Chapter 8 – sea level rise causes, impacts, and scenarios for change

Chapter 8

Sea-Level Rise: Causes,
Impacts, and Scenarios
for Change
Robert J.N. Devoy
The Coastal and Marine Research Centre, Beaufort Research and the Department of Geography,
University College Cork, Cork, Ireland

ABSTRACT
Studies of sea-level changes (SLCs) are well established, as is reflected in an extensive
research and wider publication literature. People globally have long recognized that
ancient marine shorelines and beach sediments occur inland, well away from present
day coasts. The causes of SLC center upon a series of primary controls resulting from
the operation of an integrated atmosphereeocean system. These include changes in the
total water volume of the oceans over time, temperature effects on ocean water
expansion, alterations in the shape of the ocean basins from Earth crustal movements,
and changes in their accommodation space through sediment accumulation. Satellite
surveys also show that the topography of the oceans and shelf seas is influenced by
gravitational controls, which define the Earth geoid, with an up to 180 m difference in
interregional sea surface heights. In the Quaternary, past glaciations have driven sea
levels to approximately À120 m global mean sea level (gmsl). The significance of SLC
today lies in establishing an understanding and projections of how sea levels will
change in the future. Climate warming during the twenty-first century will result in the
melting of the Earth’s remaining ice masses, with gmsls likely to rise at rates of
4e5 mm/year by 2050, reaching levels of 0.5e0.9 m by 2100 under differing Intergovernmental Panel on Climate Change (IPCC) global warming scenarios. The impacts
of these movements will potentially result in initial losses of up to 30% of coastal
wetlands and an increasing “squeeze” of people and biological systems into the reorganizing coastal zone.

8.1 INTRODUCTION
Sea-level and sea surface changes (SLCs) are a primary driver in coastal

systems’ functioning; their analysis and quantification forms a critical element
in the study of coastal environments (e.g., Devoy, 1987; Warrick et al., 1993;
Coastal and Marine Hazards, Risks, and Disasters. />Copyright © 2015 Elsevier Inc. All rights reserved.

197


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Carter and Woodroffe, 1994; Barthel et al., 1999; Smith et al., 2000; Church
et al., 2010; Nicholls, 2010; Nicholls et al., 2007, 2011; Bengtsson et al., 2012;
Wong and Losada, 2014; Church et al., 2014; Masselink and Gehrels, 2014;
Pugh and Woodworth, 2014). The fact that SLCs have occurred regularly over
varying timescales, both in vertical extent (height range) and consequent
spatial exposure of the world’s continental margins, is well established (e.g.,
Devoy, 1979, 1987, 1997; Hallam, 1981; Haq et al., 1987; Haq, 1991;
Pirazzoli, 1996; Peltier, 1998, 2004; Cronin, 1999; Edwards, 2006; Bindoff
et al., 2007; Haq and Schutter, 2008; Muller et al., 2008; Murray-Wallace and
Woodroffe, 2014; Church and Clark et al., 2014) (Figure 8.1). Most recently in
Earth’s history, the Quaternary glacial and interglacial cycles have been the
main control in these vertical movements of coastal positions worldwide.
Changes in ice-mass volumes (from glaciers and ice sheets) through
CrolleMilankovitch forcings of Earth temperature and the linked operation of
the hydrological cycle, have caused a wide range of sea-level movements. Icemass changes have driven the position of the coastal zone (CZ) regularly
across the continental shelves at timescales of 103e4 years over the last
0.5e1.0 M years (Lowe and Walker, 1997; Siegert, 2001) (Figure 8.2). This
close correlation of glaciations with SLC and its crustal impacts was recognized
early by the first Quaternary and Earth scientists, such as Agassiz, Jamieson, and

later Daly and Zeuner (Smith and Dawson, 1984; Devoy, 1987; Dawson, 1992).
The erosion associated with these CZ movements may have created the low-angle
surfaces (<1 ) and wide extent of many passive ocean-continental shelf margins
(Open University, 1998). Low-stand sea-level positions of À120 m to À130 m
global mean sea level (gmsl) (e.g., Fairbanks, 1989; Bard et al., 1990, 1996;
Peltier, 1998), during glaciations and high stands in interglacials of þ2 to
6 m gmsl have occurred during the Quaternary, with changed coastal positions
identified from most world coastal regions (Figure 8.1(b)) (e.g., Shackleton,
1987; Chappell et al., 1996; Church et al., 2010; Murray-Wallace and Woodroffe,
2014). Since the ending of the last glacial stage (10, 000 years ago), the subsequent Holocene rates of sea-level rise (SLR) have fallen from high values of
20e40 mm/year (Bard et al., 1996), following the major deglaciations in the
Northern Hemisphere, to relative mean global rates of <1 mm/year during the
late Holocene, post-3000 years BP (e.g., Bindoff et al., 2007; Delaney et al.,
2012; Church and Clark et al., 2014).
The Earth history connection between continual SLC and CZ movements
has been highlighted in the five Intergovernmental Panel on Climate Change
reports since 1990 (1990e2014) (e.g., IPCC WGI, 2001, 2007, 2014; IPCC
WGII, 2014). People ignore the message of the past scales and cycles of
rapid environmental changes at their peril: SLC is a good analog in developing an understanding of the dynamic and complex functioning of the
Earth. Detailed studies of former SLCs have occurred since the 1970s, with
the development under UNESCO of the International Geological Correlation
Programme’s projects (IGCP) on “Sea levels.” These have continued


Chapter j 8

Exxon Sea Level Change (m)

(a)

199

Sea-Level Rise

400
Global Sea Level Fluctuations
300

Hallam et al.

200

100

0

–100

N Pg

K

50

100

0

(b)

Exxon Sea
Level Curve

Last
Glacial
Episode

0

20

J
150

40

Tr

P

C

200 250 300 350
Millions of Years Ago

D

S

400

Thousand Years Before Present
60
80
100
120

O

450

Cm
500 542

140

160

20
0
Sea Level (m)

–20
–40
–60
–80

–100
–120
–140

FIGURE 8.1 (a) The pattern of long-term (106À8 years) global-relative, SLCs, interpreted from
Exxon-type continental Shelf sedimentary coastal on-lap and off-lap data (Source: Open web, after
Vail et al. (1977), Hallam (1981), and Haq (1991).). (b) Late Quaternary RSL changes for Marine
Isotope Stages 1e5, based upon geomorphological shoreline, coral-terrace data from the Huon
Peninsula, Papua New Guinea (dots), and oxygen isotope records from benthic foraminifera
(squares) (Source: Open web source figure, based upon data from Peltier (1998), after Shackleton
(1987) and Chappell et al. (1996).).

unbroken to the present day. Many of the projects, for example, Projects 61,
200, 274 (Greensmith and Tooley, 1982; van de Plassche, 1986; Pirazzoli,
1991, 1996) have concentrated on the proxy records and operations of past
SLCs. More recent work, for example, Project 588, “Preparing for Coastal
Change” have made links to CZ process and applied research (IGCP 588,


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Coastal and Marine Hazards, Risks, and Disasters

(a)

Hydrological cycle
Glaciers
Atmosphere–ocean
interaction

Ground water

Relative
sea level

Ice sheets
and shelves

Ocean properties
Ocean circulation
Geocentric
sea level

(b)
Storms

External
Marine
Influences

Waves

CLIMATE CHANGE
Sea Level
Temperature

Natural
System

Human
System

CO2 conc.

