a narrow strip of steep-faced reef backed by a lagoon with a beach or reef behind it. Offshore reefs may be twodimensional barrier reefs or three-dimensional island reefs usually with steep faces. Atolls have relatively narrow reefs with steep faces, enclosing a lagoon of sheltered water. Narrow reef islands (motus) prevent waves reaching the atoll lagoon except through the gaps in between the islands or passages through the reef. Some reefs are submerged at all times, and in many cases the coral growth on their surface provides a rough porous matrix which interacts strongly with the waves. In other cases reefs have reached the sea surface and have developed relatively smooth planar reef tops with different zones: smooth coralline algae, coral with many depressions but few projections, and even sandy reef flats. Where there is significant tidal variation of water levels the planar reef surface is generally at about the same elevation as the mean low tide level. Several types of waves break on or travel around or across reefs. First, there are wind waves and swell from the surrounding ocean. Then there are short period waves, locally generated in shallow lagoons on the reef top or in larger deeper lagoons of atolls. Waves breaking on the reef rim release both shorter (secondary) waves and longer (infragravity) waves that alter the properties of waves traveling across the reef top. Wave transformation over a reef varies depending upon which part of the reef is under consideration. Five zones can be distinguished. These are as follows: (1) The outer reef face where waves are affected by shoaling water but may not break (2) The reef edge where the larger waves generally break (3) The reef rim, between the reef edge and the reef crest, where smaller waves break and the breaking surge dissipates most of its energy (4) The reef flat or lagoon where the waves reform (5) The beach of either a reef island, continental island, or mainland, where the reformed waves finally break

On the outer reef face waves are subject to the processes of Wave Shoaling and Refraction. However, reef face slopes are very variable and so are their effects on waves. In some cases such as the ribbon reefs on the outer northern Great Barrier Reef (GBR) the reef face is almost vertical (Young, 1989). In these situations the effects of refraction and shoaling before breaking will be small, but significant reflection of wave energy may occur. Energy dissipation prior to breaking will be negligible but the actual breaking action on the reef edge will be very intense. Both field and laboratory observations indicate that on reefs with very steep faces the type of breaker is different from those observed on normal sloping beaches, being neither a plunging breaker nor a spilling breaker. The waves do not peak up significantly before breaking nor do they plunge but rather collapse into the strong backwash flowing seaward from the reef top. On the flatter seaward faces of the Caribbean fringing reefs the bottom is much more irregular than on a sandy coast, and so significant energy dissipation may occur before the waves break (Roberts et al., 1975). Even where the reef face is steep but the water depth over the reef crest is relatively small, breaking waves may shape the seaward side of the reef crest into a ramp with a slope of the order of 1 in 8 to 1 in 15 where the waves break and dissipate a significant amount of energy before the breaking surge reaches the reef crest. In other situations the seaward part of the reef rim may develop a particular morphology created by the interaction between the waves and the growing reef structure. This is the characteristic spur and groove system (Spurs and Grooves) often found on the windward edges of reefs. Normally waves break on the reef rim unless the water depth over it is large. For any given water depth over the reef the wave height on the reef increases with increasing ocean wave height up to a maximum limiting value. Thereafter, any further increase in ocean wave height does not increase wave heights on the reef. Indeed they may even reduce somewhat (Kono and Tsukayama, 1980). In some situations waves will pass over the reef rim at higher tide levels and travel across the reef flat or through the reef lagoon without breaking until they reach a cay or mainland beach where they finally break. Submerged reefs may only experience breaking wave conditions with large waves, and all wave action may pass over the deeper reefs without any breaking. The width of the surf zone on the rim of a near horizontal reef is of the order of a few wave lengths. Laboratory data (Nelson and Lesleighter, 1985; Gourlay, 1994) show that the initial breaking process on a reef with a steep face is completed within one wave length from the reef edge. The distance from the reef edge to the end of the surf zone where the waves reform varies from three wave lengths upward depending upon the ratio between the ocean wave height and the water depth over the reef. For typical conditions within the GBR region, for example, waves of 3 m height and 6 s period breaking on a reef with 2 m of water over it, this zone would be 100–125 m wide. In exposed conditions on the outer barrier with a 3 m swell of 10 s

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period in the same water depth the surf zone would be 160–200 m wide. Extreme waves of 8 m height and 10 s period in the same depth would create a surf zone 300–400 m wide. The height of the reformed waves at the leeward side of the surf zone is limited by the water depth. Field studies at John Brewer Reef in the central GBR indicate that after the completion of wave breaking the maximum significant wave height is 0.35–0.4 times the water depth over the reef flat and that the maximum wave height in most cases does not exceed 0.6 times the water depth (Hardy and Young, 1996). Laboratory data (Nelson, 1987; Gourlay, 1994) confirm that the maximum H/h value for waves propagating over horizontal bottoms or very flat slopes is about 0.55. However, higher order approximations of nonlinear wave theories indicate that, when waves propagate in water of constant depth, the maximum possible value of H/h is close to 0.78. This value is commonly used in coastal engineering practice to determine where waves break on beaches outside coral reef regions. When the presence of higher harmonics generated on a horizontal bottom or at an abrupt depth change is considered, wave theory predicts a maximum H/h < 0.6 (Massel, 1996). Field observations on various reefs in different parts of the world show that waves are very much changed as they travel across a reef flat. Not only do wave heights become smaller because of breaking and frictional resistance but the wave spectrum broadens as energy is transferred to both higher frequencies (shorter periods) and lower frequencies (infragravity waves). Breaking waves are highly nonlinear, and secondary waves, which are harmonics of the primary wave period, develop during the shoaling and breaking process. Consequently, the reformed waves generally have shorter periods than the incident ocean waves as the primary wave and one or more secondary waves travel across the reef (Figure 9). This may be particularly the case for long swells (Wiegel, 1990). The lower frequency infragravity waves are responsible for the surf beat that causes fluctuations in the mean water level (dynamic set-up) in the surf zone (see Chapter Wave SetUp). Where water is ponded on the reef flat at low tide, local winds may generate short period waves (1 < T < 3 s) on the ponded water. Significant energy losses occur as waves propagate across coral reefs. Observations in both the field and laboratory models indicate that the energy dissipation of waves breaking on coral reefs varies between 72 and 97% of the incoming wave energy (Roberts et al., 1992; Gourlay, 1994; Brander et al., 2004). This energy dissipation is caused by both bottom friction and wave breaking (Gerritsen, 1981). As reef surfaces are much rougher than those of sandy beaches by a factor of at least 10–20 times (Nelson, 1996), a relatively greater proportion of energy dissipation over coral reefs is caused by bottom friction. With relatively small waves passing over a shallow reef, friction dissipates a larger proportion of the wave energy than wave breaking (Lowe et al., 2005). Bottom friction also will have a relatively greater effect than wave breaking

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Waves and Wave-Driven Currents, Figure 9 Multiple crested secondary waves on reef flat, Rarotonga, Cook Islands, South Pacific Ocean. Photo: Michael Gourlay, March 1995.

on flatter reef face slopes and wider reef flats. As the size of the waves increases the proportion of wave energy dissipated by breaking increases, and eventually dissipation by wave breaking will dominate when waves are large. While the beaches of reef-protected shorelines vary in form from place to place, these beaches are often relatively steep (tan a > 0.1) with a plane face and are formed of medium to coarse coral sand. At low tide, no waves reach the beach since the waves break on the reef rim. However, as the tide rises, some reformed waves travel across the reef flat or shallow lagoon and break directly on the beach face. If these waves are steep ones they will plunge onto the beach and their energy will be dissipated in the uprush–backwash cycle on the upper beach face. If, on the other hand, waves of low steepness, such as long swells, pass over the reef rim and the reef flat without decomposing into secondary waves, the breakers will surge up the beach with little energy being dissipated. The latter waves have considerable potential to overtop the beach crest and move sand landward, so increasing the height of the beach crest. Moreover, on narrow reef flats the longer period infragravity waves can produce resonant oscillations of water level, which dominate the wave uprush process on a reef-protected beach (see Chapter Wave Set-Up). In a study based upon three coral reefs with different morphologies and hydrodynamic (wave and tidal) conditions, Kench and Brander (2006) found that the effect of incident wave energy in reshaping reef-protected beaches depends upon a reef energy window index c = (mean reef flat water

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depth at high spring tide)/(reef width). A narrow reef width with a low reef flat elevation has a large value of c, indicating that more wave energy will cross the reef and break on the island shoreline, whereas a wide reef with a high reef flat elevation has a low value of c, indicating that most wave energy will be dissipated on the reef rim and relatively little of it will reach the island. Rising sea levels and disintegration of reef top corals following coral bleaching events have the potential to increase the depth of water over a reef flat, thus increasing c and making reef islands more vulnerable to erosion (Sheppard et al., 2005). In the tropical latitudes where coral reefs flourish, winds and waves vary seasonally. For example, South Maalhosmadulu Atoll in the Maldives lies at latitude 5 N. It is subject to both westerly and northeasterly monsoon conditions with the former lasting longer (8 months) and also producing greater incident wave energy on the atoll and its reefs than the latter. Underlying swell from the south also refracts around the atoll delivering greater energy to its western side, independently of monsoon conditions. On an atoll scale these seasonally changing wave energy conditions control the formation, shape, and location of reef islands on the various small reefs forming the atoll, whereas on a reef scale they cause seasonal fluctuations in the positions of island beaches (Kench et al., 2006). Waves on the windward sides of reefs and cays are generally larger and shorter than those on the leeward sides, resulting in higher windward reef flat and beach ridge elevations (Kench et al., 2009; Samosorn and Woodroffe, 2008). At Raine Reef, one of the outer reefs of the northern GBR (latitude 11 360 S), southeasterly trade winds dominate during the winter months and northwesterly monsoon winds occur during most summers. Swells from distant tropical cyclones in the Coral Sea also reach this reef during summer, and occasional, generally small, tropical cyclones may affect it directly. Wave transformation across this relatively small reef platform is very much a three-dimensional process (see Figure 7 in Chapter Wave Shoaling and Refraction) and has resulted in the formation of a large cay located at the leeward end of the reef under the influence of the dominant southeasterly waves. Northwesterly monsoon waves transport sediment eastward along the northern beach, often creating a recurved sand spit at the eastern end of the island. During the subsequent winter southeasterly waves transport much of the sand from this spit along the southern shore of the island to its southwestern corner. Hence, seasonal variations in winds generate waves that cause a clockwise movement of sand around this reef island. Many of the field experiments involving waves breaking and propagating across coral reefs have been undertaken at relatively sheltered locations or at times of low energy conditions. The duration of measurements in most cases is no more than a week or two. Conclusions based upon the experimental data obtained from such experiments may not be valid when more energetic conditions, that is, cyclonic/hurricane wave conditions, occur. In many cases significant sediment movement and morphological changes on reefs or reef-protected shorelines only occur during

extreme events and even then only when reef top water levels are elevated above normal levels by wave set-up, high tides, storm surge, or tsunami.

