narrow, less sediment accumulates and more coral development is possible. Conversely, where the terrace is wide,
a larger sand reservoir can prohibit spur development
completely.
The sand terrace is a prominent feature seen in imagery
from all oceans (Figure 6) and has been widely reported as
the “ten fathom” terrace (Emery, 1961; Garrison and
McMaster, 1966). Where it is not covered by excessive
reef development, the terrace has been documented in
roughly the same 20 Æ 5 m depth range from the following
areas: the Marshall Islands (Tracy et al., 1948; Emery
et al., 1954), the Carolline Islands (Tracey, 1968), the
Tuamotus and Society Islands (Newell, 1956; Chevalier
et al., 1968), Johnson Island in the central Pacific (Keating, 1985), Clipperton in the east Pacific (Glynn et al.,
1996), the Solomon Islands (Stoddart, 1969), the CocosKeeling Islands and the Maldives in the Indian Ocean
(Williams, 1994; Bianchi et al., 1997), the Bahamas and
Caymans in the Caribbean (Blanchon and Jones, 1995;
Newell, 1951), and Bermuda in the Atlantic (Stanley and
Swift, 1968). In short, the slope break and the sandcovered 10-fathom terrace are ubiquitous geomorphological features and clearly represent the terrace and riser
(cliff) of an abandoned submerged shoreline (Blanchon
and Jones, 1995).
Formation of terrace-scarp reef-front features
Hopley (1982) argued that similar terraces on the Great
Barrier Reef represented common levels of reef growth
during glacial interstadials. A similar incremental explanation was suggested by Paulay and McEdward (1990),
who simulated reef development over the last 125 ka.
Both dismissed marine erosion as insignificant based on
micro-erosion-meter measurements of fairweather coastal
erosion (Spencer, 1985). However, mathematical modeling using more realistic erosion rates has demonstrated
that marine planation is capable of producing terraces in
all but the most resistant lithologies in less than 5 ka
(Trenhaile, 1989). Furthermore, the presence of the sand
terrace around islands with a history of uplift during the
Pleistocene such as Barbados in the Caribbean and raised
atolls in the Vanuatu chain, south Pacific, indicates that it
must have developed during the postglacial rise in sea
level (Blanchon and Jones, 1995). In fact, several drowned
early Holocene breakwater reefs have been discovered
associated with the terrace (Lighty et al., 1978; Blanchon
et al., 2002). The crest of an 8-ka old reef was discovered
around Grand Cayman along the inner edge of the terrace
at a depth of $20 m, and its elevation corresponded with
the position of an intertidal notch preserved in the midshelf scarp on the other side of the island (Blanchon
et al., 2002). A correlation with similar elevations of
drowned breakwater reefs, notches, and terraces developed around other islands in the Caribbean led Blanchon
et al. (2002) to propose that these reefs and their shorelines
were drowned by a meltwater pulse (Mwp-1c) during
the early Holocene that was associated with the final
disintegration of the Laurentide Ice Sheet. The timing of
this pulse is also consistent with further discoveries of
drowned reefs in the Gulf of Carpentaria (Harris et al.,
2008) and has been confirmed in several other non-reefal
records of early Holocene sea level (e.g., Hijima and
Cohen, 2010).
Much of the emphasis in these various explanations has
been placed on the terrace-forming process. Yet, it is difficult to prove from these data alone if the terrace resulted
from marine erosion or from incremental reef accretion.
Also, it is no surprise that the terrace elevation differs
slightly between reefs, given the variation in sediment
and reef deposition during the Holocene; however, neither
of these factors is particularly important to explain. This is
because the notched scarp at the head of the terrace clearly
shows it formed the shoreline behind reef crests that
drowned during the 8-ka meltwater pulse. As a consequence, the presence of this scarp separating terraces in
reef-front zones in other areas indicates that they too were
subject to same shoreline development $8 ka ago. The
surprising thing is that the notched scarp survived
shoreface erosion during the subsequent sea-level rise
and went on to influence reef development (Blanchon
and Jones, 1995). Indeed, the preservation of notched
scarps and abandoned reef crests in different oceans is testament to the sudden acceleration of sea level during that
8-ka meltwater pulse. It also explains why modern breakwater reef development in these areas did not start until
after that 8-ka event (Blanchon et al., 2002).
Reef slope
In many areas, the reef slope is the most poorly known of
the major geomorphic zones. It represents the outermost
zone of reef construction and can be arbitrarily divided
into two subzones: an upper subzone from the shelf break
to the limit of effective coral-framework development
at $60 m, and a lower subzone down to 120 m or more
that may be dominated by a sciaphyllic framework of
sclerosponges, coralline algae, and microbiallites. The
geomorphology of this lower subzone has been modified
by reef accretion but still bears a strong imprint of sealevel change during the last glacial–interglacial hemicycle
(e.g., Anderson, 1998).
Shelf-edge chute and buttress (shelf-edge reefs)
Slope breaks associated with the mid-shelf scarp and shelf
edge are often sites of preferential reef development
because sediment bypasses these steeper areas on its way
to depocenters on sand terrace and at the base of the reef
slope, respectively; this is known as the “edge effect”
(Porter, 1972). The shelf-edge site in particular is commonly associated with spectacular reef development (as
shown in Figures 1 and 7), which has been described from
various locations as either, sill reefs, shelf-edge reefs, or
mesophotic reef systems (Goreau and Goreau, 1973;
Burke, 1982; Blanchon and Jones, 1997; Lesser et al.,
2009). This reef system develops from the brow of the
GEOMORPHIC ZONATION
479
a larger sand reservoir can prohibit spur development
completely.
