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The above example illustrates the hazards of trying to estimate

eustatic amplitudes from oxygen isotopic records. Though conti-

nental margin stratigraphic records are complicated by changes

in subsidence, it is possible to evaluate subsidence histories using

a method called backstripping. This method progressively

removes the effects of compaction, sediment loading, and ther-

mal subsidence; the residual provides an estimate of eustatic

sea level (e.g., Kominz et al., 1998; Watts and Steckler, 1979).

Backstripping requires estimates of age, sediment type (for

decompaction), and paleowater depth; the latter is obtained from

benthic foraminifera and lithologic data and is the greatest source

of uncertainty. Kominz et al. (1998) and Van Sickel et al.

( 2004) backstripped individual onshore New Jersey boreholes

( “1-dimensional” backstripping) and found amplitudes of 20–

40 m. Two-dimensional backstripping (backstripping an entire

profile using several boreholes) of ten latest Eocene to earliest

Miocene sequences provided estimates of 20–60 m eustatic

lowerings (Kominz and Pekar, 2001). Both one and two dimen-

sional backstripping yielded Oligocene-Miocene eustatic esti-

mates that are lower than those published by the EPR group

(e.g., Haq et al., 1987) by a factor of two or more (Kominz and

Pekar, 2001; Van Sickel et al., 2004).

Studies of sea-level change must be global in scope and

ODP has drilled on the Australian margin (Legs 133, 182,

and 194) and the Bahamas (Leg 166). Despite the fundamental

difference in sedimentation style between carbonate and silici-

clastic margins, ODP drilling documented similar-Miocene

unconformities in these diverse settings ( Figure S6). Drilling

in the Bahamas showed that the flanks of a carbonate bank

environment develop sequences that are remarkably similar in

character to those of siliciclastic margins (Eberli et al., 1997).

Though tectonic effects (e.g., rapid subsidence) influence much

of the northeast Australian region (Davies et al., 1993), conse-

quences of global sea-level change on the evolution of the Mar-

ion Plateau and the Great Barrier Reef were evaluated. Drilling

on the Australian margin by Leg 194 (Isern et al., 2002) pro-

vided an estimate of a major late middle Miocene eustatic low-

ering of 33  12 m based on Airy (point) loading to 85  30

assuming rigid crust (flexural loading). These wide ranges

again illustrate the difficulty in estimating global sea-level

amplitudes.

Cause s of sea level change

Drilling on the New Jersey, Bahamas, and Australian margins

have firmly demonstrated that ice-volume changes have been

one of the primary controls on sea-level change during the Ice-

house world of the past 34 Myr. It is not surprising that this is

true because ice-volume changes are the only known mechan-

ism for causing the large ( >10s m), rapid changes in sea level

reported for this time interval. The close correspondence

between the oxygen isotopic and continental margin records

is testament to this linkage ( Figure S6). However, what about

the vast amounts of geologic time in which the Earth was pre-

sumably ice-free?

Sea level during the Cretaceous (144–65 Ma) was much

higher than at present because of long-term changes in sea

level. These long-term (10 –100 Myr scale) sea-level changes

have been largely controlled by tectonics. For example, during

the warm interval of the middle Cretaceous (ca. 80–90 Ma),

global sea level was between 100 (Miller et al., 2005) and

170 m (Müller et al., 2008) higher than today. However, such

high sea levels cannot be explained by a warmer planet and

the absence of ice sheets: even if all of the modern ice sheets

were melted (an “ice-free Greenhouse world”), then sea-level

would be only 73 m higher than today. Although the Cretac-

eous oceans were 10–12



C warmer than today, thermal expan-

sion could only account for another 10 –13 m of rise. The

generally accepted explanation for the high mid-Cretaceous sea

levels is high sea-floor spreading rates; higher rates mean hotter

crust, which expands, displacing seawater and causing long-

term flooding of the continents (Hays and Pitman, 1973;

Kominz, 1984). This pulse of mid-Cretaceous high seafloor

spreading has been recently disputed (Rowley, 2002), though

various stratigraphic data unequivocally show that relative

to the continents, sea level was 100–250 m higher, based on

backstripping in the former case (Watts and Steckler, 1979)

and continental flooding in the latter (Sahagian et al., 1996).

If not caused by high-sea floor spreading rates, the long-term

flooding and subsequent retreat of the sea over the past 180

Myr could have been the result of the breakup of the supercon-

tinent Pangea, beginning at 180 Ma, which may have led to

overall subsidence of continents relative to the oceans during

the Jurassic to mid-Cretaceous.

Even more puzzling than the long-term record is the obser-

vation that large (10s to 100 m), rapid (less than 1 Myr) sea-

level changes occurred during the Triassic to middle Eocene

(ca. 250–50 Ma), a time considered to be an ice-free Green-

house (Haq et al., 1987; Hallam, 1992). Ice-volume changes

(glacioeustasy) are the only known mechanism for producing

large, rapid sea-level change (Pitman and Golovchenko,

1983). Although it has been believed in general that there were

no significant ice sheets prior to the middle Eocene, Haq et al.

