Odekon M. Encyclopedia of paleoclimatology and ancient environments

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wind system, as today, was prevalent. As the mountains rose

progressively through the Late Cretaceous and especially the

Paleogene, a more variable wind system developed. Updrafts

carried moisture to heights where they produced rain. Progres-

sive runoff into river systems formed a freshwater cap at the

riverine mouths. On the western side of the Seaway, large del-

tas formed from Coniacian-Maastrichtian time and into the

Paleogene, reflecting the rising Rocky Mountains. Paleoenvir-

onments became marginal marine during deltaic deposition,

and fully marine offshore. Along the eastern shoreline, deltaic

deposition was restricted both spatially and temporally.

In essence, the presence of the Western Interior Seaway

brought an abundance of tropical moisture from the Gulf of

Mexico northward,and cool-temperate waters from ArcticCanada

and Alaska southward across a large portion of North America.

This conduit, 3,000 km long by 1,000 km wide, and 500 m to

1,000 m at its deepest during maximum transgression, provided

a region for mixing of atmospheric and oceanic air masses

and ocean currents, with potential for CO

exchange across the

atmosphere-oceanic boundary.

During the past several decades, the Cretaceous world was

viewed as one in which global atmospheric circulation was

sluggish due to a low thermal gradient from the equator to

the poles. In this scenario, meridional heat transport from the

low to high paleolatitudes was apportioned more to the oceans

than to the atmosphere (Barron et al., 1993). Newer interpreta-

tions from the oceanic realm, however, indicate that the tropics

superheated, probably several times, between the middle and

latest Cretaceous interval. In this revised scenario, a thermal

gradient strong enough to induce atmospheric circulation may

have been established between the superheated tropics and

the cool polar regions (Norris et al., 2002).

Oceans

Sea level

Progressively through the Lower Cretaceous, sea level rose and

advanced with each pulse, and the climate warmed across the

globe. In North America, the Upper Albian marked the first

north-south connection of the Western Interior Seaway. In South

America and Africa, vast warm seas were connected in the Late

Albian. In Europe, northern cold-temperate and subtropical seas

linked up in at least two broad areas, one in western Europe during

the Albian and later, and the other in western Russia, Poland,

Czechoslovakia, Hungary, and Bulgaria.

Sea level reached a peak near the Cenomanian-Turonian

boundary, fell in the Upper Turonian, and rose again at the

Turonian-Coniacian boundary. Sea level remained high through

the Coniacian and Santonian and until the Lower Campanian,

as expressed in the Niobrara Formation and its equivalents

where limestones and/or chalks, chalky shales, and calcareous

shales were deposited continuously through the interval. Sea

level fell in the Lower Campanian, rose moderately in the

Middle Campanian, fell in the Upper Campanian, transgressed

even more moderately in the Lower Maastrichtian, underwent a

mild regression in the Middle Maastrichtian, and after a relative

rise in the early Late Maastrichtian, fell at the end of the latest

Maastrichtian. The extent of sea level fall in the Maastrichtian

is much debated, although a small Antarctic ice sheet has been

proposed (Miller et al., 1999; see also Sea level change, last

250 million years).

In the absence of ice during the Cretaceous, variations in

sea level are attributed to variable rates of sea floor spreading

and volumetric changes in the ocean basins. The net effect of

high sea level on climate in epicontinental seaways is in the

production of maritime climatic conditions within and adjacent

to the seaways, especially in the direction of the prevailing

winds.

Pulses of elevated seas brought tropical surface waters from

the Gulf Coast through the Western Interior Seaway to Canada.

Most important in this migration of surface currents was the

northward transport of tropical larvae, humidity, and relatively

high temperatures. Likewise, in the western European region

of the Northern Hemisphere, conduits for surface currents

developed during high stands of the sea and carried heat north-

ward. These simplified scenarios of heat export from the tro-

pics by surface currents were complicated by regional and

local tectonic episodes.

Elevated sea levels that flooded continents and tectonic sub-

sidence that occurred in active basins combined to create silled

basins with restricted circulation (Hay, 1988). Large portions of

the temperate water columns became oxygen-restricted and had

limited faunas. The higher the anoxic conditions, the more

depauperate the fauna was in calcareous shales, limestones,

and chalks. Ultimately, during the most anoxic of conditions,

there was little or no macrofauna or burrowing, and organic

carbon was preserved in the rock record as oceanic anoxic

events.

The rise and fall of sea level affected climatic conditions in

the atmosphere, which in turn affected hydrologic conditions

on the land and deposition in the sea. The atmosphere con-

tained latent moisture; rainy and dry seasons alternated. During

the rainy season, the land was flooded and streams and rivers

rose, with the runoff resulting in deltaic construction, down-

cutting, and an increase in sediment load and the rate of sedi-

mentation for quite a distance offshore. Increased rainfall

resulted in a reduction in the salinity of the sea and lack of

clarity of the water. Marine chemistry, sediment types, and bio-

tic compositions were affected. Likewise, during dry seasons,

evaporative conditions prevailed and moisture returned to the

atmosphere. The strength of the hydrologic cycle during the

Cretaceous was reported as being more vigorous than that of

today (White et al., 2001).