Run-off

External
Terrestrial
Influences

Coastal System
FIGURE 8.2 (a) Operation of the coupled OceaneAtmosphere systems and their linkage to RSL
changes at the land margins, in which the large-scale masseenergy fluxes and the ocean water
properties, for example, of temperature, salinity, and density, will influence RSL values and the
positioning of the CZ (Source: Church et al. (2014).). (b) Definition and functioning of the CZ in the
context of RSL changes at the landeocean margins (Figure 8.2(a)) and showing the links to climate
forced coastal drivers (Source: IPCC WGII, 2007, Nicholls et al., 2007.).

2014). In the late 1990s, the IGCP projects were joined by the International
Geosphere and Biosphere Programme (IGBP), with the study of SLC and
coastal issues represented by the Land and Ocean Interactions in the Coastal
Zone (LOICZ) program (LOICZ, 2014). These international SLC studies and
many others worldwide, together with their incorporation into the understanding of coastal processes functioning (e.g., Carter and Woodroffe, 1994;
Hardisty, 1990, 1994; Duffy and Devoy, 1999; Sanchez-Arcilla et al., 2000;
Davison-Arnott, 2010; Masselink et al., 2011), have produced an extensive
sea level and related literature. Examples include, the NATO Advanced


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Science Institutes Research Workshop Series (as per the titling in the book
Sabadini et al., 1991); the European Science Foundation colloquium series,
Interglacial Sea-level Changes in 4D (Devoy, 1999); the European “Environmental Framework Research Programmes IIIeVII” (EU CORDIS Website, 2014; Barthel et al., 1999; Smith et al., 2000); and in the United States,
Sea Grant (federally sponsored by the National Oceanic and Atmospheric
Administration), the Environmental Protection Agency and many other
Federal and States’ funding programs (e.g., Barthel and Titus, 1984; Titus
et al., 1991, 2010). It should be noted that although much of this work began
in the mid-twentieth century in the Euro-North American “centered world,”
including Australia and New Zealand, it has only recently been owned and
developed by other societies, for example, India, Japan, China and Southeast
Asia, Africa, and South America (e.g., viz. IGCP Projects; IGBPeLOICZ;
Devoy, 1987; Tooley, 1993; Murray-Wallace and Woodroffe, 2014). Current
concerns over future SLR under climate warming, as a cause of coastal
change, also now link SLR intimately to the fields of atmosphereeocean
systems modeling (e.g., IPCC, 2001a; IPCC WGI, 2007, 2014). Work in this
field is involved in two key areas relevant to coastal vulnerability issues.
First, in establishing the ranges of projected SLR to 2100 and beyond, with
the past sea-level records for use as analog and in the validation of projected
SLC. Second, in helping develop the viable management options for future
coasts (e.g., Barthel et al., 1999; Devoy, 2008; Cooper and Cummins, 2009;
Titus et al., 1991, 2010; Gray et al., 2014; Cooper and Pilkey, 2012; Cooper
et al., 2014; CoastAdapt, 2014).

8.2 SEA-LEVEL DEFINITIONS, MEASUREMENTS,
AND CAUSES
Sea level is the primary reference surface in establishing the point of the
landeocean divide in Earth systems and consequently in the definitions of the
CZ (Figure 8.2) (Woodroffe, 2002; Cooper and Cummins, 2009; Nicholls
et al., 2007, 2010; Wong and Losada, 2014). SLCs determine the immediate
position of coastlines and the effective extent and wider spatial movements of

the CZ over time. Further, it is the surface level or “platform,” as by analogy
with a medieval “siege tower,” from which hydrodynamic processes operate in
developing shorelines (Carter and Devoy, 1987; Carter, 1988; Hardisty, 1990,
1994; Carter and Woodroffe, 1994; Stone and Orford, 2004; Masselink et al.,
2011; Tibaldi et al., 2012; Ranasinghe et al., 2013; Kremer et al., 2013; Devoy,
in press). The sea surface is constantly perturbed by wind-driven (gravity)
waves, tides, storm surges, tsunamis, many “miscellaneous” wave forms, the
effects of internal and long-period waves and by ocean currents, steric, density,
and meteorological dynamic effects (Warrick et al., 1993; Smith et al., 2000;
Open University, 1998; Pugh, 2004; Church et al., 2014). The surface is
spatially (localeregional) and temporally complex in its form and functioning,


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Coastal and Marine Hazards, Risks, and Disasters

differing in shape and elevation from microscale to macroscale (1 m2 to
103 km and seconds to 106e8 years). In paleoenvironmental reconstructions of
SLCs (e.g., the IGCP projects; Peltier, 1998, 2009; Lambeck, 1991, 1995,
2001; Church et al., 2010; Murray-Wallace and Woodroffe, 2014), and from
the contemporary monitoring of sea levels, it is clear that no globally uniform
sea level exists (van de Plassche, 1986; Warrick et al., 1993; Smith et al., 2000;
Pugh, 2004).
The sea surface is commonly defined at localeregional scales in relation to
a mean sea level (msl) (i.e., the averaged position of all states of sea level over
a period of time, generally greater than one year), as a mean tide level, or
referenced to a bathymetric, chart datum on marine maps; in the British Isles
taken commonly as the lowest point of astronomical tides (e.g., Woodworth,
1990, 1993, 1999; Woodworth et al., 1991, Woodworth and Blackman, 2004;