(1987) delineated numerous Cretaceous to early Eocene

sequence boundaries with associated large (>50 m), rapid

(<1 Myr) sea-level lowerings. There are four explanations for

this apparent paradox (Browning et al., 1996): (a) the Cretac-

eous to early Eocene sequences summarized by Haq et al.

(1987) were restricted to local basin(s) and do not reflect eus-

tasy (this is unlikely considering that many have been widely

recognized; e.g., Aubry, 1985; Mancini and Tew, 1991, 1995;

Olsson, 1991); ( b) the sequences were controlled by low-

amplitude sea-level changes (e.g., 10 m of lowering in 1 Myr

can be explained by numerous mechanisms; Donovan and

Jones, 1979); (c) mechanisms of sea-level change are not fully

understood; and (d) there were ice sheets throughout much of

the Cretaceous to early Eocene (e.g., Stoll and Schrag, 1996).

Drilling in New Jersey has recently provided a record of

sea-level changes during the warm Late Cretaceous. These stu-

dies firmly document that large (>25 m), rapid (1 Myr)

eustatic variations occurred in the Late Cretaceous, hinting at

the presence of ice sheets in this supposedly ice-free Green-

house world (Figure S7). Continuous coring recovered 11–14

Upper Cretaceous sequences at Bass River and Ancora, New

Jersey that were dated by integrating Sr-isotopic and biostrati-

graphic data (age resolution 0.5 Myr) and subsequently back-

stripped taking into account sediment loading, compaction,

paleodepth, and thermo-flexural subsidence, providing a new

sea-level estimate. The timing of Late Cretaceous sequence

boundaries in New Jersey is similar to sea-level lowerings

found by EPR (Haq et al., 1987), NW European sections

(Hancock, 1993), and Russian sections (Sahagian et al.,

1996), indicating a global cause (Figure S7

). However, New

Jersey and Russian Platform eustatic estimates are half of the

EPR amplitude; the EPR record also differs in shape from the

backstripped estimates. The use of the EPR record for the Late

884 SEA LEVEL CHANGE, LAST 250 MILLION YEARS

Figure S7 Comparison of Late Cretaceous deep-sea benthic foraminiferal d

O records, New Jersey composite sequences, relative sea level from NW Europe (Hancock, 1993), backstripped

Russian platform record (Sahagian et al., 1996), the EPR eustatic estimate (Haq et al., 1987), backstripped R2 eustatic estimates for Bass River (black) and Ancora (gray) and our best

estimate of eustatic changes (heavy black line constrained by data, dashed inferred). Pink arrows indicate positive d

O inflections (inferred cooling and /or ice volume increases) (after

Miller et al., 2004).

SEA LEVEL CHANGE, LAST 250 MILLION YEARS 885

Cretaceous should be abandoned; the New Jersey backstripped

records provide the best substitute at this time.

The firm documentation of large (>25 m), rapid (<Myr) sea-

level changes (Figure S7) indicates some glacioeustatic control

during the Late Cretaceous. The apparent inconsistency with evi-

dence for warm high latitudes during the Late Cretaceous can be

explained by ephemeral (i.e., lasting 1 Myr) ice sheets (pre-

sumably in East Antarctica (EAIS)) with volumes approaching

one-third to one-half of the modern EAIS (Miller et al., 1999)

and intervening warm, ice-free intervals (Miller et al., 2003).

Oxygen isotopic comparisons with Late Cretaceous

sequence boundaries have not attained the resolution needed

to link the two datasets unequivocally, as has been done for

the past 42 Myr (Miller et al., 1998a). Nevertheless, compari-

sons between Late Cretaceous sequence stratigraphy and d

records are intriguing (Figure S7), further suggesting small

ice sheets in this alleged Greenhouse World: (a) a major mid-

Cenomanian sequence boundary (hiatus ca. 96–97 Myr ago)

(see also Gale et al., 2002) correlates with a major (>1%)

O increase; (b) two minor d

O increases spanning the

Cenomanian/Turonian boundary may correlate with sequence

boundaries; (c) a mid-Turonian sea-level lowering (91.5–

92 Ma) may correlate with a major increase in benthic forami-

niferal d

O values (1.0%), although additional data are

needed to determine the precise timing of the increase (Fig-

ure S7). Several other Coniacian-Campanian d

O increa-

ses (dashed arrows, Figure S7) may be related to sequence

boundaries, but the data are too sparse to provide a firm

correlation.

The EPR sea-level records (Vail et al., 1977; Haq et al.,

1987) are major achievements for the Earth sciences. However,

it is not surprising after 15–25 years, that we conclude that

these previous sea-level estimates are not entirely correct in

their shape and amplitudes. Methods for extracting sea level

from stratigraphic data have advanced during this period and

the scientific community has begun to document amplitudes

of sea-level change with increased precision. Despite their lim-

itations, the EPR curves stand the test of time as an approxi-

mate chronology (1 Myr) of eustatic lowerings. Studies on

the New Jersey and other continental margins have generally

confirmed the number and approximate ages of sea-level low-

erings of Haq et al. (1987) for the past 100 Myr (Figures S6,

S7). In the absence of other datasets for the Triassic-Early Cre-

taceous, the Haq et al. (1987) record still provides the best esti-

mate for the timing of eustatic lowerings. Although the timing

can be estimated, the amplitudes of sea-level change during the

Triassic-Early Cretaceous remain poorly constrained and the

causal mechanisms for the large, rapid sea-level changes for

this interval can only be explained by glacioeustasy. Is our

understanding of causal mechanisms for global sea-level

change flawed, or is our interpretation of a largely ice-free pla-

net prior to 42 Ma incorrect?