Wave-driven currents Surfers and fishermen have been aware for a long time of the existence of wave-driven currents, such as alongshore currents, rip currents, and undertow, within the surf zone of beaches. Earth scientists and engineers concerned with coastal structures also have been aware of the transport of sand and gravel along coasts by waves and the problems this transport can cause when breakwaters are constructed to form new harbors or to stabilize river entrances. However, it was not until mathematical physicists such as LonguetHiggins and Stewart (1964) extended the sinusoidal (small amplitude) wave theory to include the concept of wave thrust or, as they described it by analogy with electromagnetic waves, radiation stress, that a proper scientific understanding of the mechanism driving the various surf zone currents was developed. This theoretical analysis also predicted that the waves would produce changes in the mean water level as they approached a beach: wave set-down offshore of the surf zone and wave set-up within the surf zone (see Chapter Wave Set-Up). Initially wave-driven current systems were observed on fringing reefs on various islands in the Caribbean and Pacific regions. In some locations the reefs were relatively narrow and were backed by lagoons so that the reefs functioned as submerged breakwaters. At Grand Cayman Island in the Caribbean the waves discharged water over the reef rim into the lagoon where it flowed parallel to the reef crest and shoreline and discharged seaward through a channel at the end of the reef (Roberts et al., 1975; Roberts, 1981). Similar wave-driven current systems were observed at Guam in the northern Pacific Ocean where the water returned to the ocean through channels in the reef rim, associated with rivers (Marsh et al., 1981). Tidal ranges were generally less than 1 m. The extensive Ningaloo Reef is located in the Indian Ocean on the edge of the continental shelf close to the Western Australian coast. Waves breaking on this westward-facing reef were observed to generate a landward flow across the reef with return circulation in the lagoon between the reef and the mainland. Water returned seaward through channels through the reef north and south of the inflow zone. The spring tide range was almost 2 m (Hearn et al., 1986). In most of these cases, waves breaking on the reef rim transport water landward across the reef flat into shallow lagoons between the reef and the mainland. The flow then runs parallel to the shoreline until it reaches a gap in the reef or the end of the reef, where it flows seaward back to the ocean. Where the outflowing current is sufficiently strong, sediments on the reef flat or lagoon bottom are likely to be transported offshore and deposited in deep water sinks from which they cannot be returned to the reef system (Roberts, 1981). Navigation channels dredged

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through the rim of coral reefs always carry this risk, unless specific action is taken to prevent it occurring. The boat access channel at Heron Island in the southern GBR is an example of this type of situation (Gourlay and Hacker, 2008b) (see Chapter Infrastructure and Reef Islands). On offshore barrier or platform reefs, waves breaking on the windward reef rim transport water across the reef flat either over the leeward reef rim or in the case of atolls into the lagoon behind the reef. Such currents have been observed at John Brewer Reef and other locations in the GBR system (Young, 1989; Hardy et al., 1991). The velocities of these wave-driven currents have been found to be approximately proportional to the height of the waves causing them (Hearn and Parker, 1988; Symonds et al., 1995). When waves approach the reef rim at an angle, they will be refracted as they travel across the sloping reef face and break either there or on the reef rim. Breaking waves that cross the reef rim at an angle will generate an alongshore current similar to that generated when waves break at an angle to a beach (see Chapter Wave Shoaling and Refraction). The actual directions of flow of wave-driven currents on reef tops depend upon several factors, including the direction of the ocean waves relative to the reef platform, the water depths over the reef top, the reef morphology, the roughness of various zones on the reef, etc. In general, currents flow from the breaker zones where wave set-up is highest to the lower portions of the reef top or to gaps in the reef surrounding a lagoon. Laboratory models of small platform reefs in the GBR demonstrated these principles quite clearly (Gourlay, 1993; 1995). A model of Raine Reef on the northern outer barrier showed how wave set-up at the windward eastern end of the reef drives a current down the center of the reef and around the coral cay at its leeward end. A model of North Reef in the southern Capricorn region showed how the presence of a windward shingle bank obstructs the longitudinal flow, particularly at low tide levels, forcing it to follow a path defined by the lower levels on the reef top.

Modeling wave-driven flow across reefs Researchers modeling wave-driven flow across reefs have assumed a two-dimensional reef profile with a seaward sloping reef face and a horizontal reef top. Waves break on the sloping reef face, dissipating their energy and generating wave set-up. The maximum set-up
r occurs at or near the reef edge, that is, at the seaward side of the horizontal reef top. This maximum set-up creates the pressure gradient required to drive the wave-driven flow across the reef top since the water level at the downstream side of the reef is assumed to be the same as that seaward of the surf zone. The various approaches for modeling wave-driven flow across reefs have been reviewed by Gourlay and Colleter (2005) and Monismith (2007). The mathematical models are based on the depthintegrated (one-dimensional) equations of motion for steady free surface flow, that is, the momentum equation incorporating the wave thrust, and the continuity equation.

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The various researchers make different assumptions in applying these equations to wave-driven flow across a reef. Symonds et al. (1995) assume that
 1.5 m), continuous inflow to the harbor occurs throughout the tidal cycle (Figure 14). For large waves and small tides the westward wave-driven flow varies with the waves rather than the tide. For large tides and smaller waves, the westward flow velocity varies with the tidal cycle and the maximum velocity occurs when weir control develops over the bund walls at low tide. Under commonly occurring wave (Hos < 1.5 m) and tidal conditions, the maximum wave-driven currents on the reef top were usually no larger than 0.3–0.4 m/s. However, during the three storm events recorded during this 12-month period ocean waves reached 3 m height and currents were of the order of at least 0.6 m/s (Figure 14). The observations of wave-driven currents on Heron Reef confirmed the prediction of Gourlay and Colleter’s

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Waves and Wave-Driven Currents, Figure 13 Wave-driven flow into Heron Island boat harbor, southern GBR (reprinted from Gourlay and Colleter, 2005, with permission from Elsevier).

theory that the wave-driven current velocity increased with increasing Hos and decreasing reef top water depth. No consistent relationship was found between current velocity and wave period. A direct comparison of measured and calculated current velocities at the northern current meter site for tide levels when the tidal current was zero gives a relationship of the form V = C(Hos  Hoscr), where Hoscr = 0.30.5 m is the threshold condition for a wavedriven current and the constant C is independent of tidal range (Gourlay and Hacker, 2008b, pp. 222–249). The measured velocities are somewhat lower than the calculated ones, and this could be attributed to the reef top being rougher than it was assumed to be and/or to the ocean wave heights used in the calculation not being reduced to allow for refraction as they propagated from the deeper water surrounding the reef platform onto the shallow reef top. Most recently Hench et al. (2008), working at Moorea, French Polynesia, have observed wave-driven currents on a coral reef with a shallow lagoon separating it from a continental island and a narrow deep passage connecting the lagoon to the ocean. Tidal ranges during these observations were small (0.3 m), so the water circulation was primarily the result of the set-up caused by waves breaking on the reef. Hench et al. were unable to predict their observed wave set-up using either the theory of Hearn (1999) or that of Gourlay and Colleter (2005), but an empirical relationship was established between wave setup on this reef and the wave height, wave period, and

the water depth over the reef. Circulation and exchange of water between this coral reef system and the adjacent ocean was largely determined by episodes of larger waves generated by remote weather events rather than periodic tidal exchange mechanisms. Field measurements of both wave set-up and the levels of the reef and beach topography where they are made must be accurately related to a common fixed land datum. If this is not done, the water depths cannot be determined to sufficient accuracy to reliably compare observed values of set-up and current velocities with those derived from various theories.

Summary Knowledge of ocean waves comes from observation and measurements in the field, laboratory experiments, mathematical theories, and numerical models. Engineering works, modern warfare, and environmental management have all stimulated and contributed to scientific research concerning ocean waves and their influence on coral reefs. Waves are one of the most important physical phenomena shaping the morphology and influencing the ecology of modern coral reefs and reef top islands. There is a large spectrum of ocean waves, ranging from wind waves, swell, tides, storm surges to tsunamis. This entry is concerned with wind waves and swell. Small amplitude wave theory is a first order linear theory for oscillatory waves of sinusoidal form. It provides

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Waves and Wave-Driven Currents, Figure 14 Wave-driven currents dominate over tidal currents at Heron Reef, southern GBR – 27 July, 1996 (adapted from Gourlay and Hacker, 2008b, p. 179).

relationships between the wave length, wave period, wave celerity, and wave height as waves propagate from deep water, where waves are unaffected by the bottom, through intermediate depths, where the depth has an increasing effect, into shallow water, where the depth dominates the propagation of the waves (see also Chapter Wave Shoaling and Refraction). As waves propagate across the water surface, water particles disturbed by them move in closed orbits with negligible forward transport but energy is transmitted by the waves with a velocity less than the wave celerity.