The sand terrace is a prominent feature seen in imagery
from all oceans (Figure 6) and has been widely reported as
the “ten fathom” terrace (Emery, 1961; Garrison and
McMaster, 1966). Where it is not covered by excessive
reef development, the terrace has been documented in
roughly the same 20 Æ 5 m depth range from the following
areas: the Marshall Islands (Tracy et al., 1948; Emery
et al., 1954), the Carolline Islands (Tracey, 1968), the
Tuamotus and Society Islands (Newell, 1956; Chevalier
et al., 1968), Johnson Island in the central Pacific (Keating, 1985), Clipperton in the east Pacific (Glynn et al.,
1996), the Solomon Islands (Stoddart, 1969), the CocosKeeling Islands and the Maldives in the Indian Ocean
(Williams, 1994; Bianchi et al., 1997), the Bahamas and
Caymans in the Caribbean (Blanchon and Jones, 1995;
Newell, 1951), and Bermuda in the Atlantic (Stanley and
Swift, 1968). In short, the slope break and the sandcovered 10-fathom terrace are ubiquitous geomorphological features and clearly represent the terrace and riser
(cliff) of an abandoned submerged shoreline (Blanchon
and Jones, 1995).
Formation of terrace-scarp reef-front features
Hopley (1982) argued that similar terraces on the Great
Barrier Reef represented common levels of reef growth
during glacial interstadials. A similar incremental explanation was suggested by Paulay and McEdward (1990),
who simulated reef development over the last 125 ka.
Both dismissed marine erosion as insignificant based on
micro-erosion-meter measurements of fairweather coastal
erosion (Spencer, 1985). However, mathematical modeling using more realistic erosion rates has demonstrated
that marine planation is capable of producing terraces in
all but the most resistant lithologies in less than 5 ka
(Trenhaile, 1989). Furthermore, the presence of the sand
terrace around islands with a history of uplift during the
Pleistocene such as Barbados in the Caribbean and raised
atolls in the Vanuatu chain, south Pacific, indicates that it
must have developed during the postglacial rise in sea
level (Blanchon and Jones, 1995). In fact, several drowned
early Holocene breakwater reefs have been discovered
associated with the terrace (Lighty et al., 1978; Blanchon
et al., 2002). The crest of an 8-ka old reef was discovered
around Grand Cayman along the inner edge of the terrace
at a depth of $20 m, and its elevation corresponded with
the position of an intertidal notch preserved in the midshelf scarp on the other side of the island (Blanchon
et al., 2002). A correlation with similar elevations of
drowned breakwater reefs, notches, and terraces developed around other islands in the Caribbean led Blanchon
et al. (2002) to propose that these reefs and their shorelines
were drowned by a meltwater pulse (Mwp-1c) during
the early Holocene that was associated with the final
disintegration of the Laurentide Ice Sheet. The timing of
this pulse is also consistent with further discoveries of
drowned reefs in the Gulf of Carpentaria (Harris et al.,
2008) and has been confirmed in several other non-reefal
records of early Holocene sea level (e.g., Hijima and
Cohen, 2010).
Much of the emphasis in these various explanations has
been placed on the terrace-forming process. Yet, it is difficult to prove from these data alone if the terrace resulted
from marine erosion or from incremental reef accretion.
Also, it is no surprise that the terrace elevation differs
slightly between reefs, given the variation in sediment
and reef deposition during the Holocene; however, neither
of these factors is particularly important to explain. This is
because the notched scarp at the head of the terrace clearly
shows it formed the shoreline behind reef crests that
drowned during the 8-ka meltwater pulse. As a consequence, the presence of this scarp separating terraces in
reef-front zones in other areas indicates that they too were
subject to same shoreline development $8 ka ago. The
surprising thing is that the notched scarp survived
shoreface erosion during the subsequent sea-level rise
and went on to influence reef development (Blanchon
and Jones, 1995). Indeed, the preservation of notched
scarps and abandoned reef crests in different oceans is testament to the sudden acceleration of sea level during that
8-ka meltwater pulse. It also explains why modern breakwater reef development in these areas did not start until
after that 8-ka event (Blanchon et al., 2002).
Reef slope
In many areas, the reef slope is the most poorly known of
the major geomorphic zones. It represents the outermost
zone of reef construction and can be arbitrarily divided
into two subzones: an upper subzone from the shelf break
to the limit of effective coral-framework development
at $60 m, and a lower subzone down to 120 m or more
that may be dominated by a sciaphyllic framework of
sclerosponges, coralline algae, and microbiallites. The
geomorphology of this lower subzone has been modified
by reef accretion but still bears a strong imprint of sealevel change during the last glacial–interglacial hemicycle
(e.g., Anderson, 1998).
Shelf-edge chute and buttress (shelf-edge reefs)
Slope breaks associated with the mid-shelf scarp and shelf
edge are often sites of preferential reef development
because sediment bypasses these steeper areas on its way
to depocenters on sand terrace and at the base of the reef
slope, respectively; this is known as the “edge effect”
(Porter, 1972). The shelf-edge site in particular is commonly associated with spectacular reef development (as
shown in Figures 1 and 7), which has been described from
various locations as either, sill reefs, shelf-edge reefs, or
mesophotic reef systems (Goreau and Goreau, 1973;
Burke, 1982; Blanchon and Jones, 1997; Lesser et al.,
2009). This reef system develops from the brow of the
GEOMORPHIC ZONATION
479