ODP drilling has determined ages of sequence boundaries

on continental margins to better than 0.5 Myr and provided

a chronology of eustatic lowering for the past 42 Myr

(Figure S6). ODP has also validated the transect approach of

drilling passive continental margins and carbonate platforms

(onshore, shelf, slope). Additional integrated geological and

geophysical studies of transects of margins are needed to

address the history magnitude, and mechanisms of global sea

level change over the past 250 Myr.

Kenneth G. Miller

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Cross-references

Cenozoic Climate Change

Foraminifera

Glacial Eustasy

Greenhouse (Warm) Climates

Icehouse (Cold) Climates

Ocean Drilling Program (ODP)

Oxygen Isotopes

Sea Level Change, Post-Glacial

Sea Level Change, Quaternary

Sea Level Rise, in Encyclopedia of World Climatology

SEA LEVEL CHANGE, POST-GLACIAL

The late Quaternary glacial period was characterized by mas-

sive ice sheets that covered parts of North America, northern

Europe, and several other high latitude and high altitude

regions. This volume of ice was sufficient to lower global sea

level by over 100 m. As the ice sheets began to melt and retreat

over 20,000 years ago, sea level rose rapidly. However, the

time history and magnitude of sea level rise were widely

debated, as data from radiocarbon-dated paleosea level indi-

cators began to accumulate in the 1950s and 1960s. A major

point of discussion at that time centered on whether or not

the post-glacial marine transgression followed a smooth versus

a fluctuating curve that reflected world-wide sea level changes

(e.g., Shepard and Suess, 1956; Fairbridge, 1961, 1976). Sup-

port for both points of view could be found as recently as

1991 in a compilation of Holocene sea level curves (Pirazzoli,

1991). The acquisition of more accurately dated sea level data

and development of a theoretical framework to account for

interactions among various geophysical and climatological pro-

cesses has led to a much clearer understanding of post-glacial

sea level rise (SLR) within the last 30 years. The deglacia-

tion was accompanied by a global increase in sea level

of 120 m. Post-glacial sea level did not rise smoothly, but

instead jumped in several sharp and rapid spikes, termed “melt-

water pulses” (Fairbanks, 1989). However, the exact number,

timing, and magnitude of these major pulses, and their relation

to paleoclimate events have yet to be resolved. The extent and

significance of minor Holocene sea level oscillations also needs

further investigation.

Processes affecting sea level

Sea level integrates the effects of diverse geological and clima-

tological phenomena that operate over a wide range of spatial

and temporal scales. The dominant processes that have affected

post-glacial sea level involve changes in ocean water volume

caused by melting of ice sheets (glacial eustasy) and changes

in ocean basin volume due to re-adjustment of the lithosphere

in response to removal of the ice load (glacial isostasy) and

glacial meltwater loading of ocean basins (hydro-isostasy).

While of lesser importance during this period, major

factors controlling global sea level over geologic timescales

(10

–10

years) include changes in volume of ocean basins

due to variations in rates of seafloor spreading and subsidence

of cooler, oceanic crust away from spreading centers (tectono-

eustasy). Changes in ocean water density, caused by changes

in temperature or salinity (steric changes) become more sig-

nificant contributors to sea level on much shorter timescales

(10–10

years). A twentieth century global sea level rise

of 1.7–1.8 mm yr

1

(Church et al., 2004; Holgate and Wood-

worth, 2004) appears to be related to recent climatic warm-

ing, which has been attributed to anthropogenic additions of

SEA LEVEL CHANGE, POST-GLACIAL 887

greenhouse gases into the atmosphere (e.g., Church et al., 2001,

IPCC, 2007). Around 0.4 mm yr

1

,or22–24%, derives from

penetration of this atmospheric heat signal into the oceans and

the consequent thermal expansion (Antonov et al., 2005); the

balance is due to melting of mountain glaciers and ice sheets.