Finite amplitude wave theories give a better representation of the form of real waves, that is, the peaks are sharper and the troughs are flatter than those of sinusoidal waves. Stokes theory applies for deep water waves, and the second order theory predicts a small net movement of water particles or mass transport in the direction of wave propagation. Cnoidal wave theory represents shallow water finite amplitude waves. Waves break when their crests become unstable. In deep water this appears as white caps at the front of the

WAVES AND WAVE-DRIVEN CURRENTS

wave crests. In shallow water there are a series of breaker forms, spilling, plunging, collapsing, and surging, depending upon the steepnesses of both the bottom and the waves. The region between the break point and the shore is known as the surf zone, the final section of which is the swash zone on the beach face. Real ocean waves are not periodic but irregular. The characteristics of real waves measured by a wave recorder are represented by the significant wave height or root mean square wave height, which are related to one another by the Rayleigh distribution. An alternative or complementary representation of irregular waves is the wave spectrum that measures amount of energy at different frequencies in a given short-term wave record. The interaction of physical processes on coral reefs is complex, and wave transformation occurs in water depths that are constantly changing with the tides in different locations with varied reef morphology. The types of waves propagating on reefs include sea and swell from the surrounding ocean, short locally generated wind waves, secondary waves, and infragravity waves. Waves are affected by the rough reef face before breaking on the reef rim, after which they reform and travel across the reef flat and ultimately may break on the beach of a cay or other land mass. For any given water depth over a reef there will be a limiting wave height that can cross the reef without breaking. Both field and laboratory experiments indicate that the significant wave height of reformed waves on a reef does not exceed 0.35–0.4 times the water depth on the reef, and the maximum individual wave height is no more than 0.55–0.6 times the water depth. Reef surfaces are 10–20 times rougher than sandy beaches, and 72–97% of wave energy is dissipated as waves propagate across reefs. High and wide reefs dissipate much more energy than low and narrow reefs. Winds and the waves they generate vary seasonally, and these variations have significant implications for the formation, shape, and location of reef islands on individual reefs. Wave Set-Up causes variations in mean water level on reefs, and the resulting water level gradients drive reef top current systems. Wave-driven currents flow from the exposed reef edge, where waves are breaking, toward the sheltered side of the reef and/or to lower parts of the reef top or atoll reef. Outflowing currents through both natural and artificial channels have the potential to remove sediment from the reef. Mathematical models of wave-driven currents are based upon the depth-integrated equations of motion for steady free surface flow, that is, momentum equation incorporating wave thrust, and continuity equation. These models predict that the wave-driven flow over a reef increases from zero, as the ocean water level rises above the reef top level, to a maximum with increasing water depth and then decreases as wave breaking and wave set-up reduce, becoming zero again when waves pass over the reef without breaking. Such models also can simulate wave-driven circulation in the presence of tidal circulation. However, they are restricted to reefs with relatively flat reef face slopes and generally

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do not produce meaningful results when water depths over the reef are very shallow. The wave pump analogy regards the waves breaking on the reef as a pump lifting water from the ocean and discharging it onto the reef top or into the lagoon behind the reef. The pumped water then flows back over the leeward side of the reef in accordance with open channel flow theory. The wave pump can be applied to and is most efficient on steeper faced reefs. It can also be applied where reef top water levels are very shallow or even emergent, that is, above ocean water level. Reef top characteristic curves for a given reef show the increase in wave-driven flow from zero at very small depths to a maximum at much larger depths and then decreasing to zero when waves pass over the reef without breaking. They also show the different flow conditions when reef top control and reef rim control occur and the critical submergence when one form of control changes to the other. Field measurements show that wave pumping provides a very effective mechanism for flushing atoll lagoons, particularly those that are completely enclosed by a reef with no channels through its rim. In open ocean situations, even when tides are relatively large (range 3 m), wave-driven currents significantly influence the reef top circulation system on platform reefs. On a particular reef, under commonly occurring wave (1.5 m height) and tidal conditions, the incoming tidal flow was completely reversed by wave-driven flow even though the maximum wave-driven currents on the reef top were usually no larger than 0.3–0.4 m/s. However, during storm events with ocean waves 3 m height, currents were at least 0.6 m/s. When tidal currents were insignificant, the wave pump theory gave a good prediction of reef top current velocities.

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Gourlay, M. R., 1994. Wave transformation on a coral reef. Coastal Engineering, 23, 17–42. Gourlay, M. R., 1995. Wave-generated currents at Raine island: laboratory model experiments. University of Queensland, Department of Civil Engineering. Report CH44/95. Gourlay, M. R., 1996a. Wave set-up on coral reefs. I. Set-up and wave-generated flow on an idealised two dimensional horizontal reef. Coastal Engineering, 27, 161–193. Gourlay, M. R., 1996b. Wave set-up on coral reefs. II. Set-up on reefs with various profiles. Coastal Engineering, 28, 17–55. Gourlay, M. R., and Colleter, G., 2005. Wave-generated flow on coral reefs – an analysis for two-dimensional horizontal reeftops with steep faces. Coastal Engineering, 52, 353–387. Gourlay, M. R., and Hacker, J. L. F., 2008a. Reef-top currents in vicinity of Heron Island Boat Harbour, Great Barrier Reef, Australia. I. Overall influence of tides, winds and waves. University of Queensland, Civil Engineering Report CH72/08. http:// espace.library.uq.edu.au/view/UQ:159070. Gourlay, M. R., and Hacker, J. L. F., 2008b. Reef-top currents in vicinity of Heron Island Boat Harbour, Great Barrier Reef, Australia. II. Specific influences of tides, meteorological events and waves. University of Queensland, Civil Engineering Report CH73/08. http://espace.library.uq.edu.au/view/UQ:159075. Hardy, T. A., and Young, I. R., 1996. Field study of wave attenuation on an offshore coral reef. Journal of Geophysical Research, 101(C6), 14311–14326. Hardy, T. A., Young, I. R., Nelson, R. C., and Gourlay, M. R., 1991. Wave attenuation on an offshore coral reef. In Proceedings 22nd International Coastal Engineering Conference, Delft, July 1990. New York: American Society of Civil Engineers, Vol. 1, pp. 330–344. Hearn, C. J., 1999. Wave-breaking hydrodynamics within coral reef systems and the effect of changing relative sea level. Journal of Geophysical Research, Series C, 104, 30007–30019. Hearn, C. J., and Parker, I. N., 1988. Hydrodynamic processes on the Ningaloo Coral Reef, Western Australia. In Proceedings of the 6th International Coral Reef Symposium, Townsville, Australia, August, 2, pp. 497–502. Hearn, C. J., Hatcher, B. G., Masini, R. J., and Simpson, C. J., 1986. Oceanographic processes on the Ningaloo Coral Reef, Western Australia. University of Western Australia, Centre for Water Research, Report Number: ED-86-171. Hench, J. L., Leichter, J. J., and Monismith, S. G., 2008. Episodic circulation and exchange in a wave-driven coral reef and lagoon system. Limnological Oceanography, 53, 2681–2694. Kench, P. S., and Brander, R. W., 2006. Wave processes on coral reef flats: implications for reef geomorphology using Australian case studies. Journal of Coastal Research, 22(1), 209–223. Kench, P. S., Brander, R. W., Parnell, K. E., and McLean, R. F., 2006. Wave energy gradients across a Maldivian atoll: implications for island geomorphology. Geomorphology, 81, 1–17. Kench, P. S., Brander, R. W., Parnell, K. E., and O’Callaghan, J., 2009. Seasonal variations in wave characteristics around a coral reef island, South Maalhosmadulu atoll, Maldives. Marine Geology, 262, 116–129. Kono, T., and Tsukayama, S., 1980. Wave transformation on reef and some consideration on its application to field. Coastal Engineering in Japan, 23, 45–57. Kraines, S. B., Yanagi, T., Isobe, M., and Komiyama, H., 1998. Wind-wave driven circulation on the coral reef at Bora Bay, Miyako Island. Coral Reefs, 17, 133–143. Longuet-Higgins, M. S., and Stewart, R. W., 1964. Radiation stress in water waves, a physical discussion with applications. DeepSea Research, 11, 529–562. Lowe, R. J., Falter, J. L., Bandet, M. D., Pawlak, G., Atkinson, M. J., Monismith, S. G., and Koseff, J. R., 2005. Spectral wave