Zerbini et al., 2000; Pugh, 2004; Church et al., 2010; Cronin and Devoy, 2010;
Admiralty Tide Tables, 2014). These levels have been measured and monitored by land-based tide gauges at reference locations that are situated ideally
in “stable” locations, free of Earth crustal subsidence and uplift movements.
This is often difficult to determine accurately; hence, the relocation of datum
points at times (e.g., in Britain, from Liverpool to Newlyn in 1921, or in the
Republic of Ireland from Dublin to Malin Head in 1965). Historically, these
expressions of sea level (e.g., msl) have been used as the reference point for
geodetic land height surveys Ordnance Datums, for example, OD (Britain),
NAP (the Netherlands), NGF (France), and will vary between regions, countries, and continents (IOC, 1985, 2009; Woodworth et al., 1991, 2002, 2007,
2009a; Woodworth and Player, 2003). These datums do not allow for easy
comparison of sea surface height changes over long distances, or even the
within-region recognition of the complexity in the operation of SLC (e.g., at
scales of >50e100 km). Consequently, the advent of satellite altimetry and its
integration into more extensive and precise land-based tide gauge networks,
and whole earth spatial coverage, have led to a much better understanding of
sea-level variability (e.g., Global Sea-level Observing System, GLOSS, 2014;
Merrifield et al., 2009; IOC, 2009) and the development of geocentric systems
of measuring Earth shape (Carter et al., 1993; Zerbini et al., 2000)
(Figure 8.3). These survey systems now form the basis for measurements of
SLC and particularly for definition of the first-order, long-term “static-semistatic” component of sea surface levels, namely, differences in the Earth geoid,
that is, the gravitational equipotential surface shape of the Earth (Warrick
et al., 1993; Zerbini et al., 2000; Pugh, 2004; Masselink and Gehrels, 2014).
The geoid surface results primarily from the gravitational effects of land
masses, rock density differences with Earth geophysical functioning, and
extraterrestrial gravitational forces.
Although predicted earlier (e.g., Lisitzin, 1974; Carey, 1981), satellite
surveys clearly show the geoid distortions of Earth shape, with the distribution
of relative high and low points of the sea surface (Figure 8.4). These evidence

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FIGURE 8.3 (a) Definition of sea-level surfaces in relation to the geoid and of other key
reference parameters, using satellite altimetry (Source: Zerbini et al. (2000).). (b) Land-based tide
gauge and satellite altimetry-linked, monitoring station (Source: Open Web, National Tidal Centre,
Australia.).


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Coastal and Marine Hazards, Risks, and Disasters

(a)

(b)

(c)

FIGURE 8.4 (a) The Earth’s geoid, measured relative to the reference ellipsoid and showing areas of
maximum surface elevation (orangeered) of þ85 m and minimum levels (blueepurpleeblack) of
approximately 106 m. The differences result primarily from the uneven mass and density distributions
within the Earth (Source: Pugh (2004).). (b) Earth geoid shape (Source: Open web). (c) Dynamic sea
surface as gmsl variations from the geoid, measured in dynamic centimeters, caused by meteorological,
ocean density, and current changes with time. The black arrows show the main ocean current pathways
and are associated with some of the largest height differences (Source: From Pugh, Changing Sea
Levels, (2004).).

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205

significant relative sea-level (RSL) height differentials of up to 180 m between
the “lows” and “highs,” for example, of southern India compared to the
centraleeastern North Atlantic. Earlier conjecture that these “static” RSL
positions migrate in long-term timescales (>106 years) (e.g., Mo¨rner, 1987a,b)
lack support (Devoy, 1987; Tooley, 1993; Peltier, 1998). Further, it is recognized that at continental land margins (Clark et al., 1978; Clark and Primus,
1987; Clarke et al., 2005; Church et al., 2014), and also around major ice
masses, other mass/gravity distortions of the geoid exist (scales of >1-m sea
surface set up). Melting of the present day major ice masses (e.g., for the
Greenland and Antarctica ice sheets) will lead to the relaxation of the icemarginal sea surfaces and to regional and wider redistributions of the gravitationally held water “bulges” (e.g., Peltier, 1998, 2004; Lambeck, 2001;
Milne et al., 2009; Mitrovica et al., 2009, 2011; Bindoff et al., 2007; Church
et al., 2010; Church et al., 2014).
Additionally, the return of land-based melt water to the oceans under
Earth rotation results in centrifugal movements of water and large-scale time
and spatial differential SLCs. In the lateglacial deglaciation of the Northern
Hemisphere, early rapid rises of sea level occurred in the Southern Hemisphere in response to the ice melt, as reflected in the Holocene “family” of
sea-level curves (Thom and Roy, 1983; Devoy et al., 1994; Murray-Wallace
and Woodroffe, 2014), with RSL reaching or exceeding present levels by
6,000 BP (Figure 8.5(a) and (b)). Subsequent equilibriation of water levels
took place, with return flows northward later in the Holocene and RSL
continuing to present levels (Clark and Lingle, 1977; Clark and Primus,
1987; Tooley, 1993; Pirazzoli, 1996; Peltier, 1998, 2004, 2009; Milne and
Mitrovica, 1998), with overall rapid early rises of gmsl (Bard et al., 1996).

These changes in ice and ocean water masses have feedback effects on
changes in Earth tilt, “wobble” and rate of spin, which also undergo other
periodic changes over time, and together result in variations in sea surface
levels (Church et al., 2014). Under modeled ice melt with future climate
warming (to 2100 and beyond), then similar mechanisms of water flows and
the recording of SLC in the Northern and Southern Hemispheres will occur.
These will be consequent upon the differences in the expected timings of
Greenland and Antarctic downwasting, though rapid global SLC responses to
ice melting are likely (Church et al., 2014; Mitrovica et al., 2011; Rahmstorf,
2007, 2010, 2012).
More significant at these large-regional spatial scales are the short term
(101 years) to rapid (seconds to days) movements in sea levels created by
the dynamic sea surface (Lisitzin, 1974; Zerbini et al., 2000; Pugh, 2004;
Church et al., 2010) (Figure 8.4(c)). These are caused mainly by meteorological and coupled Earth atmosphereeocean energy drivers (including,
e.g., ocean currents and steric changes, atmospheric pressure fields, winds,
open ocean rainfall, and large-scale river discharges). Resultant phenomena
(e.g., El Nino Southern Oscillation (ENSO), El Nin˜o, and La Nin˜a events;


206

Coastal and Marine Hazards, Risks, and Disasters

FIGURE 8.5 (a) The Holocene “family” of RSL curves from Northwest Europe, showing the
different coastal zonal “signatures” in simplified sea-level trends resulting from Earth crustal
isostatic responses to loading changes following deglaciation (Source: From Carter, 1992, with
permissions from the Quaternary Research Association.) and (b) the Holocene SLC patterns from
areas of Australia and New Zealand, indicating RSLs at or above present day levels by 6,000 years
BP (Source: Devoy et al. (1994); Thom and Roy (1983).). (c) Approaches in the construction and
identification of Holocene sea-level patterns, trends and tendencies. (Source: From Carter, 1992,

with permissions from the Quaternary Research Association.).