The relative magnitudes of the processes influencing sea

level vary from place to place, thus tending to obscure any

globally coherent sea level signal. Therefore, the derivation of

a eustatic sea level curve requires an ability to decouple the

effects of land motions (uplift/subsidence) from processes

related to changes in continental ice extent or ocean water

volume. Away from convergent plate margins, the main cause

of long-wavelength elevation differences during the Holocene

is related to changes in ice and water loading. Geophysical

models of glacial isostatic adjustment (GIA) account for time-

varying gravitational interactions among ice sheets, land, and

ocean (Peltier, 2001, 2002, 2004; Lambeck et al., 2002; Mitrovica

and Milne, 2002). GIA models calculate the observed sea

level change over time at any given location as a function

of the volumetric addition of glacial meltwater and contri-

butions from both glacial isostasy and hydro-isostasy. Input

parameters include values for maximum ice sheet thickness

and extent, a history of ice sheet retreat, and mantle viscosity

profiles. The GIA models are “tuned” by fitting the results to

radiocarbon-dated sea level records (see Sea level indicators)

and adjusting the mantle viscosity (Peltier, 2001, 2002, 2004;

Lambeck et al., 2002). GIA models predict several zones that

display similar sea level patterns (Figures S8–S9). In areas

formerly covered by ice that are isostatically rebounding,

sea level falls continuously (Figure S8a). Areas marginal to

ice sheets that had been upwarped during glaciation (the

“peripheral forebulge”) subside following removal of the ice

load (Figure S8b). Sea level rises continuously in this zone.

Transitional zones between glacial rebound and collapse

experience a more complex history of sea level fall followed

by a rise (Figure S8c). On mid-ocean islands far removed

from former ice sheets (the “far field”), after attaining its

maximal height, sea level drops once deglaciation has ceased

in the mid-Holocene due to “siphoning” of water toward

subsiding forebulges (Figure S9a). Finally, far field continen-

tal coastlines show regions of sea-level fall, with rising ocean

levels further offshore (due to “continental levering,” see

below; Mitrovica and Milne, 2002; Figure S9b). Thus, the

observed sea level trend at any given locality depends on

the interplay between the effects of glacial eustasy and glacial

isostasy. The composite eustatic sea level curve (Figure S10)

is based on careful selection of sites where glacial isostatic

effects are minimal.

Sea level rise after the last glaciation

A composite curve of post-glacial sea level (Figure S10a) has

been derived from radiocarbon- and

230

Th/

234

U-dated sea level

indicators, including corals, plant material, foraminifera, and

mollusks from a number of regions with minimal glacial

isostatic influence. Figure S10b shows the late Quaternary –

Holocene sea level curve calculated from the ICE-5G (VM2)

GIA model together with Th-U-dated corals from Barbados

(after Peltier and Fairbanks, 2006). Note that the coral data

suggest a minimum sea level close to 26,000 years before

present, significantly older than the commonly accepted date

of 20,000

BP (Table S1).

The post-glacial eustatic sea level rise, although continuous,

was punctuated by several periods of more rapid transgression,

or meltwater pulses. While the exact number and timing of

these events has been debated, recent studies support the pro-

position that rates of sea level rise have varied over this period

(see Table S1). Sea level stood close to 120 m below present at

the Last Glacial Maximum (LGM) (Table S1). An initial pulse

of rapid SLR may have started  19,000 years ago, at which

time ocean levels rose 10–15 m within 500 year, herein

referred to as “meltwater pulse 1A

” (Yokoyama et al., 2000;

Clark et al., 2004). However, this meltwater pulse does not

appear in all records (e.g., Hanebuth et al., 2000; Peltier and

Fairbanks, 2006), although other paleoclimate proxies suggest

a sudden influx of meltwater with accompanying changes in

ocean circulation at around this time (Clark et al., 2004). Other

Figure S8 Schematic near field Holocene sea level curves from

different zones of glacial isostatic adjustment. The vertical scale

represents relative sea level change in meters (negative values are

below present sea level, positive values above). These generalized

curves do not represent specific localities. (a) Former ice-covered

regions–continuous sea level fall (land emergence); e.g., Scandinavia;

Canada. (b) Collapsed peripheral bulge–continuous sea level rise (land

subsidence); e.g., the Netherlands, Belgium, southern England, East

Coast, USA, south of New York. ( c) Transitional zone–early Holocene sea

level fall followed by sea level rise; e.g., Maine, Canadian Maritimes.

888 SEA LEVEL CHANGE, POST-GLACIAL

data and theoretical models suggest instead that the post-glacial

marine transgression may have begun as early as 26,000 years

ago (Figure S10b; Peltier and Fairbanks, 2006).

Another accelerated phase of SLR occurred between 14,600

and 13,500 y

BP (meltwater pulse 1A; Table S1), when sea level

increased by some 16 –24 m (Fairbanks, 1989; Hanebuth,

2000). Depending on its exact timing, this event has been

variously associated with the first cold phase during the

Bølling-Allerød interstadial or the Older Dryas (Bard et al.,

1996), the Bølling warming (Kienast et al, 2003; Peltier,

2005), or the end of the Bølling (Hanebuth et al., 2000; Liu and

Milliman, 2004). Kienast et al. (2003), for example, connect a

sharp drop in terrigenous organic matter delivered to the South

China Sea, a rapid marine incursion, and synchronous rise in

ocean temperatures between 14,600 and 14,300 y

BP with

the Bølling warming.