dissipation over a barrier reef. Journal of Geophysical Research. 110, C04001. Lugo-Fernandez, A., Roberts, H. H., and Wiseman, W. J., 2004. Currents, water levels, and mass transport over a modern Caribbean coral reef: Tague Reef, St Croix, USVI. Continental Shelf Research, 24, 1989–2009. Marsh, J. A., Ross, R. M., and Zolan, W. J., 1981. Water circulation on two Guam reef flats. In Proceedings of the Fourth International Coral Reef Symposium, Manila. Vol. 1, pp. 355–360. Massel, S. R., 1996. On the largest wave height in water of constant depth. Ocean Engineering, 23, 553–573. Massel, S. R., 1999. Fluid Mechanics for Marine Ecologists. Berlin: Springer. Massel, S. R., and Brinkman, R. M., 2001. Wave-induced set-up and flow over shoals and coral reefs: Part 1. Oceanologia, 43(4), 373–388. Monismith, S. G., 2007. Hydrodynamics of coral reefs. Annual Review of Fluid Mechanics, 39, 37–55. Munk, W. H., and Sargent, M. C., 1954. Adjustment of Bikini Atoll to ocean waves. U.S. Geological Survey Professional Paper, 260-C, pp. 275–280. Nelson, R. C., 1987. Design wave heights on very mild slopes – an experimental study. Australian Civil Engineering Transactions, CE29, 157–161. Nelson, R. C., 1996. Hydraulic roughness of coral reef platforms. Applied Ocean Research, 18, 265–274. Nelson, R. C., and Lesleighter, E. J., 1985. Breaker height attenuation over platform coral reefs. Preprints 1985 Australasian Conference on Coastal and Ocean Engineering, Christchurch, New Zealand. Conference Organising Committee, Christchurch, New Zealand, 2, pp. 9–16. Nielsen, P., 2009. Coastal and Estuarine Processes. Advanced Series on Ocean Engineering. Singapore: World Scientific, Vol. 29. Roberts, H. H., 1981. Physical processes and sediment flux through reef-lagoon systems. In Proceedings 17th International Coastal Engineering Conference, Sydney 1980. American Society of Civil Engineers, Vol. 1, pp. 946–962. Roberts, H. H., Murray, S. P., and Suhayda, J. N., 1975. Physical processes in a fringing reef system. Journal of Marine Research, 32, 233–260. Roberts, H. H., Wilson, P. A., and Lugo-Fernandez, A., 1992. Biologic and geologic responses to physical processes: examples from modern reef systems of the Caribbean-Caribbean Reefs region. Continental Shelf Research, 12, 809–834. Samosorn, B., and Woodroffe, C. D., 2008. Nearshore wave environments around a sandy cay on a reef platform. Continental Shelf Research, 28, 2257–2274. Sheppard, C., Dixon, D. J., Gourlay, M., Sheppard, A., and Payet, R., 2005. Coral mortality increases wave energy reaching shores protected by reef flats: examples from the Seychelles. Estuarine, Coastal and Shelf Science, 64, 223–234. Symonds, G., Black, K. P., and Young, I. R., 1995. Wave-driven flow over shallow reefs. Journal of Geophysical Research, Series C, 100, 2639–2648. Tartainville, B., and Rancher, J., 2000. Wave-induced flow over Mururoa Atoll Reef. Journal of Coastal Research, 16, 776–781. Wiegel, R. L., 1990. Transformation of a swell across a reef. Shore and Beach, 58(2), 31. Young, I. R., 1989. Wave transformation over coral reefs. Journal of Geophysical Research, 94(C7), 9779–9789.

Cross-references Cay Formation Engineering On Coral Reefs With Emphasis On Pacific Reefs Fringing Reef Circulation Hydrodynamics of Coral Reef Systems

WEST INDIAN CORAL REEF CLASSIFICATION

Infrastructure and Reef Islands Lagoon Circulation Reef Front Wave Energy Spurs and Grooves Tidal Effects on Coral Reefs Tropical Cyclone/Hurricane Tsunami Wave Set-Up Wave Shoaling and Refraction

WEST INDIAN CORAL REEF CLASSIFICATION Jörn Geister Naturhistorisches Museum Bern, Bern, Switzerland

Definition Coral reef classification: Assignments of single coral reefs or of parts thereof to reef types distinguished by their zonal structure. The present reef classification is ecological and based on the occurrence and distribution of certain coral associations (“facies zones”) that develop as a result of differential exposure to waves. Introduction As a side result of his longtime taxonomic studies in the northern Red Sea, Benjamin Klunzinger (1870) described for the first time in great detail the zonal distribution of reef biota parallel to the front of a fringing reef. Almost 85 years later, a well-developed zonal pattern was reported again by Wells (1954) from Bikini atoll in the western Indo-Pacific Ocean. Wells (1954, 396) defined the “ecological zone” of a coral reef as “an area where local ecological differences are reflected in the species association and signalized by one or more dominant species”. Thus, distribution of reef builders is essentially controlled by “ecological differences” (primary environmental factors) such as substratum (soft or hard), water depth (diminution of luminosity with increasing depth), wave exposure and intensity of seasonal abrasion. All of these fore-mentioned factors are entirely physical in nature (Geister, 1975, 1980a). From Jamaican reefs, Goreau (1959) presented a first detailed description of coral zonation in the western Caribbean Reefs region. Studying southwestern Caribbean reefs, Geister (1975) analyzed zonal patterns of the reef complex surrounding San Andrés Island, the reef zones being interpreted as facies units. The nature and causes of this and other benthic zonations, found both in Caribbean Reefs and Indo-Pacific reefs, were reviewed by Done (1983). Composite structure of shallow Caribbean coral reef types The monograph of Geister (1975) provided a first attempt of an ecological classification of shallow West Indian reefs

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based on the eye-catching zonal pattern of the framebuilding biota (facies zones) corresponding to graduated exposure to waves. Waves in the West Indies are essentially generated by the trade winds blowing with rather constant force from easterly directions. When incoming waves approach over the fore-reef slope of coral reefs, a gradual increase of bottom water movements is observed as an effect of shallowing. Maximum water movement is observed in the reef crest area where waves normally break (“breaker zone”). Decreasing hydrodynamic energy on the bottom behind the topographic crest is due to energy loss by wave friction and by the effect of gradual deepening. Graduated wave exposure on ridge-shaped coral reefs is mirrored by striking sub-parallel facies belts of the reef building biota (Figure 1). These are both ecological and facies zones defined by dense growths of predominant hermatypic corals, mixed coral associations or other reef biota adapted to different degrees of wave exposure. The wave zones recognized in Caribbean reefs are listed below, following a sequence of decreasing wave exposure: (1) Melobesieae wave zone (“algal ridge”) Red algae forming massive algal encrustations (2) Palythoa-Millepora wave zone Millepora spp. framework associated with the colonial zoanthid Palythoa sp. (3) strigosa-palmata wave zone Diploria strigosa heads with Acropora palmata thickets (4) cervicornis wave zone Acropora cervicornis thickets, almost monospecific (5) porites wave zone Porites porites thickets, almost monospecific (6) annularis wave zone Montastrea annularis (and related species M. franksi and M. faveolata) massive heads in mixed faunal complex including some branching species. Where well-developed, an overall symmetrical arrangement of subparallel “wave zones” will be recognized from the fore-reef (“front wave zones”) to the back-reef area (“rear wave zones”). This composite zonal structure is best developed in crest-shaped subtidal reefs facing the incoming waves (Figure 1). It permits to define six basic reef types. Each reef type is recognized by and named after the wave zone community on the reef crest (“breaker zone community”). This breaker zone community reflects the maximum wave exposure characteristic for each reef. We distinguish the following six basic “reef types” (Figure 1) arranged according to decreasing wave exposure: Melobesieae reef > Palythoa-Millepora reef > strigosa-palmata reef > cervicornis reef > porites reef > annularis reef Sequences of wave zones in front and to the rear of the breaker zone (reef crest) reflect successively more sheltered conditions on the sea bottom. If well developed, their

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WEST INDIAN CORAL REEF CLASSIFICATION

Palythoa-Millepora Reef strigosa-palmata Reef

strigosapalmata annularis

porites

annularis

BREAKER Z. PalythoaMillepora

Melobesieae

cervicornis

porites

cervicornis

strigosapalmata

FRONT WAVE ZONES

PalythoaMillepora strigosapalmata

porites annularis

Palythoa-Millepora

cervicornis cervicornis annularis

cervicornis

strigosapalmata porites

strigosa-palmata

annularis

porites cervicornis porites porites

annularis

cervicornis

cervicornis Reef

annularis

porites annularis porites annularis

porites Reef

annularis

annularis Reef

Increasing Exposure of Reefs to Wave Action

Melobesieae Reef

REAR WAVE ZONES

annularis

West Indian Coral Reef Classification, Figure 1 Idealized complete zonal sequences of the six basic reef types present in West Indian reef complexes. Increase in relative degree of wave exposure is indicated by simple to multiple arrows. The uppermost profile represents a reef exposed to maximum wave energy, the lowest to a minimum. The scheme shows a successive dropping out of high-energy associations with decreasing wave exposure and a shift of deeper water associations to shallower locations on sheltered reefs. Zonal sequences are normally incomplete, commonly with zonation gaps in the fore-reef and back-reef areas. Modified from Geister (1983).

zonal sequence may be rather complete as shown in Figure 1. But more frequently, parts of the front and rear wave zones are missing leaving empty zonal gaps. Incomplete or irregular series of wave zones result from irregular bottom topography in front or rear of the crest, from steep break-offs, unusually shallow reef flats, substrate changes from reef to lagoon facies (unsuitable substratum of loose sediments for settling corals!) and from rocky surfaces exposed to extreme seasonal storm waves causing abrasion (see Geister, 1977: Figure 1). The distribution of these six basic reef types within the reef complex of San Andrés Island was presented in a map (Geister, 1975: Figure 23). Comparable studies in the Indian Ocean (Rosen, 1975) suggest that the same principles of wave zonation are effective both in Caribbean Reefs and Indo-Pacific coral reefs. Exact comparison, however, is necessarily bound to reefs of the same faunal realm.