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207

FIGURE 8.5 cont’d

Pacific Decadal Oscillation (PDO); North Atlantic Oscillation (NAO)),
produce changes in mean sea surface heights of 1e2 m from the geoid and
locally >0.3 m (Figure 8.6). These sea surface variations may be interannual, or longer-term periodic to quasicyclical in occurrence. In the case of
ENSO events, return periods of 3e7 years in the nineteenth/twentieth
centuries (McGregor, 1992; IPCC WGI, 2014). Paleo-El Nin˜o events from
proxy ocean core records show different frequencies during the Holocene
(Ortlieb and Maclane, 1993; Fagan, 2000). Most importantly, the warming
of ocean water results in the expansion of the water column and thermostatic (steric) SLR. This effect is regionally varied, due to changes in water


208

FIGURE 8.5 cont’d

Coastal and Marine Hazards, Risks, and Disasters


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Sea-Level Rise

(a) Normal conditions

EI Nino conditions

(b)

FIGURE 8.6 (a) Equatorial cross-section of the Pacific Ocean, showing the development of an El
Nino event, showing normal conditions in January 1997 to the full El Nino development in
November 1997 in the eastern Pacific, with raised sea surface levels. Changed sea surface
topography is shown linked to the temperature changes, 30  C as red to 8  C as dark blue (Source:
Open web, NASA.). (b) El Nin˜o event developing in 1992, showing the height detail of the surface
topography (Source: Open Web.).

density and linked characteristics. Estimates indicate that during the period
from 1954e2014, steric changes have caused averaged rates of SLR of
0.33 mm Æ 0.7 mm/year, to as high as 0.52 mm/year (in the top 700 m of
the oceans), with rates of 0.4 mm Æ 0.9 mm/year indicated for the global
oceans to depths of 3000 m. Maximum effects of these changes are
concentrated in the subtropical North Atlantic and tropical eastern Pacific
regions (IPCC WGI, 2001, 2007, 2014).
Dynamic sea surface changes, particularly where coupled with storminess
patterns, have important consequences on coasts, in terms of increased rates of
erosion, flooding from coupled marine inundation, and river discharges and


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Coastal and Marine Hazards, Risks, and Disasters

through wider impacts in sediment movements and coastal processes functioning (Duffy and Devoy, 1999; Woodroffe, 2002; Nicholls et al., 2007; Saito
et al., 2007; Kremer et al., 2013). Along East Pacific coasts, El Nin˜o events
result in elevated sea surfaces of >0.1 m and La Nin˜a episodes can cause
lowered surfaces (Figure 8.6). When these temporary changes are associated
with storminess, significant phases of increased erosion and beach change can
occur, as recorded along western coasts of the United States, eastern Australia,
and in the wider central Pacific region (Short et al., 1995; Storlazzi and Griggs,
2000; Soloman and Forbes, 1999; USGS, 2011). Widespread impacts can also
occur upon offshore marine systems in sea surface temperature (SST) changes,
ocean productivity, and famine consequences for people throughout the South
American Pacific regions (Open University, 1998; Fagan, 2000; IPCC, 2001a;
IPCC WGI and WGII, 2007, 2014). Operation of the NAO and linked changes
in the position of the Icelandic Low Pressure system, influenced possibly by
variations in Arctic sea-ice distributions, have similar erosion and coastal
process impacts in the North Atlantic (Stone and Orford, 2004). Periods of
increased storminess in the late Holocene have been identified from this region, associated with barrier breaching, coastal flooding, and other sedimentaryehydrodynamic consequences resulting from these teleconnection
types of factors (Gilbertson et al., 1999; Dawson et al., 2004; Delaney et al.,
2012).

8.2.1 Larger-Scale Causes for SLCs
The longer-term causes of SLCs have been a topic of study since at least the
eighteenth Century and subject to varying changes in paradigms and understanding through the twentieth Century (Smith and Dawson, 1984; Devoy,
1987; Shennan and Tooley, 1987; Smith et al., 2000; Murray-Wallace and
Woodroffe, 2014). Much debate has centered on the definition of the term
eustasy, which formerly implied absolute changes of water level within the
oceans and controlling the position of shorelines (e.g., Suess, 1906; Daly,
1934; Visser, 1980; Tooley, 1993). Attached to the concept initially was that

SLCs were global in nature and the term was synonymous with the “water
volume” of the oceans. The work of the IGCP Sea Levels Projects 61, 200 and
274 (e.g., Greensmith and Tooley, 1982; van de Plassche, 1986; Pirazzoli,
1996) particularly, altered this understanding. They established clearly the
worldwide variability of Holocene and Quaternary interglacial SLC, knowledge of the component drivers to SLCs and the range and heterogeneity of data
used in the reconstruction of past sea levels. Further, this IGCP project series
(200, 274, 367, 437.588), among other research programs, raised questions
on the value and the approaches needed in using paleosea levels as predictors
of future SLCs and of the limitations of these classic geologically based sealevel studies in establishing coastal functioning (Devoy, 1987; Shennan and
Tooley, 1987; Pirazzoli, 1996; Smith et al., 2000). What became clear in the


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FIGURE 8.7 Primary controls in the development of relative sea-level changes. Source: From
Dawson (1992); after Mo¨rner (1980).

1980s is that eustasy, and the other linked controls on sea-level positions, are
not spatially and temporarily uniform, or worldwide in operation. This understanding is supported by the work on the causes and drivers to longer-term,
pre-Quaternary RSL changes (Vail et al., 1977; Hallam, 1981; Haq et al., 1987;
Haq, 1991; Haq and Schutter, 2008; Conrad, 2013) (Figure 8.1(a)).
As morphodynamic and hydrodynamic studies of coastal processes
emphasize (e.g., Carter and Woodroffe, 1994; Duffy and Devoy, 1999;
Short, 1999; Cooper et al., 1995, 2001, 2007; ICS, 2002, 2009, 2011, 2013;
Ranasinghe et al., 2013; O’Shea and Murphy, 2013; Cooper and Jackson,
2014) coasts operate at local scales. In turn, the same is true in the deconstruction of the causes of SLCs and their use in determining shoreline positions. SLCs comprise many macroscale to microscale temporal and spatial