Marine temperatures

Cautions exist regarding our ability to predict pre-Pleistocene

mean annual temperatures because of salinity uncertainties

and potential problems with alkalinity/diagenesis (Crowley

and Zachos, 2000). Yet, new fossil discoveries and excellent

preservation of fossils, precision in analytical equipment, and

further understanding of diagenetic effects allow for continued

refinement of interpretations regarding temperatures from the

marine realm. For example, based on oxygen isotopes derived

from well-preserved planktic foraminifers, estimated sea sur-

face temperatures (SSTs) for the Late Albian and Early Ceno-

manian were 30–31



C (Norris and Wilson, 1998). From

calculations, the authors determined that SSTs of 35–37



would be necessary to support hypersaline conditions. A few

years later, isotopic determinations from planktic foraminifers

yielded SST interpretations of 33 –34  2



C (Norris et al.,

2002), thus moving into the range of hypersaline conditions.

A SST proxy that is free from diagenetic effects and salinity

corrections has emerged recently from organic chemical analyses

(Schouten et al., 2003). TEX

(tetraether index of 86 carbon

atoms) is based on membrane lipids of marine crenarchaeota. Cre-

narchaeota comprise 20–30% of the picoplankton component

CRETACEOUS WARM CLIMATES 215

in today’s ocean; they adjust the composition of their tetraether

membranes lipids in response to SST. This response, quantified

as TEX

, shows a strong linear relationship to today’sSSTand

was used to evaluate middle Cretaceous sea surface temperatures.

Results yielded high SSTs (e.g., Early Aptian equatorial Pacific

27–32



C, Late Albian 30



NAtlantic35



C, Cenomanian/Turo-

nian equatorial Pacific 34–36



C, Cenomanian/Turonian 30



Atlantic 34–36



C), indicating extreme warmth from the low lati-

tudes for the middle Cretaceous and giving support to the concept

of a more vigorous atmospheric circulation than that envisioned

10 years ago.

Subsurface circulation from hypersaline waters in low lati-

tudes was a prediction of Chamberlin (1906), with support

from Brass et al. (1982). Newest model results from the Cam-

panian indicate that deep water formation likely occurred at

high southern latitudes (Deconto et al., 2000a), rather than in

saline pools at low latitudes, and the process of subsurface

circulation was more like today’s thermohaline circulatory pro-

cess during this stage.

The land

Investigations into the role of vegetation and its effect on the dis-

tribution of warmth across the Cretaceous globe are evolving

rapidly, especially with numerical models. Otto-Bliesner and

Upchurch (1997) examined the interrelationships between high-

and middle-latitude forests and surface temperature regulation.

Model results for the K/T indicated that forest vegetation warmed

the global climate by 2.2



C. Forests caused warming in the high-

latitude land areas, heat was transferred to the oceans, sea-ice

formation was delayed, and winter temperatures over coastal land

increased. DeConto et al. (2000a) used an OGCM that was forced

by the Campanian climate simulated by the climate-vegetation

model. The inclusion of plant community information into numer-

ical models thus appears to assist with determinations of meri-

dional thermal gradients and ocean heat-transport mechanisms,

long-standing endeavors in the Cretaceous research community.

Correlation of events from terrestrial to marine sequences

was difficult in the past due to the relatively poor temporal pre-

cision extracted from terrestrial rocks and fossils. Recently,

however, analyses of co-equal events occurring between the

two realms have been accomplished through correlation of

the isotopic excursions of terrestrial and marine organic carbon,

and marine carbonate carbon (Hasegawa, 2003). For example,

based on a comparison of these species of carbon, it was pro-

posed that the climatic optimum of warm and humid conditions

occurred during the latest Cenomanian in northeastern Asia.

This is significant in that data from the large continental land-

mass can now be integrated into the global database and used

not only for climatic interpretations, but also for purposes of

model-data comparisons.

Pulses of warming associated with Deccan volcanism and

release of CO

are among the most current avenues of

marine-terrestrial investigation, and these are achieved by ana-

lysis of the

187

Os/

188

Os record. The seawater osmium ratio was

declining during the Maastrichtian and prior to the K/T bound-

ary. The drop in this ratio was identified from several Deep Sea

Drilling Project (DSDP) sites and shows an abrupt decline

more than 100,000 years before the boundary. The decline is

coincident with Deccan volcanism from continental flood

basalts (leading to release of CO

), and the Late Maastrichtian

warming of 3–5



C that is separate from that of the K/T

boundary extinction (Ravizza and Peucker-Ehrenbrink, 2003).