Regional and local distribution of reef types Reef classification based on ecological zonation facilitates comparative studies of regional reef distribution and reef mapping at scales ranging from single small patch reefs up

to large reef complexes. Thus, a follow-up study (Geister, 1977a) analyzed the geographic range of these reef types over the entire West Indian faunal province. Regional distribution of reef types mainly concerns seaward reefs around oceanic islands, atolls, and in front of continental coasts but also fringing and patch reefs within the lagoons of barrier reefs and atolls (Geister, 1975, 1977a, 1983, 1992; Geister and Diaz, 1997; Diaz et al., 2000; Geister and Diaz, 2007). Reefs of extreme wave exposure (Melobesieae reefs) are regionally restricted to areas facing a maximum windward fetch of the trade winds reaching up to 2,000 or more kilometers in length. Thus, these reefs are best developed along the eastern shores of the Lesser Antilles (Adey and Burke, 1976), where they receive the full impact of the Caribbean Reefs swell. Melobesieae reefs at the western ends of the Caribbean Sea and Gulf of Mexico are restricted to particular local topographic conditions where wave refraction and interference enhance local wave intensity (San Andrés Island, Providencia, San Blas/Panama, etc.). Reefs of high wave exposure are those of the Palythoa-Millepora type. They are best developed in the unprotected windward tracts of the central Caribbean (Pedro Bank etc.) and in the atolls and island barrier

WEST INDIAN CORAL REEF CLASSIFICATION

reefs along the Nicaraguan Rise and western Gulf of Mexico where the fetch of the trade winds attains around 500–2,000 km in length. Seaward reefs of the strigosa-palmata type reflect medium wave exposure. They are found protected from the full impact of oceanic swell by a wide shelf area or land mass as in Jamaica (Goreau, 1959), or lie in leeward positions facing a continental coast (Caribbean coast of Colombia: Diaz et al., 2000). They also flourish in the wave shade of higher wave exposure reef complexes as at Belize or in the lagoons of barrier reefs and atolls. Reef types reflecting moderate (Acropora cervicornis), low (Porites porites) and minimum wave exposure (Montastrea annularis) are mainly bound to protected lagoonal environments of major reef complexes. In addition, they flourish in the quieter waters of the lagoon floor at several meters of water depth (see Geister, 1975: Figure 23; Geister, 1992; Wallace, and Schafersman, 1977; Diaz et al., 1997). Where protected from oceanic swells, they may also characterize wellsheltered leeward sectors of atoll reefs, as at Belize. Local variation in distribution patterns of reef types can best be observed within the lagoons of single atoll or barrier reef complexes (Geister, 1975; Wallace, and Schafersman, 1977; Diaz et al., 1997). There, patch reefs will be found distributed corresponding to the local wave exposure pattern (see Geister, 1975: Figures 6 and 23).

Lateral transitions along a single reef tract from one to another reef type Along a same seaward reef tract, wave exposure will decrease from exposed windward towards more protected leeward positions. It will also decrease at locations where the reef crest is submerged for several meters, or where diffraction of swell occurs. Hence, single seaward reef tracts of barriers or atolls may reveal zonation patterns of more than one reef type as here defined. Vertical transitions in time from one to another reef type and sea-level history The three-dimensional zonal structure of modern reefs records Holocene sea-level history (Geister, 1983: Figures 49–54). This is exemplified by a standard reef composed of the following three wave zones: M. annularis, A. cervicornis and A. palmata that developed under conditions of stable, slowly rising, rapidly rising and falling sea level. Intermediate situations are common: Horizontal zonation pattern (Type A) As soon as a fairly stable sea level position was achieved for some extended period after the Holocene transgression, reef growth resulted in an essentially horizontal pattern of internal zonation: lateral growth widened the established facies zones. Growing vertically into higher energy levels, they were succeeded by the wave zone corresponding to the next higher energy level. Thus, vertical transitions from an annularis Reef, to a cervicornis

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Reef and finally to a strigosa-palmata Reef are possible (Geister, 1983: Figure 49, 50).

Vertical zonation pattern (Type B) As soon as renewed sea level rise equaled the rate of reef growth, a pattern of vertical parallel facies zones developed from a Type A reef into an overall U-shaped zonal structure. The facies zone of highest wave exposure is found in the center, and successively lower wave exposure facies will grow towards the outer margins of the reef body (Geister, 1983: Figure 51). This geometry corresponds to the “keep-up reef ” of Neumann and Macintyre (1985). Lens-shaped zonation pattern (Type C) When sea level rise eventually surpassed the growth rate of the Type A reef, the latter was gradually enveloped by facies zones of successively lower hydrodynamic energy, thus resembling an onion peel structure (Geister, 1983: Figure 53). Characteristic for this reef type is an upward sequence of facies zones, which reflects decreasing wave exposure and final reef drowning (Blanchon and Shaw, 1995). Breaker zones are covered by facies of successively lower wave exposure. This geometry corresponds to the “give-up reef ” of Neumann and Macintyre (1985). Cap-on-cap zonation pattern (Type D) The internal facies geometry of a reef growing during a period of relative sea level lowering is characterized by a cap-on-cap arrangement of facies zones indicating downward increase of wave exposure with time. It may eventually lead to the emergence of the crest as a rocky reef islet. The best modern examples from the Caribbean Sea are the emerging reefs of the Rosario Islands chain and neighboring Baru Peninsula in Colombia (Geister, 1983: Figure 54; Pl. 26/4). Occurrence and distribution of reef types in the Pleistocene of the West Indies Of the six reef types defined herein for the modern Caribbean Sea, five were found in Late Pleistocene (mostly Sangamonian) reefs of the West Indies. Thus, modern reef types may also serve as models for comparative palaeoecological studies and for mapping of Pleistocene reef tracts (Geister, 1975: Figure 27; Geister, 1980b). However, during a preliminary survey comprising the emergent reefs of the whole West Indian faunal province, no reefs corresponding to the Palythoa-Millepora type were encountered (Geister, 1980b). The reason may be that the fire coral Millepora was rare at that time and locally absent from Sangamonian reefs. On the other hand, the annularis to porites wave zones in Sangamonian backreefs were locally dominated by extensive thickets of Pocillopora sp., a branching reef coral genus completely absent from modern West Indian reefs (Geister, 1977b, 1984: Figures 13 and 14).

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Bibliography Adey, W. H., and Burke, R., 1976. Holocene bioherms (algal ridges and bank-barrier reefs) of the eastern Caribbean. Geological Society of America Bulletin, 87, 95–105. Blanchon, P., and Shaw, J., 1995. Reef drowning during the last deglaciation: Evidence for catastrophic sea-level rise and icesheet collapse. Geology, 23, 4–8. Diaz, J. M., Barrios, L. M., Cendales, M. H., Garzon-Ferreira, J., Geister, J., Lopez-Victoria, M., Ospina, G. H., Parra-Velandia, F., Pinzon, J., Vargas-Angel, B., Zapata, F. A., and Zea, S., 2000. Areas coralinas de Colombia, Instituto de Investigaciones Marinas y Costeras “José Benito Vives de Andreis” INVEMAR. Sta. Marta, Colombia: Serie Publicaciones Especiales, Vol. 5, 1–175. Diaz, J. M., Sanchez, J. A., and Geister, J., 1997. Development of lagoonal reefs in oceanic reef complexes of the southwestern Caribbean: geomorphology, structure and distribution. In Proceedings 8th international Coral Reef Symposium, Vol. 1, pp. 779–784. Done, T. J., 1983. Coral Zonation: Its Nature and Significance. In: Barnes, D. J., (ed.), Perspectives on Coral Reefs. Brian Clouston Publisher, pp. 107–147. Geister, J., 1975. Riffbau und geologische Entwicklungsgeschichte der Insel San Andrés (westliches Karibisches Meer, Kolumbien). Stuttgarter Beiträge zur Naturkunde, Serie B (Geologie & Paläontologie), 15, 1–203. Geister, J., 1977a. The influence of wave exposure on the ecological zonation of Caribbean coral reefs. In Proceedings, Third International Coral Reef Symposium, University of Miami, Vol. 1, pp. 23–29. Geister, J., 1977b. Occurrence of Pocillopora in Late Pleistocene Caribbean coral reefs. Mémoíres du Bureau de Recherches géologiques et miniéres, 89, 378–388. Geister, J., 1980a. Morphologie et distribution des coraux dans les récifs actuels de la mer des Caraïbes. Annali dell’Università di Ferrara (Nuova Serie), Sez. IX Scienze Geologiche e Paleontologiche. VI(supplemento), 15–28. Geister, J., 1980b. Calm-water reefs and rough-water reefs of the Caribbean Pleistocene. Acta palaeontologica polonica, 25, 541–556. Geister, J., 1983. Holocene West Indian coral reefs: geomorphology, ecology and facies. Facies, 9, 173–284. Geister, J., 1984. Récifs pléistocènes de la mer des Caraïbes: aspects géologiques et paléontologiques. In: Geister, J., and Herb, R., (eds.), Géologie et paléoécologie des récifs. Berne: Institut de Géologie de l’Université de Berne, pp. 3.1–3.34. Geister, J., 1992. Modern reef development and Cenozoic evolution of an oceanic island/reef complex: Isla de Providencia (Western Caribbean Sea, Colombia). Facies, 27, 1–70. Geister, J., and Diaz, J. M., 1997. Field guide to the oceanic barrier reefs and atolls of the southwestern Caribbean (Archipelago of San Andrés and Providencia, Colombia). In Proceedings 8th International Coral Reef Symposium, Vol. 1, pp. 235– 262. Geister, J., and Diaz, J. M., 2007. Reef environments and geology of an oceanic archipelago: San Andrés, Old Providence and Sta. Catalina Islands. Caribbean Sea, Colombia. Bogota: INGEOMINAS. Goreau, Th. F., 1959. The ecology of Jamaican coral reefs. I. Species composition and zonation. Ecology, 40, 67–90. Klunzinger, C. B., 1870. Eine zoologische Excursion auf ein Korallriff des rothen Meeres. Verhandlungen zoologischbotanische Gesellschaft Wien, 20, 389–394. Neumann, A. C., and Macintyre, I., 1985. Reef response to sea level rise: Keep-up, catch-up or give-up. In Proceedings Fifth international Coral Reef Congress, Tahiti, Vol. 3, pp. 105–110.

Rosen, B. R., 1975. The distribution of reef corals. Reports of the Underwater Association, 1(N.S.), 2–16. Wallace, R. J., and Schafersman, S. D., 1977. Patch-reef ecology and sedimentology of Glovers Reef Atoll, Belize. In Reefs and related carbonates – ecology and sedimentology. American Association of Petroleum Geologists, Studies in Geology Vol. 4, pp. 37–52. Wells, J. W., 1954. Recent corals of the Marshall Islands. United States Geological Survey Professional Paper, 260-I, 285–486.