Earth environmental factors. Understanding of these SLC components helps
define the term eustasy as, “absolute sea-level changes regardless of causation
and including the main family of vertical and horizontal geoid and dynamic
changes” (Mo¨rner, 1987a,b) (Figure 8.7 and Table 8.1). The significance of
this approach to SLCs has been in making the linkage to problem solving in
the study of coastal processes and associated management issues (e.g.,
Jennings et al., 1998; Cowell and Thom, 1994; Cooper et al., 2001, 2007) and


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TABLE 8.1 Summary of the Components Involved in Defining Eustasy and
Relative SLCs

VOLUME

OCEAN

DISTRIBUTION

DYNAMIC

TECTONICS

SEA FLOOR SUBSIDENCE
OTHER EARTH MOVEMENTS
LOCAL ISOSTASV
HYDRO - ISOSTASV

INTERNAL LOADING ADJUSTMENT

WATER IN SEDIMENT, LAKES
AND CLOUDS, EVAPORATION,
JUVENILE WATER

VOLUME

OCEAN

MID-OCEANIC RIDGE GROWTH
PLATE TECTONICS

GLACIAL EUSTASY

WATER

MASS / LEVEL

OROGENY

SEDIMENT IN-FILL
ISOSTASY

BASIN

TECTONO-EUSTASY

OCEAN

GEOIDAL EUSTASY

VERTICAL AND HORIZONTAL GEOID C H A N GE S

C H A N GE S
LE V E L
OC E A N

DYNAMIC
CHANGES

E

U

S

T

A

S

Y

EARTH-VOLUME CHANGES

GRAVITATIONAL WAVES
TILTING OF THE EARTH
EARTH’S RATE OF ROTATION

DEFORMATION OF GEOID RELIEF
(DIFFERENT HARMONICS)

METEOROLOGICAL

SEA LEVEL

HYDROLOGICAL

CHANGES

OCEANOGRAPHIC

Source: After Tooley (1993).


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213

in the development of numerical models in simulations of coastal changes (de
Vriend, 1991; de Vriend et al., 1993; Peltier, 1998; Stive, 2004, 2006; Stive
and de Vriend, 1995; Stive and Wang, 2003; Wang et al., 2010; Cronin et al.,
2009; Cronin, 2010; Dan et al., 2011; Ranasinghe et al., 2013). The approach
emphasizes the limitations of concentrating on the short-term characterization
of coastal systems and the need to incorporate feedbacks and inheritance from
mesoscale controls to macroscale coastal boundary controls, such as SLCs.
Operation of these longer-term factors will continually alter the projections of

coastal-system outcomes provided by the numerical process models used to
establish coastal functioning, such as those of Delft 3-D, Deltares; MIKE 21,
Danish Hydraulics Institute (e.g., Stive et al., 1991; Stive and de Vriend, 1995;
Stive, 2004, 2006; Cronin, 2010; O’Shea and Murphy, 2013; Devoy, in press).
Although varying modeling approaches have been used to help resolve these
scale and feedback issues, the problems remain, particularly in the inclusion of
SLC factors (IPCC WGI, 2007, 2014).

8.2.2 Shorter-Term SLCs and Related Drivers to Coastal
Vulnerability
Assessments and numerical modeling of physical coastal-system changes
under climate warming indicate the critical role of SLR as a forcing control
(Smith et al., 2000; Church et al., 2014). For cohesionless, “soft” sedimentary
coasts (e.g., sand and cobble-sized beachs and dune-barrier systems) the Bruun
Rule relationship of SLR as a “proportional” driver in the landward retreat of
shorelines, will account for 25e50% of likely coastal changes (Ranasinghe
et al., 2013; Nicholls et al., 2007; Devoy, 2008; Church et al., 2010). The
remaining 50% of coastal “retreat” will be influenced by climate-driven
rainfall and river-discharge factors, with iterative repercussions in linked
coastal process operations. In this, SLR will be important in contributing to
feedbacks in coastal-sediment flux and particularly in the sedimentary infilling
of coastal-accommodation spaces (e.g., Cowell and Thom, 1994; Ranasinghe
et al., 2013; de Groot et al., 2012). These types of responses to SLRs are
connected closely to the recognized impacts of people in controlling the
pathways and volumes of material mass (i.e., sediments and water) moving to
coasts from land sources. The storage of water in reservoirs and other artificial
water basins accounts for 3% of potential future SLR (Church et al., 2014).
The impacts of this disruption to drainage systems, including major modifications of drainage basins, from vegetation clearances, changed land uses, and
sediment retention, have had widespread significance for the rates of observed
coastal changes (Bindoff et al., 2007; Nicholls et al., 2007). Where river and

sediment flows have been interrupted, estuaries, deltas, and other coastal
systems have become sediment starved. This has caused feedbacks in with
consequent feedbacks in the effectiveness of SLR in driving increased rates
of coastal erosion and the observed landward retreat of coastal


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Coastal and Marine Hazards, Risks, and Disasters

wetlandsdecosystems. Examples of this phenomenon are common, from the
Mississippi River, where this was in part a cause of the 2004 Hurricane Katrina
disaster in New Orleans, to the GangeseBrahmaputra river catchments and the
large delta systems of Southeast Asia (Saito et al., 2007). Similarly, the impact
of large-scale groundwater pumping, for coastal megacities and in estuaryedelta zones, has speeded land subsidence in these areas since the 1950s and
added to the acceleration of SLRs (Smith et al., 2000; Nicholls et al., 2007;
Church et al., 2010; Church et al., 2014).

8.3 HUMAN LINKS AND DRIVERS: IMPACTS OF SLCs ON
PEOPLE
Understanding the components and causes of SLCs has its intrinsic value as
part of science and the study of Earth Systems’ functions (Smith et al., 2000;
Schwartz, 2006). Yet the primary significance of SLCs is in their likely impacts on the risks and vulnerabilities for people living in the CZ and linked
marine environments (e.g., ISOS, 1991; Barthel et al., 1999; de Groot and
Orford, 2000; Lim et al., 2005; Adger et al., 2005; Oft and Tsuma, 2006;
Olsson et al., 2004, 2007; IPCC WGII SPM, 2007; Alcamo et al., 2007;
Devoy, 1992, 2008; Nicholls et al., 2007, 2011; Lange et al., 2010; Marchand,
2009; Nicholls, 2010; European Commission, 2011; Gray et al., 2014; Wong
and Losada, 2014; CoastAdapt, 2014; Cooper et al., 2014). SLCs have long
influenced where and how people live and interact in the CZ (Cooper and