In a similar manner, terrestrial evidence for a warming is

also derived from carbon and oxygen isotopes from paleosol

carbonates. Through this method, it was determined that

atmospheric warming events occurred between 70 and 69

Myr, and also 65.5 and 65 Myr ago (Nordt et al., 2003),

indicating that dramatic climatic fluctuations occurred several

million years before as well as during the K/T boundary

extinction interval. Although a wealth of empirical data has

been accumulated from the K/T boundary extinction interval,

definitive processes causal to the mass extinction continue to

be debated vigorously.

Summary and conclusions

Empirical, analytical, and experimental techniques have been

applied toward understanding the development and persistence

of processes that produced Cretaceous warm climates. Used in

combination, biotic distributions, isotopic and/or geochemical

analyses, and numerical climate models yield the most compre-

hensive results and the most robust interpretations on Cretac-

eous climates.

The role of the atmosphere, and of the oceans, in distribut-

ing the warmth across the globe is as yet unresolved. The pro-

cess of atmospheric circulation transporting heat across the

latitudes may have been the same for icehouse and greenhouse

intervals, but the actual temperatures for each steady state sys-

tem were quite distinct. Cretaceous oceans distributed warmth

across latitudes via surface currents, but subsurface circulation

commenced in high latitudes, at least during the Campanian.

Determination of sea surface temperatures and the interaction

of the sea surface with atmospheric CO

exchange remain

active research areas.

Thermal changes in the Cretaceous tropics dispute the theory

of time-stable biotic evolution within equable tropical environ-

ments, and suggest an important role for large-scale disturbance

in the evolution of tropical, and perhaps temperate, ecosystems

(Johnson et al., 1996).

Erle G. Kauffman and Claudia C. Johnson

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Cross-references

Animal Proxies, Invertebrates

Animal Proxies, Vertebrates

Cretaceous/ Tertiary (K-T) Boundary Impact, Climate Effects

Flood Basalts: Climatic Implications

Geochemical Proxies (Non-Isotopic)

Greenhouse (Warm) Climates

Heat Transport, Oceanic and Atmospheric

Mesozoic Climates

Ocean Anoxic Events

Ocean Paleotemperatures

Organic Geochemical Proxies

Paleoclimate Modeling, Pre-Quaternary

Sea Level Change, Last 250 Million Years

Stable Isotope Analysis

CRETACEOUS/TERTIARY (K-T) BOUNDARY

IMPACT, CLIMATE EFFECTS

Introduction

Collisions of asteroids and comets with the Earth’s surface are rare

events that punctuate the geologic record. The collision of a large

asteroid or comet with a planetary surface produces a tremendous

explosion, equivalent to thousands of times the explosive energy

of the world’s combined nuclear arsenal. The end result is a crater

tens to hundreds of kilometers in size. Although the existence of

large impact structures on Earth is undisputed, the possible cli-

matic effects of an impact were not seriously considered until

1980, when a team led by the famous physicist Luis Alvarez and

his son, the geologist Walter Alvarez, suggested that the profound

end-Cretaceous mass extinction might have been caused by the

impact of a 10-km diameter asteroid or comet (Alvarez et al.,

1980). Dated to 65 million years ago, this extinction is the last of

the large, known mass extinctions on Earth and defines a major

geologic boundary between the Cretaceous and Tertiary (or Paleo-

gene, as it is often referred to in the recent literature) periods, the

K/T boundary. The discovery by the Alvarez team of an anoma-

lous enrichment of extraterrestrial iridium at the K/T boundary

in the sections at Gubbio, Italy and Stevns Klint, Denmark opened

a whole new era of reanalysis of the geologic record in search of

more clues that could confirm the impact hypothesis. Since

1980, the list of “clues” has become long and overwhelming, cul-

minating with the discovery of the K/T boundary crater, the Chic-

xulub structure, in the Yucatán Peninsula, Mexico (Figure C77).

That region is a partially submerged continental platform consist-

ing of a thick sedimentary sequence of carbonates and evaporites

overlying continental crust. At the time of impact, a shallow sea

a few tens to several hundreds of meters deep covered the impact

region. The crater is roughly 180–200 km in diameter and cur-

rently buried under about 1 km of Tertiary sediments. Such a large

impact event would affect the environment and climate world-

wide. However, while several impact structures larger than 100

km in diameter have been identified on the Earth’s surface, to this

date Chicxulub is the only large, known crater that closely coin-

cides in time with a mass extinction event.

Several short-term and long-term environmental effects

result from a large impact event (e.g., see Toon et al., 1997).