Cross-references Bahamas Bermuda Eastern Caribbean Coral Reefs Florida Keys Geomorphic Zonation Ecomorphology Lagoons Patch Reefs: Lidar Morphometric Analysis Wave Set-Up Wave Shoaling and Refraction Western Atlantic/Caribbean, Coral Reefs

WESTERN ATLANTIC/CARIBBEAN, CORAL REEFS Bernhard Riegl Nova Southeastern University, Dania, FL, USA

Definition and introduction Scleractinian corals occur all throughout the western Atlantic; however, reefs are only formed in a belt between 32.3 N (Bermuda) and 17.5 S (Brazil). The continuous reef belt occurs between 26.5 N (Florida) and 9.8 N (Colombia) with an isolated southern extension in Brazil, separated from the Caribbean reef belt by the Orinoco and Amazonas rivers, and Bermuda as an outlier to the north. The western Atlantic forms its own biogeographic province and harbors a unique coral fauna that diversified primarily after the closure of the Isthmus of Panama, but the evolution of which may have already begun in the Eocene. All reef types known from other reef areas also occur in the western Atlantic, but some reef types are unique. The coral fauna The Caribbean/western Atlantic region is home to a unique coral reef fauna and therefore makes up a distinct biogeographic province. Within this province subtle difference further subdivides it into a Bermudian (lacking Acropora), northern (E-Florida/Bahamas), north-western (Gulf of Mexico, W-Florida), central (Mesoamerica, Antilles), and Brazilian subregion (Veron, 1995). The most distinct region is the Brazilian with several endemic species and an endemic genus (Mussismilia braziliensis, M. hispida, M. hartii, Favia gravida, F. leptophylla, Siderastrea stellata, Millepora braziliensis, M. nitida, M. laboreli; Leao et al., 2003; Amaral et al., 2008). The Brazilian fauna has

WESTERN ATLANTIC/CARIBBEAN, CORAL REEFS

affinities to the Tethyan and Paratethyan Miocene faunas, where these genera were common (Leao and Kikuchi, 2005). Within the Atlantic Ocean, only the western region supports coral reefs. The easternmost Atlantic harbors corals, but no true reef building, although some dense coral growth is reported from the Gulf of Guinea and the Cape Verde Islands (Laborel, 1974; Moses et al., 2003). The Mediterranean, up to the Miocene a locus of vigorous coral growth, has been essentially devoid of any reefal development since the Messinian Salinity Crisis, during which the Mediterranean desiccated or at least became hypersaline (Brachert et al., 1996). It later refilled with cryospheric Atlantic middle water, thus displacing its warm water fauna with one of cold water, that has never been fully displaced even though the Mediterranean has warmed since. In the Plio and Pleistocene, and in some places also in the Holocene, Cladocora caespitosa formed sizeable (meters thick) banks (Aguirre and Jimenez, 1998). Algal-built hardgrounds (the coralligène, Bosence, 1985a) and Cladocora banks presently occur, but are at best analogues to more ancient, but not equivalent to recent W-Atlantic reefs. The end-Miocene, early Pliocene faunal crisis in the eastern Atlantic left an isolated but highly diverse fauna with mainly Indo-Pacific species in the western Atlantic that was fully isolated with the formation of the Panamanian land bridge in the lower Pliocene (3.5 Ma = million years ago; Collins et al., 1996). During the period of 6–1 Ma, the Indo-Pacific fauna was progressively reduced by a relatively slow step-down extinction and replaced by pulses of modern Caribbean fauna. Seventeen of 41 genera living in the Caribbean in the Pliocene became extinct, 11 of these genera continued to exist in the Indo-Pacific. No new genera originated and, with the exception of Pocillopora which disappeared again, no species immigrated (Budd and Wallace, 2008). Prior to this turnover, shallow areas were dominated by Stylophora and Pocillopora. One of the most iconic species of the Caribbean coral fauna, Acropora palmata, is first reported from Costa Rica (Qebrada Chocolate Formation) about 3.6–2.6 Ma (McNeill et al., 1997). A. palmata was rare at the beginning of its rise. A. cervicornis, however, commonly cooccurred with Stylophora in several late Pliocene formations (e.g., in Costa Rica, Jamaica, and Curaçao; Budd and Wallace, 2008). Several other Acropora (two of the subgenus Isopora which presently only occurs in the Indo-Pacific) had existed in the Caribbean but had all disappeared during the last extinction pulse around 2 Ma. Molecular analyses have shown that several traditional families of coral may be polyphyletic (e.g., Faviidae, Mussidae), and may be even many genera (e.g., Montastraea, Favia, Scolymia). A distinct clade in Atlantic members of the Faviidae and Mussidae suggests that the relationship between Atlantic and Indo-Pacific corals may indeed not be very close and divergence could predate final closure of the Central American Isthmus in the Pliocene to as early as the Eocene. Nine of the 27

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genera of reef-building Atlantic corals belong to this previously unrecognized lineage, which probably diverged over 34 Ma (Fukami et al., 2004) (Figure 1).

Loss of long-term ecological stability Remarkable ecological stability has been observed in Caribbean coral assemblages throughout the Pleistocene and Caribbean coral reefs exhibit a clear zonation pattern throughout the region (Geister, 1983). It has been shown that coral community structure on Caribbean coral reefs reassembled after global sea-level changes in similar ways throughout a 500-ka interval (Pandolfi and Jackson, 2006). Six common coral species (A. palmata, A. cervicornis, three Montastraea, and Diploria strigosa) dominated the reefbuilding episodes. While not unique to the Caribbean, this is the longest demonstrated such interval of community stability. Similarly, Pleistocene community recurrence patterns were found in New Guinea (Huon Peninsula) over 115 ka (Pandolfi, 1996). In the Holocene, community stability has been demonstrated in the Caribbean (Aronson et al., 2004) and Brazil (Leao and Kikuchi, 2005). However, in stark contradiction to the situation in the Pacific, this persistence of community structure has recently been lost throughout most of the Caribbean by the almost regionwide demise of the A. palmata zone and equally dramatic reduction of A. cervicornis (Pandolfi and Jackson, 2006). Also in Brazil, community patterns and coral density that have shown relative stability on millennial scales have recently changed (Leao and Kikuchi, 2005). Types of reefs and frameworks The W-Atlantic/Caribbean region exhibits reef growth in all the classic carbonate sedimentological settings (Schlager, 2005; Hine et al., 2008): ramp (W-Florida, NE Yucatan), unrimmed shelf (Yucatan, Brazil), attached rimmed shelf (SE Florida and Florida Keys, Cuba, Belize), and rimmed shelves on unattached (or isolated) platforms (Bahamas banks, Chinchorro, Belize banks, etc.). A specialty of Atlantic rimmed shelves is coral reefs often cooccurring and alternating with grainstone shoals (frequently ooids, i.e., Bahamas and Yucatan). In the shallowest regions of the shelf, typical fringing reefs (e.g., around St. Croix, Vieques, Lesser Antilles, Cayman Islands, Jamaica, Roatan, etc.) or bank–barrier reefs (such as developed in the Florida Keys or the SE Puerto Rico shelf) occur. Many of these fringing reefs are built by A. palmata frameworks in situ, although some are not and consist mainly of rubble (see below). Many such structures initiated about 8–6 ka and either continued growing until today or, as in the high-latitude SE Florida continental reef tract, experienced a switch off around 6–4 ka. On the middle shelf, mid-shelf reefs, situated at a depth level intermediate between shelf-edge and fringing reefs, are reported from many Caribbean localities (Puerto Rico, Vieques, Florida, Cayman, etc.). Blanchon et al. (1997) explained mid-shelf reefs at Cayman as a result of stepped Holocene sea-level change, i.e., periods

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Western Atlantic/Caribbean, Coral Reefs, Figure 1 (a) Acropora palmata, an iconic species for the Caribbean coral fauna (here at Lee Stocking island, Bahamas), has recently experienced significant die-back throughout its range (b) long-term ecological stability throughout the Pleistocene (outcrop in Curacao), (c) stromatolites can form small fringing reefs in the Bahamas or subtidal structures, such as these at Lee Stocking island, Bahamas. Measuring stick = 1 m. The stromatolite with the stick shows artichoke-like leaves, (d) rubble is an important constituent of Caribbean reef frameworks and now all that is left of the former A. palmata zone throughout much of its range, (e) islands on the northern Caribbean banks do not have fringing reefs. Shown is Lee Stocking Island on the Great Bahama Bank, the only area of shallow A. palmata reef is shown in a circle, the area of subtidal stromatolites within a dotted polygon, and the trend of the mid-shelf reef (at about 10 m depth) with a dashed line. (f) The western end of Roatan (Bay Islands, Honduras) with a typical fringing reef and three offshore bank reefs to the south.

of unusually rapid and then slowing sea-level rise. The realism of this sea-level assumption (as that of many others) is, however, being debated (Hubbard et al., 2008; Blanchon et al., 2009) and mid-shelf structures can be found at variable depth and with variable morphology (Banks et al., 2007; Hubbard et al., 2008), making their explanation purely by sea-level difficult since antecedent topography may have also played a role. Shelf-edge reefs are a typical feature throughout the Caribbean and have been described in detail from Barbados (Macintyre, 1988), Florida (Lighty et al., 1978, Banks et al., 2007), St. Croix, the Lang Bank, and Puerto Rico (Hubbard et al., 1990, 2008). They are typically built by A. palmata, are situated at depths around 20 m, initiated around

12–10 ka, and flourished until 8 ka. Tracking rising sea level, they then stepped back throughout the Caribbean to form the wide variety of shallower shelf reefs. In Belize and Mexico, the Caribbean also has one of the world’s biggest barrier reefs, a largely structurally controlled feature where reef development is guided and determined by fault lines (Purdy et al., 2003), lowstand karst topography (Purdy, 1974), and paleo-river beds (Esker et al., 1998). Other large barrier reefs exist in the Bahamas at Andros, in the Turks and Caicos, Venezuelan off-shore banks, etc. Deep, submerged bank reefs occur in the Gulf of Mexico (half a dozen banks of which the Flowergarden banks are the most famous, Schmahl et al., 2008), where they are associated with salt diapirism, on the Nicaragua rise, and