Cummins, 2009). Numerous examples exist (e.g., North Sea, The Netherlands,
Black Sea and Mediterranean Sea margins, Nile, northeast Arabian Sead
Indus delta) of drowned land surfaces and of forced changes in human and
biotic system’s living spaces and habitats following the postglacial SLR (e.g.,
Jelgersma, 1966; Roeleveld, 1974; Jelgersma et al., 1979). Migrations of
people, plants, and animals have also been conditioned by SLCs in the
operation of landbridges (e.g., Fairbridge, 1961; Kurte´n, 1968; Devoy, 1985,
1995; Edwards and Brooks, 2008). The causes of SLC impacts range from the
immediate effects of RSLs and linked coastal process operations, for example,
tsunami, storms, earthquake, and volcanicity-induced land movements
(Figure 8.8), to those of longer-term SLCs from deglaciation and the consequent return of water to the oceans, with marine inundation and earth crustal
land uplift or subsidence (e.g., Pirazzoli, 1996; Sabadini et al., 1991; Peltier,
1998; Lambeck, 1995, 2001, Lambeck et al., 1996, 1998, 2012). Expected
RSL rise of >1 m by 2100 through climate warming, and to higher levels by
2300, will cause increasing rates of SLR (4e5 mm/year by 2030) and the
onshore movement of coastal systems (Bindoff et al., 2007; Devoy, in press;
European Space Agency, 2014; Church et al., 2014; Rahmstorf, 2010, 2012;
Rahmstorf et al., 2007, 2012a; Rahmstorf and Coumou, 2011; Royal Society &
US National Academy of Sciences, 2014).


Chapter j 8

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215

(a)

(b)

FIGURE 8.8 (a) Impacts of the Durban tsunami, 2007 on coastal structures (Photograph Source:
Denis Linehan, Department of Geography, UCC.). (b) Storm damage in January 2014 on coastal
road and infrastructure at Rossbehy, southwest Ireland (Source: Valerie O’Sullivan, Killarney,
Ireland.).


216

Coastal and Marine Hazards, Risks, and Disasters

Currently, 25e30% of the world’s population lives “at the coast” (Æ1 m of
msl), with this figure estimated to rise to 50% by 2100 (Nicholls et al., 2007,
2011;,Wong and Losada, 2014). People and coastal systems will be forced into
diminishing land spaces, as controlled by the continental land margin gradient,
with all the potential effects of “coastal squeeze” established in the coastal
management literature (e.g., Barthel et al., 1999; de Groot and Orford, 2000;
Nicholls et al., 2007; Cooper and Cummins, 2009, 2014, Cooper and Jackson,
2012; Cooper and Pilkey, 2012). In this future onshore movement of coastal
systems, 20e30% of the world’s coastal wetlands will be initially “lost”
through erosion, or inundated by 2050 (Church et al., 2010; Church et al.,
2014; Royal Society & US National Academy of Sciences, 2014). Together
with the physical coastal systems, these wetlands and linked biotic environments will have to adjust within the new spaces developed under SLRs in the
current back-CZ areas. During this process, both people and biological systems will be stressed and vulnerable to the impacts of the increasing magnitudes, and possibly frequencies, of coastal events, such as storm surges,
sediment movements, and erosion (Figure 8.8) (Sanchez-Arcilla et al., 2000;
Lozano et al., 2004; Lowe et al., 2009; Jenkins et al., 2009; Church et al.,
2014). In the case of storms, although such events are themselves causes of
temporary increases in sea surface levels, the effects of storminess will be
increasingly magnified on coasts with the immediate background drivers in
SLR into the twenty-first Century under climate warming and have an intimate

linkage as such to SLCs per se (Wong and Losada, 2014). More widely, the
effects of SLRs with other coastal process functions have had significant
impacts on the world’s coasts. The effects of El Nin˜o-driven coastal erosion for
the western USA, eastern Australia, and central Pacific islands have been
noted earlier (Section 8.2). In southeast Asia, the effects of direct SLRs,
coupled with sediment starvation on coasts from river catchment changes, are
having a profound impact on the major river and delta systems of the region
(Saito et al., 2007; Hijioka et al., 2014). Globally, large delta and estuaries are
similarly recording the impacts of SLRs magnified by the contemporary
increased effects of sediment starvation, land subsidence through groundwater
abstraction, urban developments, and linked changes in wetland extent and
functioning (Nicholls et al., 2007; Wong and Losada, 2014).
Consequently, radical adjustments to these forcings will have to be made
through mitigation and adaptation measures (see Nicholls, this volume). This
will continue to require active political decision making and governance, in the
development of Coastal Zone Management, linked Marine Spatial Planning
policies and legislation, and, importantly, in economic-business-industry and
lifestyle responses (Cooper and Cummins, 2009; European Commission,
2011). For coastal communities, awareness of and responses to these issues of
SLRs and linked coastal changes are often difficult and complicated by a lack
of technical knowledge of integrated coastal-system functioning and of the
wider science. The extensive research literature now established shows the


Chapter j 8

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217

need for coastal inhabitants, and societies as a whole, to build capacity to
respond to future SLRs and coastal changes. This may mean abandoning the
traditional approaches of built, shoreline-protection measures (Cooper and
Pilkey, 2012) and adopting increasingly soft-engineering techniques in defense
strategies and of hazard zoning, together with structured shoreline retreat of
the most vulnerable coasts, as in the large delta areas, such as the Mississippi
River. Equally, coastal dwellers will need to engage with the realities of where
future coasts will be prone to SLR and learn new techniques for adaptation
(e.g., Titus et al., 2010; Cooper and Cummins, 2009; Lim et al., 2005;
European Commission, 2011, 2013; MCCIP, 2012; Gray et al., 2014; Cooper
et al., 2014; Devoy, in press; Wong and Losada, 2014).

8.4 SEA-LEVEL PATTERNS, TRENDS, AND MODELS
Sea-level patterns and trends are a complex subject, but one that underpins the
development and reliability of the modeled projections of future SLRs, for
example, from natural irradiative forcings of Earth climate, as in
CrolleMillankovitch planetry change factors in climate functioning, to
anthropogenically induced atmosphere warming and scenarios for change
(e.g., IPCC WGIII SRES, 2000; IPCC WGI, 2007, 2014; IPCC WGI SPM,
2014). It is not the intention here to present the detail of the different models
used in establishing the patterns and trends in SLCs at the macroscales to
microscales, but to provide a brief guide to the different approaches.
The models used in SLCs and related studies range from conceptualebehavioral types to those of varying numerical form (Woldenberg,
1985; Kirkby et al., 1993; Raper, 1993; Raper et al., 2000; India and Bonillo,
2001; Allen, 1997, 2003; Allen and Pye, 2009; Marchand, 2009; IPCC WGI,
2007, 2014). Numerical models include empirical and deterministic, quasideterministic types, based upon mathematicalestatistical interpretations of
data, for example, sea-level index points (van de Plassche, 1986), which
assume linear relationships, to stochastic and linked simulation modeling
(e.g., de Vriend, 1991; Schlesinger, 1993; Cowell and Thom, 1994; Raper,
1993; Raper et al., 2000; Stive and Wang, 2003; Wang et al., 2010; Dan et al.,