The rate of sea level rise slowed between 14,000 and 12,000

BP during the Younger Dryas cold period. It was succeeded

by a faster phase (meltwater pulse 1B) between 11,500 and

11,000 years ago. Fairbanks (1989) initially estimated that sea

level jumped by 28 m during meltwater pulse 1B, but subse-

quent studies suggest a significantly more modest increase, if at

all (Bard et al., 1996; Lambeck et al., 2002).

A fourth surge of rapid sea level rise, herein referred to as

“meltwater pulse 1C” has been proposed, but is less firmly

established. This meltwater pulse was first identified by a hia-

tus in coral growth in the Caribbean at 7,600 y

BP (Blanchon

and Shaw, 1995), but neither Bard et al. (1996) nor Lambeck

et al. (2002) detected any noticeable increase in sea level at

that time. Nevertheless, findings from diverse locations hint

at a possible sudden sea level jump around 8,000 y

BP. Data

from Chesapeake Bay show a relatively stable period of sea

level between 9,200 and 7,900 y

BP, followed by a short and

rapid spurt of SLR between 7,900 and 7,600 y

BP (Cronin et al.,

2002). Dates on sediment cores from the Mississippi River delta

constrain the episode of abrupt SLR to between 8.5  0.3

Figure S9 Schematic far field Holocene sea level curves, showing

mid-Holocene land emergence. These generalized curves do not

represent specific localities. (a) Equatorial ocean islands; e.g., Cocos

(Keeling) Islands, Indian Ocean; Fiji Islands, Pacific Ocean. (b)

Continental coastlines; e.g., eastern Australia, New Zealand, Brazil,

West Africa.

Figure S10 (a) Generalized post-glacial eustatic sea level curve

(based on Fairbanks, 1989; Bard et al., 1990, 1996; Hanebuth et al., 2000;

Yokoyama et al., 2000). (b) Eustatic sea level curve for the last 32 Kyr

corrected for glacial isostatic adjustments using the ICE-5G (VM2)

model (smooth curve). Coral data from Barbados are shown by error

bars. Horizontal bars indicate depth of coral sample. Vertical bars

represent the range with respect to sea level in which the coral species

could live. The shortest bars (5 m) correspond to Acropora palmata; the

intermediate bars (20 m) to Monastera annularis; longer bars to either

Asteroides or Diploria (modified from Peltier and Fairbanks, 2006).

The latter species provide only a lower limit of relative sea level.

SEA LEVEL CHANGE, POST-GLACIAL 889

and 8.1  0.1 kyBP (Törnqvist et al., 2004). Stratigraphic data

from the subaqueous delta of the Yellow River, China (Liu

et al., 2004), coastal Lancashire, England and Limfjord, north-

western Denmark (Tooley, 1989) also point to a more rapid

marine transgression at around 8,000 y

BP. Meltwater pulse 1C

has been linked to the 8,200-y

BP cold event (Cronin et al.,

2002; Törnqvist et al., 2004), which resulted from the cata-

strophic drainage of proglacial Lakes Agassiz and Ojibway

(Barber et al., 1999; Clarke et al., 2004).

Sources of meltwater

Among the proposed episodes of accelerated post-glacial SLR,

evidence for meltwater pulse 1A is probably the strongest.

This event occurred during a period of maximal ice sheet

retreat. Yet, the origin of the meltwater has remained elusive.

Meltwater from the breakup of the Laurentide Ice Sheet has

generally been assumed to be the main contributor to the rapid

SLR, mainly because of its large extent. Geophysical models

suggest that sea level changes accompanying rapid melting

(or growth) of an ice mass should lead to a spatially non-

uniform distribution of sea level change, not only because of

ice loading, but also due to gravitational attraction between

the ice sheet and ocean water (Clark et al., 2002). As the ice

sheet shrinks, its gravitational attraction on ocean water will

diminish and water will flow away from the edge of the ice.

Therefore, different sources of meltwater (e.g., the Laurentide,

Antarctic, or Barents Sea Ice Sheets) should give rise to differ-

ent geographic patterns of sea level rise. The closest match

between observed sea level records from various parts of the

world and model predictions was found when Antarctica made

a substantial contribution to meltwater pulse 1A and, in addi-

tion, a relatively high value for lower mantle viscosity was

assumed (Clark et al., 2002; Bassett et al., 2005). However,

from a new reconstruction of ice sheet retreat (Peltier, 2004)

based on a glacial systems model that was calibrated by

C-dated relative sea level records, geodetic measurements,

and geomorphological data, Tarasov and Peltier (2005) deter-

mined instead that this meltwater pulse derived largely from

North American sources. Furthermore, varved sediments con-

taining diatomaceous ooze from fiords in East Antarctica

record the onset of ice sheet recession there at 11,500–

10,500 y

BP, long after meltwater pulse 1A (Leventer et al.,

2006). Thus, Antarctica was not a likely contributor to this

meltwater surge. In addition, comparison of relative sea level

data from Barbados and the Sunda Shelf, off Indonesia, reveals

differences in magnitude of the meltwater pulse 1A at these

two far field locations that are much too small to attribute to

a specific geographical source (Peltier, 2005).