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in Brazil (Leao et al., 2003), where they often consist of bedrock (older limestones or crystalline basement). Unusual, bank-like, deep reefs also occur off SW-Florida, such as the carbonate banks of the Florida Middle Grounds and a drowned dune/coastal spit complex at Pulley ridge (Jarrett et al., 2005). These deep banks are home to unique, and frequently dense, coral assemblages (Hine et al., 2008). Several reef types unique to specific Atlantic region exist. Outlier reefs in the Florida Keys Reef Tract (Lidz et al., 2008; Hine et al., 2008) are Pleistocene reefs situated on an upper-slope terrace, separated from the present shelf edge with its shelf-edge reefs by a well-developed trough. Outlier reefs form a discontinuous and drowned (surfaces at around 10 m below MSL), largely parallel tract to the major present reef trend on the shelf edge. The outliers comprise a stacked sequence of reefs from MIS 5c at the base to 5b and 5a at the contact to Holocene overgrowth and Lidz et al. (2008) consider them testament to optimal platform margin settings for reef growth during 5c and later times. Algal cup reefs occur in Bermuda. They are somewhat reminiscent of the Brazilian Chapeiroes (mentioned below) inasmuch as the cup reefs can also take a mushroom-shaped appearance. The tops are made by calcareous algae and can have a raised rim and a slightly deeper interior. They are mostly built by calcareous algae, the gastropod Dendropoma irregulare and Millepora spp. (Ginsburg and Schroeder, 1973). Algal ridges occur throughout the Lesser Antilles that are built by several thousand years of algal growth (Adey and Burke, 1976) and off Cozumel, coralline algal microatolls occur (Boyd et al., 1963). Chapeiroes in Brazil are pinnacle reefs made of Holocene reef limestone that consist of isolated narrow pillars that expand on the top to take a mushroomshaped appearance (Leao et al., 2003). These reefs can be 5–25 m high and have diameters of 5–50 m at their tops and are mainly built by coral. This type of reef is restricted to the Abrolhos reef complex in southern Bahia State, where a wide variety of morphologies of isolated and coalescent Chapeiroes exist. On many other Brazilian reefs, most particularly close inshore, substratum for coral growth is frequently on the former beachrock, indurated dunes, or bedrock (igneous or carbonate). This situation is reminiscent of high-latitude reefs in the Indo-Pacific, with corals in SE Africa and SE Australia carpeting indurated sand dunes and beachrock ridges. Stromatolites are microbialites that are common and attain large sizes along the Exuma margin in the Bahamas. Modern stromatolites, which are layered sedimentary rocks formed by sediment-trapping microbial mats, are rare and best documented from W-Australia (Shark Bay) and the Bahamas (Reid et al., 1995). The Bahamian stromatolites can be domal, molar shaped, or elongate and have some leaf-shaped structures that are reminiscent of Precambrian stromatolites. At Stocking Island and Highbourne Cay in the Bahamas, regular small fringing reefs are dominated in their backreef areas by stromatolites and toward the beach by thrombolites (Macintyre et al., 1996; Reid et al., 1999; Andres et al., 2009). These stromatolites are 5–6 m/ka) until about 8–6 ka, when it slowed to about 1 m/ka while ascending to today’s level (Toscano and Macintyre, 2003; Hubbard et al., 2008). Throughout the Pacific, there is clear evidence of an earlyto mid-Holocene temporary sea-level highstand of about 3 m, centered at 4 ka (the Brazilian record suggests a maximum of nearly þ5 m at 5 ka, with subsequent decline to today’s level; Leao et al., 2003). The overshoot (generally ascribed to earth rheology) is documented in raised reefs, raised intertidal notches, or raised intertidal benches throughout the Pacific basin. The Caribbean region does not display any unequivocal such features, but Brock et al. (2008) used morphological features of northern Florida Keys patch reefs in combination with raised beach ridges to infer such an overshoot also for the Caribbean region.

Conclusion Coral reefs in the Atlantic/Caribbean differ from those in the Indo-Pacific with regard to fauna as well as aspects of reef framework formation. The Atlantic coral fauna began its evolution prior to the closing of the isthmus of Panama, but fully diversified only afterwards. Remarkable ecological continuity of Pleistocene coral community organization has recently been lost. Several types of reefs unique to the Atlantic/Caribbean exist (stromatolitic fringing reefs, Chapeiroes, algal cup reefs, outlier reefs).

Fringing reef and island development on the largest

Caribbean banks (Bahamas banks, Florida Keys) are frequently different than in the Indo-Pacific due to relatively high-latitude position, but more comparable within the truly tropical belt (Belize, Venezuela, etc.). No Holocene raised reefs known from Caribbean region, only in Brazil signs of early Holocene highstand.

Bibliography Adey, W. H., and Burke, R., 1976. Holocene bioherms (algal ridges and bank–barrier reefs) of the eastern Caribbean. Geological Society of America Bulletin, 87, 95–109. Aguirre, J., and Jimenez, A. P., 1998. Fossil analogues of present-day Cladocora caespitosa coral banks: sedimentary setting, dwelling community, and taphonomy (Late Pliocene, W Mediterranean). Coral Reefs, 17, 203–213. Amaral, F., Steiner, A. Q., Broadhurst, M. K., and Cairns, S. D., 2008. An overview of the shallow-water calcified hydroids from Brazil (Hydrozoa: Cnidaria), including the description of a new species. Zootaxa, 1930, 56–68. Andres, M. S., Reid, R. P., Bowlin, E., Gaspar, A. P., and Eisenhauer, A., 2009. Microbes versus metazoans as dominant reef builders: insights from modern marine environments in the Exuma Cays, Bahamas. International Association of Sedimentologists Special Publication, 41, 149–165. Aronson, R. B., Macintyre, I. G., Wapnick, C. M., and O’Neill, M. W., 2004. Phase shifts, alternative states, and the unprecedented convergence of two reef systems. Ecology, 85, 1876–1891. Banks, K. W., Riegl, B., Shinn, E. A., Piller, W. E., and Dodge, R. E., 2007. Geomorphology of the southeast Florida continental reef tract (Miami-Dade, Broward, and Palm Beach Counties, USA). Coral Reefs, 26, 617–640. Blanchon, P., Jones, P., and Kalbfleisch, W., 1997. Anatomy of a fringing reef around Grand Cayman: storm rubble, not coral framework. Journal of Sedimentary Research, 67, 1–16. Blanchon, P., Eisenhauer, A., Fietzke, J., and Liebetrau, V., 2009. Rapid sealevel rise and reef back-stepping at the close of the last interglacial highstand. Nature, 458, 881–885. Bosence, D., 1985a. The Coralligene of the Mediterranean – a recent analogue for tertiary coralline algal limestones. In Toomey, D. F., Nitecki, M. H. (eds.), Paleoalgology. Berlin: Springer, pp. 215–225. Bosence, D., 1985b. The morphology and ecology of a moundbuilding coralline alga (Neogoniolithon strictum) from the Florida Keys. Paleontology, 28, 189–206. Boyd, D. W., Kornicker, L. S., and Rezak, R., 1963. Coralline algae microatolls near Cozumel Island. Mexico: University of Wyoming Contributions to Geology, 2, 105–108. Brachert, T., Betzler, C., Braga, J. C., and Martin, J. M., 1996. Record of climatic change in neritic carbonates: turnover in biogenic associations and depositional modes (Late Miocene, southern Spain). International Journal of Earth Sciences, 85, 327–337. Brock, J. C., Palaseanu-Lovejoy, M., Wright, C. W., and Nayegandhi, A., 2008. Patch-reef morphology as a proxy for Holocene sea-level variability, northern Florida Keys, USA. Coral Reefs, 27, 555–568. Budd, A. F., and Wallace, C. C., 2008. First record of the IndoPacific reef coral genus Isopora in the Caribbean region: two new species from the Neogene of Curacao, Netherlands Antilles. Paleontology, 51, 1387–1401. Collins, L. S., Budd, A. F., and Coates, A. G., 1996. Earliest evolution associated with closure of the Tropical American Seaway.