2011; Rahmstorf, 2007, 2012; Tebaldi et al., 2012; Meehl et al., 2007; IPCC
WGI, 2014). Numerical modeling and simulation techniques are the most
relevant in developing projections of SLR under climate warming scenarios.
As with all models, these often suffer from a lack quality, environmentalboundary data as inputs for model forcings, and validation. Also, in the
direct recording of SLCs, good data may be lacking in defining the different
contributors to SLRs, and particularly in helping understand the feedbacks
involved and in the validation of model outputs (Raper, 1993; Raper et al.,
2000; Zerbini et al., 2000; Lowe et al., 2009; IPCC WGI, 2007, 2014; Church
et al., 2010, 2013). However, the quality of these complex models has
improved steadily, particularly with the introduction of “ensemble”


218

Coastal and Marine Hazards, Risks, and Disasters

Atmosphere and Ocean General Circulation Models (AOGCMs) and the
wider use of Regional Circulation Models (RCMs). This modeling and better
knowledge of the constraints on systems contributing to SLC, for example, of
glacier mass balance and ice melt and ocean mixing, have provided greater
confidence in model results (IPCC WGI, 2014). Many would argue that
numerical model projections should still be used only as a guide to possible
environmental system outcomes (e.g., of SLR), rather than certainties
(Cooper and Pilkey, 2007, 2008).

8.4.1 Macroscale to Meso-scale Changes
At the marcro- and meso-scales, the patterns and trends of SLCs, as evidenced
in sedimentary and other Earth proxy records, are controlled by domaindominant, environmental-forcing controls (Devoy, 1982; Carter and Woodroffe, 1994; Jennings et al., 1998). The long-term (106e8 years) SLC drivers of
Earth tectonics, such as ocean ridge and plate growth decay, mountain
building, ocean-sediment accumulation, and geoid changes together, are

shown in Figure 8.7 and Table 8.1 (Devoy, 1987; Ota et al., 1992; Woodroffe,
2002). They are joined by changes in the total volume of water available in the
oceans; through additions of juvenile water via volcanicity and other Earth
exogenic processes, especially in the early- to mid-Phanerozoic, coupled with
water storage and exchanges with epicontinental seas and ocean basins. The
formation of the Mediterranean and the Messinian salinity crisis forms a good
example of this (e.g., Open University, 1998; Muller et al., 2008).
Understanding long-term SLC patterns (106À8e104 years), as measured
against the land “freeboard” margins (Schubert and Reymer, 1985), have been
based on seismic and sea-bed drilling data from the continental shelves (e.g.,
Vail et al., 1977; Hallam, 1981; Haq, 1991; Haq and Schutter, 2008). These
have generated “signature” logs of sedimentary cyclical coastal on-laps and
off-laps (Figure 8.1(a)). The links of these changes to RSLs, and to their
causative geophysical and Earth crustal movements, together with other largescale environmental drivers, have been developed into conceptualebehavioral
models and shown in deterministic numerical expressions. These provide
primarily first-order projections of SLCs over geological timescales (Pitman,
1978; Devoy, 1987; Cronin, 1999).
Similar approaches have developed in the study of the Quaternary
(Figure 8.1(b)) and particularly in the Holocene SLC. Sedimentary and other
data sets of inferred sea-level positions (index points) have been used to
develop timeedepth sea-level curves (Figure 8.5(c)) (Bloom, 1977; van de
Plassche, 1986; Pirazzoli, 1991, 1996). Statistical analyses of these data
(Devoy, 1982, 1987; Shennan and Tooley, 1987; Shennan et al., 1983, 2002;
Shennan and Andrews, 2000) show that empirical, deterministic models provide only a first-order and generally weak explanation of sea-level records,
because of the complexity of controls involved in developing the observed


Chapter j 8

219

Sea-Level Rise

SLC patterns. Stochastic and environmental feedback mechanisms are shown
as becoming increasingly important during the Holocene following reduction
in the initial rapid rises of RSLs (Bard et al., 1990, 1996) and in generating
local to regional patterns of SLCs (Jennings et al., 1998; Roy et al., 1994;
Allen, 1997, 2003; Allen and Pye, 2009). A significant problem in using these
“timeedepth” data are that they invariably represent RSLs, often with significant error margins (van de Plassche, 1986). Analytical approaches have
been developed (e.g., Jelgersma, 1966; Jelgersma et al., 1979; Andrews, 1970;
Tooley, 1978; Mo¨rner, 1980,, 1987a,b; Shennan et al., 1983; Shennan and
Tooley, 1987; Carter, 1992) to address and help reduce these limitations. This
has resulted in attempts at isolating the tectonic, crustal isostatic (uplift and
subsidence), and “ocean water volume” components involved (Figure 8.7),
thereby generating more reliable explanations of the sea-level data as actual
SLC and the reconstruction of shoreline positions. Andrews’ (1970), and
Mo¨rner’s (1971, 1976, 1980, 1987a,b) related conceptualeempirical approaches form early examples of this, expressed generically as:
Ut ¼ Rt À Et

(8.1)

where Ut is the isostatic uplift over time, R is the position of RSL at a specific
time t as measured in reference to present sea level, and Et is the position of
eustatic sea level over time (after Devoy (1987) and Tooley (1993)).
A related method, which introduced a new approach in isolating “sea-level
tendencies,” challenged the use of the height component of proxy sea-level
data, with their many error problems (Figure 8.5(c)). These included those
of measurement errors, the use of inhomogeneous data sets, the compaction
and consolidation of cored sedimentary data, the different possible water-level
meanings of index points and paleotidal changes at different sites (Shennan,

1982; Hennan & Tooley, 1987; Long et al., 1998; Zong and Tooley, 1996;
Shennan and Andrews, 2000; Shennan et al., 2002). In this approach, only the
chronological elements of sea-level data are used, emphasizing the dependence on the number of index point dates available and their precision (e.g., of
radiometric and other “absolute” dating evidence). Analyses provide a basis
for the numerical study and modeling of the patterns of Holocene SLCs at
regional scales, but again, not for the projection of future SLCs:
E ¼ 0:5 sin ðTp=500Þ if T < 300
E ¼ 0:5 sin ðTp=500Þ þ 50ðcos ððT À 3100Þp=2000Þ À 1Þ