The signals for meltwater pulse 1B and meltwater pulse 1C

are less distinct than that of meltwater pulse 1A. Meltwater

from glacial Lake Agassiz has long been thought to have

triggered thermohaline circulation changes that led to the

Younger Dryas cold period (12,800–11,600 y

BP), just prior to

meltwater pulse 1B. However, new dates on moraine deposits

show that regional deglaciation occurred nearly 1,000 radiocar-

bon years later than the Younger Dryas (Lowell et al., 2005),

and hence Lake Agassiz was not the likely source of meltwater

pulse 1B. Instead, glacial models show a period of increased

meltwater discharge between 12,900 and 12,800 y

BP, originat-

ing from outflow into the Arctic Ocean, the Fram Strait, and

ultimately the eastern North Atlantic (Tarasov and Peltier,

2005). They suggest that this event may have triggered the

Younger Dryas cold reversal, although it has not left a dis-

cernable sea level record. On the other hand, the timing of

deglaciation in eastern Antarctica roughly coincides with the

onset of meltwater pulse 1B (Leventer et al., 2006).

As noted above, meltwater pulse 1C can be identified in

a number of widely separated geologic records. This melt-

water pulse has been attributed to the catastrophic drainage of

proglacial Lakes Agassiz and Ojibway around 8,400 year ago

(Barber et al., 1999; Clarke, et al., 2004 ). The volume released

by the sudden outburst flood was 10

within a few years,

or even less. Although this enormous influx of freshwater to

the world’s oceans probably triggered the 8.2 kyr cold event

(see Alley and Ágústsdóttir, 2005; the 8,200-y

BP event, this

volume), the effects on global sea level appear to have been

relatively modest. The resulting increase in sea level (assuming

an even distribution of this volume over the world’s oceans)

was only on the order of 1 m or less (Barber et al., 1999; Clarke

et al., 2004; Törnqvist et al., 2004). Nevertheless, this minor

change in sea level left an imprint in the stratigraphic record.

Middle to late Holocene sea level

By the mid-Holocene, 6,000–5,000 yBP, glacial melting had

essentially ceased. However, glacial isostatic and hydro-isostatic

adjustments persisted, although at decreasing amplitudes over

Table S1 A review of major events associated with post-glacial sea level

rise

Sea level at LGM

–121  5 m (Barbados; Fairbanks, 1989)

–118 m (Barbados; Bard et al., 1990)

–115  1 m (Sunda Shelf; Hanebuth et al., 2000)

–125 to –121 m; –135 to –130 m (Bonaparte Gulf, Australia; corrected for

glacial isostasy; Yokoyama et al., 2000)

–123 to –119 m (Barbados; Peltier and Fairbanks, 2006)

Onset of post-glacial sea level rise

18,000 y

BP (Barbados; Fairbanks, 1989)

>19,000–18,000 y

BP (Barbados; Bard et al., 1990)

21,000–19,000 y

BP (Sunda Shelf; Hanebuth et al., 2000)

22,000–19,000 y

BP (Bonaparte Gulf; Yokoyama et al., 2000)

20,000 y

BP (Irish Sea; Clark et al., 2004)

26,000 y

BP (Barbados; Peltier and Fairbanks, 2006)

Meltwater pulse 1A

19,000 yBP; rate of SLR 33.3–50 mm yr

1

(Yokoyama et al., 2000)

19,120  180 y

BP; rate of SLR 20 mm yr

1

(Clark et al., 2004)

Meltwater pulse 1A

13,000 y

BP; rate of SLR, 39 mm yr

1

(Barbados; Fairbanks, 1989)

13,500 y

BP;” 37 mm yr

1

(Barbados; Bard et al., 1990)

14,000 y

BP (Tahiti; Bard et al., 1996)

14,600–14,300 y

BP; rate of SLR, 53.3 mm yr

1

(Sunda Shelf; Hanebuth

et al., 2000; South China Sea, Kienast et al., 2003)

14,300–14,000 y

BP; rate of SLR, 65  10 mm yr

1

(Liu and Milliman,

2004)

Meltwater pulse 1B

10,500 y

BP; rate of SLR, 26.4 mm yr

1

(Barbados; Fairbanks, 1989)

11,000 y

BP;”25 mm yr

1

(Barbados; Bard et al., 1990)

11,300 y

BP (Tahiti; Bard et al., 1996)

11,500–11,200 y

BP; rate of SLR 40 mm yr

1

(Liu and Milliman, 2004)

Meltwater pulse 1C

7,600 y

BP (Caribbean Sea; Blanchon and Shaw, 1995)

8,200 y

BP (Mississippi Delta; Törnqvist et al., 2004)

8,000–7,000 y

BP (Yellow River Delta, China; their “meltwater pulse 1D,”

Liu et al., 2004)

7,900–7,600 y

BP; rate of SLR  40 mm yr

1

(Chesapeake Bay; Cronin

et al., 2002)

Other periods of rapid sea level rise

9,800–9,000 y

BP; rate of SLR  45 mm yr

1

(their “meltwater pulse

1C”; Liu et al., 2004)

890 SEA LEVEL CHANGE, POST-GLACIAL

time. Therefore, sea level continued to lower in formerly gla-

ciated regions and rise in areas of the collapsing forebulge.