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coral reefs of the Florida Keys. In Riegl, B., and Dodge, R. E. (eds.), Coral Reefs of the USA. Berlin: Springer, pp. 9–74. Lighty, R. G., Macintyre, I. G., and Stuckenrath, R., 1978. Submerged early Holocene barrier reef south-east Florida shelf. Nature, 275, 59–60. Macintyre, I. G., 1988. Modern coral reefs of western Atlantic: new geologic perspectives. American Association of Petroleum Geologists Bulletin, 72, 1360–1369. Macintyre, I. G., Reid, R. P., and Steneck, R. S., 1996. Growth history of stromatolites in a Holocene fringing reef, Stocking Island, Bahamas. Journal of Sedimentary Research, 66, 231–242. McNeill, D. F., Budd, A. F., and Borne, P. F., 1997. An earlier (Late Pliocene) first appearance of the reef-building coral Acropora palmata: stratigraphic and evolutionary implications. Geology, 25, 891–894. Messing, C. G., Reed, J. K., Brooke, S. D., and Ross, S. W., 2008. Deep-water reefs of the United States. In Riegl, B., and Dodge, R. E. (eds.), Coral Reefs of the USA. Berlin: Springer, pp. 767–791. Moses, C. S., Helmle, K. P., Swart, P. K., Dodge, R. E., and Merino, S. E., 2003. Pavement of Siderastrea radians on Cape Verde reefs. Coral Reefs, 22, 506. Neumann, A. C., Kofoed, J. W., and Keller, G. H., 1977. Lithoherms in the straits of Florida. Geology, 5, 4–11. Pandolfi, J., 1996. Limited membership in Pleistocene reef coral assemblages from the Huon peninsula, Papua New Guinea: constancy during global change. Paleobiology, 22, 152–176. Pandolfi, J. M., and Jackson, J. B. C., 2006. Ecological persistence interrupted in Caribbean coral reefs. Ecology Letters, 9, 818–826. Playford, P. E., and Cockbain, A. E., 1976. Modern algal stromatolites at Hamelin Pool, a hyper-saline barred basin in Shark Bay, Western Australia. In Walter, M. R. (ed.), Stromatolites. Amsterdam: Elsevier, pp. 389–411. Purdy, E. G., 1974. Karst-determined facies patterns in British Honduras: Holocene carbonate sedimentation model. American Association of Petroleum Geologists Bulletin, 58, 825–855. Purdy, E. G., Gischler, E., and Lomando, A. J., 2003. The Belize margin revisited. 2. Origin of Holocene antecedent topography. International Journal of Earth Sciences, 92, 552–572. Rasser, M. W., and Riegl, B., 2002. Holocene coral reef rubble and its binding agents. Coral Reefs, 21, 57–72. Reid, R. P., MacIntyre, I. G., Browne, K. M., Steneck, R. S., and Miller, T., 1995. Modern marine stromatolites in the Exuma Cays, Bahamas: uncommly common. Facies, 33, 1–18. Reid, R. P., MacIntyre, I. G., and Steneck, R. S., 1999. A microbialite/algal ridge fringing reef complex, Highborne Cay, Bahamas. Atoll Research Bulletin, 466, 1–18. Roberts, H. H., 1992. Reefs, bioherms, and lithoherms of the northern Gulf of Mexico: the important role of hydrocarbon seeps. In Proceedings of the 7th International Coral Reef Symposium, Guam, Vol. 2, pp. 1121–1128. Roberts, H. H., Rouse, L. J. Jr., Walker, N. D., and Hudson, J. H., 1982. Cold water stress in Florida Bay and the northern Bahamas: a product of cold-air outbreaks. Journal of Sedimentary Petrology, 52, 145–155. Schlager, W., 2005. Carbonate sedimentology and sequence stratigraphy. SEPM Concepts in Sedimentology and Paleontology, 8, 198 p. Schmahl, G. P., Hickerson, E. L., and Precht, W. F., 2008. Biology and ecology of coral reefs and coral communities in the Flower Garden Banks region, northwestern Gulf of Mexico. In Riegl, B., and Dodge, R. E. (eds.), Coral reefs of the USA. Berlin: Springer, 221–261. Sprachta, S., Camoin, G. F., Golubic, S., and LeCampion, T., 2001. Microbialites in a modern lagoonal environment: nature and distribution, Tikehau atoll (French Polynesia). Paleogeography, Paleoclimatology, Paleoecology, 175, 103–124. Toscano, M. A., and Macintyre, I. G., 2003. Corrected western Atlantic sea-level curve for the last 11,000 years based on

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calibrated 14C dates from Acropora palmata framework and intertidal mangrove peat. Coral Reefs, 22, 257–270. Turmel, R., and Swanson, R., 1976. The development of Rodriguez Bank, a Holocene mudbank in the Florida Reef Tract. Journal of Sedimentary Petrology, 46, 497–519. Veron, J. E. N., 1995. Corals in Space and Time. Sydney: University of New South Wales Press, 321 p.

Cross-references Bahamas Banks Island: Frasnian (Late Devonian) Reefs In Northwestern Arctic Canada Barbados Belize Barrier and Atoll Reefs Bermuda Brazil, Coral Reefs Florida Keys Ooids Sediment Durability Tethys Ocean Vaughan, Thomas Wayland (1870–1952)

WESTERN AUSTRALIAN REEFS Lindsay B. Collins Curtin University of Technology, Bentley, WA

Synonyms Reefs; Western Australia Definition Wave resistant structures built primarily by coral growth. Introduction Situated in the eastern Indian Ocean along the western continental margin of Australia, coral reefs extend from the tropical Kimberley coast in the north to the temperate southwest (Figure 1). This latitudinal climatic gradient places reefs in a setting likely to become increasingly significant as a natural laboratory when predictions of global climate change are considered. The relatively small human population along several thousand kilometres of coast (44% of the Australian coastline) limits anthropogenic impacts from a largely desertic coastal hinterland bordering the continental shelf. Coral reefs include fringing reefs such as the Ningaloo Reef (at 280 km long, the world’s longest fringing reef ), mid-outer shelf reefs such as the Houtman Abrolhos reefs, the southernmost reefs in the Indian Ocean, and isolated reefs such as the Rowley Shoals and Scott Reef as some of the better known reefs, but there are many other poorly known and remote systems in the region. Regional oceanography Three oceanographic and biogeographic provinces (Figure 1) present along Australia’s west coast are: the Northern Australian Tropical Province (11–22 S), the

Western Coast Overlap Zone (22–32 S), and the Southern Australian Warm Temperate Province (south of 32 S). This regional oceanography is influenced by the South Equatorial Current, at 5–15 S latitude, driven by easterly trade winds, and the Indonesian Throughflow, which floods the North West Shelf with warm, low salinity water, resulting in sea levels in the tropics being 0.5 m higher than along the southern coast of Australia. With higher tropical sea levels, the formation of a north-south pressure gradient induces a weak easterly flow of central Indian Ocean subtropical water toward the Australian coast between 15–35 S (Pearce and Griffiths, 1991). This easterly flow is deflected south by the coastline, eventually contributing to the Leeuwin Current, a warm low salinity current which flows from the Indonesian Throughflow, southward along the adjacent shelf in winter. The Indonesian Throughflow delivers larvae of both Pacific and Asian reef species southward, and the Leeuwin Current is an important control on southward larval delivery, whilst suppressing upwelling (Hatcher, 1991). Tidal range and cyclone frequency decrease from north to south. The North West Shelf is cyclonic and tidally dominated, with a mean spring range from 9.2 m in King Sound to 1.7 m at Ningaloo Reef (Harris et al., 1991). Both Scott Reef and Rowley Shoals have semi-diurnal tides with a spring range of 4.5 m. At the southerly limits the coral growth reefs are microtidal (1 m) and swell wave dominated with low cyclone frequency (1 per 4 years).

Reef morphology Studies of Australian reefs have been dominated by work on the eastern seaboard and the Great Barrier Reef (see Hopley et al., 2007), but reefs of the western continental margin of Australia and its bordering carbonate ramps remain relatively unknown. Coral reefs of Australia’s western margin include isolated oceanic atoll-like reefs (Ashmore Reef, Seringapatam and Scott Reefs, Rowley Shoals); island-associated shelf reefs and fringing reefs of the Kimberley coast and Dampier Archipelago; Pilbara reefs (Barrow and Montebello Islands); Ningaloo Reef, a fringing reef adjacent to the North West Cape, and the mid- outer shelf Houtman Abrolhos carbonate platforms and reefs (Collins et al., 1993a, b, 1997, 2009). Mid-outer shelf reefs: the Houtman Abrolhos Between latitudes 28–29.5 S the open, low-gradient shelf is interrupted by a 150 km long, discontinuous rim, whose seaward margin is 8–10 km east of the shelf/slope break near the 100 m isobath. The emergent rim consists of three platforms (Pelsaert, Easter and Wallabi platforms) separated by channels up to 40 m deep. Each platform rises abruptly some 40 m above a flat shelf, and culminates as reef flats and low islands. Submerged banks lie to the north and south of the platforms along the same trend. The three platforms differ geomorphologically but a windward reef, leeward reef, and lagoon with a central platform is distinguishable in each case. Both the central

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Western Australian Reefs, Figure 1 Distribution of Western Australian Reefs. Note position of Northern Australian Tropical Province (white), Southern Australian Warm Temperate Province (dots) and Western Coast Overlap Zone.

platforms and leeward reefs are expressed as small islands of various types, whereas windward reefs are wave-swept. The central platforms are the last interglacial ones in age, whereas the windward and leeward reefs are Holocene (Eisenhauer et al., 1993; Collins et al., 1993a, b; Zhu et al., 1993). The islands generally rise only 3–5 m above sea level. Extensive ‘blue-hole’ terrains and reticulate reefs occur at the eastern parts of the platforms and Holocene sand sheets have developed in both the windward and leeward lagoons (Wyrwoll et al., 2006). Central platform islands are surfaced by dense, calcretized limestones, while subparallel ridges of coral rubble characterize leeward (eastern) islands. The Houtman Abrolhos region is located within the Western Coast Overlap Zone and biotic transition (Morgan and Wells, 1991). It is characterized by the gradual replacement of a tropical fauna in the north by a predominantly temperate fauna in the south, as reflected in a variety of shelf and reef biotic elements, including corals, macroalgae, molluscs, echinoderms and fishes. Coral faunal communities are highly diverse; 184 species and 42 genera are recorded

(Veron and Marsh, 1988). The number of genera is much higher than for other ‘high latitude’ coral reefs (e.g., the Solitary Islands; Kure Atoll; Bermuda Reefs). Acropora, which is missing or vary rare on many ‘high latitude’ reefs is the dominant coral. Although there are substantial numbers of temperate species and Western Australian endemic species the fauna at the Houtman Abrolhos is, on the balance, essentially tropical, and is generally considered to be at the southern limit in Western Australia of the tropical biota. Forty metres of Holocene reef buildup has occurred yielding high precision reef growth and sea level records for the Holocene and the last interglacial reefs (see Eisenhauer et al., 1993; Collins et al., 1993a, b).

Ningaloo fringing reef Situated close to North West Cape and Cape Range, this fringing reef complex consists of a narrow reef crest, which is emergent at low water, with well-developed spur and groove morphology present on most outer reef slopes; complex multiple developments of spur and groove are also present. The reef crest is backed by a reef flat

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(usually