(8.2)
if T > 3100
(8.3)

where E is regional eustasy for time T (after Shennan, 1987).
Further, this methodology is dependent on a number of boundary assumptions, particularly of former tidal regimes and vertical land-level movements (Devoy, 1987; Tooley, 1993). This has led to integration in subsequent
analyses and modeling of Holocene sea-level data with independent models of


220

Coastal and Marine Hazards, Risks, and Disasters

Earth crustal behavior and tides at regional scales (e.g., Lambeck, 1991, 2001;
Lambeck et al., 1996, 1998; Shennan et al., 2002, 2006; Edwards and Brooks,
2008; Lowe et al., 2009).
Understanding the “land movement” components of SLCs and RSLs at
varying mesotime and spatial scales has been based upon the geophysical
modeling of crustal rheology and behavior (e.g., Walcott, 1972; Clark
et al., 1978; Mo¨rner, 1980; Lambeck et al., 1996; Lambeck, 2001; Peltier,
1998, 2004; Brooks et al., 2007; Bradley et al., 2011). The models have

been used widely in paleoshoreline and coastal map reconstructions, primarily for the Holocene. However, many of these approaches use sea-level
index points in model validation and in tuning the appropriate rheological
characteristics. Consequently, the models also suffer to different degrees
from built-in error problems, in terms of the of height and time accuracy of
the coastal reconstructions (Devoy, 1991; Zerbini et al., 2000). Issues of
regional variations inherent in Earth rheology and crustal response time to
uncertainties in the calculations of former ice and water loadings also exist
(i.e., ice volumes and rates of change) and may be difficult to constrain
(e.g., Peltier, 1998; Lambeck, 2001; Ehlers and Gibbard, 2004; Ehlers
et al., 2011; Benn and Evans, 2013; BRITICE-CHRONO, 2014). The primary terms of these geophysical models for past SLCs can be expressed
most simply as:
Dzð4; l : tÞ ¼ zð4; l : tÞ Àzð4; l : t0 Þ
¼ Dze ðtÞ þ Dzr ð4; l : tÞ þ Dzi ð4; l : tÞ þ DzW ð4; l : tÞ

(8.4)

where the positions of past sea levels at latitude 4, longitude l, and time t
are calculated as the sum of four terms in relation to present day sea level
(t0), though the operation of vertical tectonics is not accounted for. The first
term in Eqn (8.4) is for the equivalent SLR, defined as the change in ocean
volume/(divided by) ocean surface area; the second term is the SLC for the
gravitational potential from changed ice and water loadings; the third is the
rebound from changes in ice loadings, and includes further gravity effects,
and the last term is the Earth response to sea floor loadings by melt water
addition or losses as water uptake to ice masses. Where a number of ice
sheets contributes to the SLC in time, as is likely, Eqn (8.4) can be
expressed as:
Dz ¼ Dze þ

N n

X






o
ðnÞ
ðnÞ
DzðnÞ
4;
l
:
t
þ
Dz
4;
l
:
t
þ
Dz
4;
l
:
t
(8.5)
W
r

l

n¼1

where Dze is the equivalent SLC for the combined ice sheet contributions.
The other terms represent the components of the position of the site at which
SLCs are being calculated relative to a particular ice mass (from Lambeck
(1991)).


Chapter j 8

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221

In spite of the problems in defining time-based changes in the model terms,
these geophysical models have been essential in helping resolve crustal responses to ice, water, and sediment loadings and redistributions during
glacialedeglacial hemicycles as part of SLC (e.g., IGCP projects 274, 367,
437 and Devoy, 1991). These form the isostatically driven uplift and subsidence components in SLC, including the operation and significance of
different regional signatures of ice-margin and far-field crustal forebulge
movements (Sabadini et al., 1991; Devoy, 1991, 1995; Peltier, 1998; Lambeck,
1991, 1995, 2001; Lambeck et al., 1996, 1998; Carter, 1992; Brooks et al.,
2007; Edwards and Brooks, 2008; Bradley et al., 2011). The different crustal
loadings can also affect changes in planetary rotation and “wobble,” with
feedbacks to the geoid and potentially to the timings and scales of SLCs (Clark
and Lingle, 1977; Clark and Primus, 1987; Peltier, 1998; Milne and Mitrovica,
1998; Mitrovica et al., 2009, 2011). Importantly, the geophysical models
established the key first order SLC “finger-prints” with ice-melt patterns for
world coasts at regional scales (Figure 8.9) (e.g., Clark et al., 1978). Linked to

this work comes an increased understanding and quantification of changes in
the geoid. Model results provide a valuable basis and analogs for simulating
the likely impacts of Greenland and Antarctic ice deglaciation with future
climate warming (Milne and Mitrovica, 1998; Mitrovica et al., 2009, 2011;
Milne et al., 2009; Church et al., 2014). Changes in the mass of the present
major ice sheets will cause gravitational “relaxation/collapse” of the sea surfaces at their margins and related far-field consequences in SLR. For the North
Atlantic, this mechanism alone may result in as much as a 0.19 m SLR (Raper,
1993, 2000; IPCC WGI, 2007, 2014).

8.4.2 Meso-scale to Micro-scale Changes
Meso-scale SLCs are determined primarily by factors of tectonic and glacial
eustasy (Figure 8.7 and Table 8.1). At this level of resolution, “oscillatory,” as
cyclical and quasiharmonic type sea-level movements, are evident and to be
expected, particularly as correlated with the Quaternary glacial and interglacial
stages (Figure 8.1(b)) (e.g., Devoy, 2005; Roe et al., 2009; Murray-Wallace and
Woodroffe, 2014). Superimposed on these glacioeustatic trends are shorter-term
sea-level movements, driven by meteorological and other dynamic sea surface
factors, together with the feedbacks from longer-term controls (Figure 8.4(b)).
These are rarely preserved clearly in the geological proxy records of SLC,
except perhaps as sea-level pattern variabilities, or “pattern noise” (e.g., Devoy,
1982; Allen, 1997, 2003; Delaney et al., 2012). At these scales, the question
arises as to whether these shorter term drivers really do result also in higher
resolution “oscillatory” patterns in sea-level behavior. This question was equally
part of the causes, form and nature of Holocene sea-levels “debate” in the early
IGCP Sea Levels Projects 61 and 274 (Figure 8.5) (Bloom, 1977; Greensmith
and Tooley, 1982; Tooley, 1993; Pirazzoli, 1996; Zerbini et al., 2000; Church