At many equatorial far field ocean island sites, sea level stood

several meters higher than present during the mid-Holocene,

6,000–4,000 y

BP, and has been falling slowly ever since

(e.g., Dickinson, 2001). This phenomenon has been explained

to be a consequence of lithospheric responses due to changes

in ice and water loading (Mitrovica and Peltier, 1991; Mitrovica

and Milne, 2002). Water is “siphoned” away from the central

equatorial ocean basins into nearshore depressions left by the

collapse of peripheral forebulges. In addition, the oceanward

tilting of far field continental coastlines (“continental levering”)

because of hydro-isostasy also leads to migration of ocean

water away from central equatorial ocean basins toward the

continents (Figure S11). Mid-Holocene sea level high-stands

have also been interpreted as eustatic in origin, likely related

to minor climate fluctuations. Support for this view comes

from well-dated sessile intertidal tubeworms from Western

Australia, which show several episodes of higher than present

sea level, ranging between 6,000 and 2,000 y

BP (Baker et al,

2005). These events are also detected in Antarctic ice cores

and several paleosea level proxies elsewhere.

The sea level record of the last 4,000–3,000 years is less-

well constrained, in spite of its importance as a baseline for

understanding contemporary sea level trends. In part, this arises

from a paucity in accurately-dated high resolution sea level

proxies, as well as difficulty in extracting a global climatic signal

from low-amplitude oscillations. Minor sea level fluctuations of

1–2 m have been reported in a number of records (e.g., Finkl,

1995; Scott et al., 1995a,b; Varekamp and Thomas, 1998). These

probably record local to regional climate-related processes, such

as variations in river discharge, storm deposition, or coastal sedi-

mentation, rather than true eustatic changes in sea level.

Recent sea level rise

Twentieth century mean global sea level, determined from

tide gauges in coastal harbors, has been increasing by 1.7–

1.8 mm yr

1

(Figure S12; Church et al., 2004; Holgate and

Woodworth, 2004; Church and White, 2006), apparently

related to the recent climatic warming trend (Church et al.,

2001, IPCC, 2007). An even higher sea level trend of 2.8

mm yr

1

has been obtained from the TOPEX/POSEIDON

satellite altimeter, since 1993 – a period strongly influenced

by major El Niño-Southern Oscillation (ENSO) events, when

sea surface temperatures and hence thermal expansion were

higher than normal over large portions of the oceans. Neverthe-

less, analysis of tide-gauge data over a longer period, from

1870 to 2004, shows a twentieth century acceleration in global

sea level of 0.013  0.006 mm yr

2

(Church and White, 2006).

The twentieth century sea level trends are substantially

higher that those of the last few thousand years, derived from

geological data (Shennan and Woodworth, 1992; Gornitz,

1995; Varekamp and Thomas, 1998; Donnelly et al., 2004).

The exact timing for the onset of the current accelerated phase

is uncertain because of limited instrumental data prior to the

1850s and lack of high resolution proxy sea level data for the

last few thousand years. Although Varekamp and Thomas

(1998) claim that the acceleration began at least 200 years

ago, other studies point to a mid-late nineteenth century to early

twentieth century onset (e.g., Donnelly et al., 2004;Gehrels

et al., 2005; Church and White, 2006).

Summary and conclusions

The melting of the ice sheets at the end of the last glaciation,

beginning at least 20,000 and possibly as early as 26,000 years

ago, produced a rapid rise in sea level of some 120 m, spanning

a period of 13,000 years. The post-glacial sea level rise was

temporally non-uniform and several periods of more rapid rise

may have occurred: 19,000 y

BP (meltwater pulse 1A

);

14,600–13,500 y

BP (meltwater pulse 1A); 11,500–11,000 yBP

(meltwater pulse 1B); and 8,200–7,600 yBP (meltwater pulse

1C). However, evidence for some episodes (e.g., meltwater

Figure S11 Schematic diagram illustrating processes contributing to

mid-Holocene sea level maximum, and subsequent fall, at far field sites

(not drawn to scale). (a) Ocean “siphoning” due to flow of water from

far field oceans toward collapsed forebulge. (b) “Continental levering”

due to ocean loading by meltwater.

Figure S12 Recent global sea level rise, 1948–2002, from tide gauge

data corrected for the effects of glacial isostatic adjustments (after

Holgate and Woodworth, 2004).

SEA LEVEL CHANGE, POST-GLACIAL 891

pulse 1A) is stronger than for others (e.g., meltwater pulses

, 1C) and further research is needed to confirm dates and

magnitudes. During the last 6,000–5,000 years, eustatic sea

level changes have been relatively minor after melting of major

ice sheets essentially ceased; however, crustal adjustments to

glacial and hydro-isostasy have continued, although at dimin-

ishing rates, leading to variations in long-term sea level change

at different localities. The mid to late Holocene emergence

(i.e., apparent sea level fall) of many tropical ocean islands

and shorelines has generally been attributed to ocean “siphoning”

and continental “levering.” Evidence for low-amplitude Holocene

sea level oscillations has been reported from a number of local-

ities, but it remains to be determined whether these represent

true eustatic (i.e., global) events or regional to local variability.

A recent acceleration in the rate of sea level rise ( 1–2mm

1

) began within the last 150–100 years and appears to be related

to recent global warming (Church et al., 2001, IPCC, 2007).

Vivien Gornitz

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Cross-references

Bølling-Allerød Interstadial

Glacial Eustasy

Glacial Isostasy

Glacial Megalakes

Hypsithermal

Last Glacial Maximum

Last Glacial Termination

Radiocarbon Dating

Sea Level Change, Quaternary

Sea Level Indicators

Sea Level Rise, in Encyclopedia of World Climatology

The 8,200-Year bp Event

Uranium-Series Dating

Varved Sediments

Younger Dryas

SEA LEVEL CHANGE, QUATERNARY

Relative to the land-masses, the position of sea level has

fluctuated throughout the geologic past, owing to changes in

both the quantity of ocean water and the geometry of the ocean

basins. The total volume in the hydrosphere (water plus ice)

appears to have been constant but the fractions held in land

reservoirs – glaciers, lakes, groundwater and, in particular,

continental ice sheets – fluctuated significantly. For example,

if the present Antarctic Ice Sheet were to melt, sea level would

rise by about 55 m; at the height of the last ice age, sea level

was about 120–130 m below its present position.

The proportions of land and sea are basically determined by

the fact that continents, being composed of rock lighter than

oceanic crust, stand about 4.5 km above the ocean floor. How-

ever, owing largely to slow variations in the rates of sea floor

spreading and plate tectonics, the average depth of the ocean

basins has varied throughout geological time and shorelines

have periodically advanced and retreated across the continental

shelves. Figure S13 illustrates examples of observed sea level

change on time scales from about 10

to 10

years. On the

longer timescale, sea level changed globally with amplitudes

up to several hundred meters, largely owing to plate-tectonic

changes in ocean basin geometry (Figure S13a). During the

Quaternary Era, periodic exchanges of mass between the ice

sheets and oceans caused sea level changes of tens to over a

hundred meters in amplitude, on time scales of tens to hundreds

of thousands of years (Figure S13b).

Global changes of sea level caused by changes in the

volumes of seawater or the ocean basins are referred to as

eustatic. Superimposed on these global signals are more regio-

nal and local changes, where the relative positions of land and

sea are affected by uplift or subsidence of the coastal zone.

Observations vary substantially from site to site, even over

relatively short distances such as in northwestern Europe,

where sea level at Ångerman has fallen nearly 200 m in the

past 9,000 years, while at Andøya the level 9,000 years ago

was near its present position and in southern England was

about 35 m below the present shoreline (Figure S14).

Several interlinked factors contribute to changes in sea level.

When ice sheets melt, the resulting sea level change is spa-

tially variable because the Earth's surface deforms under the

changing ice and water load, and the gravitational potential

of the Earth-ocean-ice system also changes. The combined

deformation-gravitational effects are referred to as glacio-

hydro-isostatic contributions to sea level, and it is these effects

that cause the spatial variability illustrated in Figure S14. Iso-

static warping and subsidence also occur in sedimentary basins,

in response to the accumulation of sediments many kilometers

deep over millions of years. Furthermore, tectonic movements

drive mountain uplift, enhancing landscape denudation and

leading to rapid accumulation in sedimentary basins; they also

contribute to changes of relative sea level through coastal uplift

or subsidence.

Knowledge of the complex history of Quaternary changes of

sea level is relevant in diverse fields, including the understand-

ing of climate changes, and the determination of deep-earth

properties such as the viscosity of the Earth’s mantle. Once

comprehensive sea level models are developed, it becomes pos-

sible to test hypotheses about Quaternary migrations of flora

and fauna across shallow seas that are now covered by ocean.

Moreover, to understand future sea level rise under atmospheric

greenhouse conditions, the background “natural” signal must

be known. The success of sea level studies depends very much

on the ability to separate the different contributions – eustatic,

isostatic and tectonic – in the observational record.

Observational evidence

Direct evidence for Quaternary sea level changes occurs mostly

in the form of sediments and biohermal reefs that were formed

in coastal and nearshore situations. Upper Quaternary sea

level changes are usually pieced together from raised or sub-

merged shoreline features, including shallow-water coral reefs

(Figure S15). Further back in time, relative sea level changes can

be deduced from sedimentary basins, in which eustatic variations

are registered in cyclic sediments (cyclothems) that show alternat-

ing transgressions and regressions of the sea (Figure S16). Using

evidence from seismic, drill-hole and outcrop data, the locus of

coastal-zone sedimentation can be traced throughout a given basin

sequence, allowing the course of sea level changes relative to the

basin to be identified (Hallam, 1992).

A relative sea level curve for a given area can be established

from age-height data for a series of ancient shorelines, or other

indicator deposits such as shallow marine sediments for which

SEA LEVEL CHANGE, QUATERNARY 893