Short-term effects extend up to few weeks after the impact

and are generally believed to have little influence on the

long-term evolution of the climate. They include the localized

direct effects of shock waves generated by the impact in the

atmosphere, like blast waves (high pressure pulses traveling

through the atmosphere at high velocity), and at the Earth’s

surface, such as earthquakes and tsunamis. Effects that are

more widespread include the production of toxic gases like

nitrogen oxides (NO, NO

) and nitric acid (HNO

) by shock

heating of the atmosphere from the entering projectile as well

as the atmospheric re-entry of material initially ejected well

beyond the stratosphere by the impact. These gases also caused

massive destruction of stratospheric ozone. The re-entering

ejecta also produced intense friction heating of the atmo-

sphere, leading to major wildfires that, in turn, filled the lower

atmosphere with smoke, dust and pyrotoxins in a scenario

reminiscent of a nuclear winter. Indication that much of the

end-Cretaceous land biomass was consumed by fire is pro-

vided by widespread evidence of soot at the K/T boundary,

CRETACEOUS/ TERTIARY (K-T) BOUNDARY IMPACT, CLIMATE EFFECTS 217

Figure C77 Artistic view of the Chicxulub impact event: (a) end-Cretaceous, 5 s before impact (Yucatan peninsula, Mexico); (b) end-Cretaceous,

1 s before impact (approaching asteroid or comet); (c) end-Cretaceous, the impact event; (d) beginning of Tertiary, about 1 month after

impact (dust and sulfate clouds blocking sunlight); (e) Early Tertiary, 1,000 years after impact (final multiring Chicxulub structure) (paintings by

William K. Hartmann).

218 CRETACEOUS/ TERTIARY (K-T) BOUNDARY IMPACT, CLIMATE EFFECTS

consisting of up to 2% of the carbon content of the end-

Cretaceous biosphere. Soot is a strong absorber of solar radia-

tion, thus preventing it from reaching the surface.

Long-term effects extend over months to decades after the

impact, and can have profound effects on the environment both

directly and indirectly by perturbing the overall climate. They

include the radiative effects due to the stratospheric loading

of small size dust (e.g., Alvarez et al., 1980; Covey et al.,

1994; Luder et al., 2003), and greenhouse gases such as carbon

dioxide (CO

) and water vapor (H

O). Unique to the Chicxulub

impact is the climatic effect of sulfur-bearing gases (Ivanov

et al., 1996; Pierazzo et al., 1998), whose importance has been

inferred from the climatic effects associated with major volca-

nic eruptions (see below).

The geologic record of the end-cretaceous climate and

its perturbation

The end-Cretaceous geologic record can provide some con-

straints on the climatic effects of the Chicxulub impact event.

The Cretaceous Period is characterized by one of the warmest

climates documented in Earth’s history (Barron et al., 1995).

No evidence of extensive glaciation has yet been found for

any time during this geologic period (although late-Cretaceous

eustatic sea level fluctuations have been interpreted as being

caused by the growth and shrinkage of minor ice sheets on

Antarctica (see Sea level change, last 250 million years, ed.).

The late-Cretaceous climate was slightly cooler than the mid-

Cretaceous (up to 8–10



C), but its estimated mean global sur-

face temperatures were still several degrees (



C) higher than

present. Oceanic and continental data suggest equatorial paleo-

temperatures similar to the present-day, but a much lower lati-

tudinal temperature gradient and thus warmer high-latitude

temperatures.

Moving across the K/T boundary, the plant records of the

Northern Hemisphere (mainly North America) indicate a major

ecological disruption and quick recovery, which has been attribu-

ted to a brief low temperature excursion and change in the hydro-

logic cycle. This excursion, however, may not have affected the

Southern Hemisphere as much, since only minor palynological

changes have been found at the few investigated Southern Hemi-

sphere locations. The marine isotopic record, on the other hand,

shows little convincing evidence for either warming or cooling

across the K/T boundary (see Bolide impacts and climate). This

suggests that if any climate change occurred as a result of the

impact event, it must have been of very short duration. The mean

early Paleocene climate (64 Ma) was very similar to that of

the latest Maastrichtian (67 Ma). It must be kept in mind, how-

ever, that the pre-Paleogene geologic record is characterized by

rather low resolution, at best of the order of a few thousands years;

time-averaging effects (like bioturbation) may further hide short-

lasting shifts in the signal. Consequently, if evidence from a

climate change spanning a few hundred years is averaged within

a 10,000-year record, it may be very difficult to resolve in the

geologic record.

Impact-related climatic effects

The most widely discussed climatic effect of a large impact is

caused by dust loading of the upper atmosphere. Alvarez et al.

(1980) suggested that dust raised by the impact would have

spread worldwide in clouds thick enough to cause an endless

night and block photosynthesis for periods longer than a year,

with catastrophic consequences for the terrestrial biota. Today,

the original Alvarez’s scenario has been scaled down to a much

less catastrophic view where only the stratospheric portion of

the fine (sub-micron) dust from the impact can affect the cli-

mate on a global scale (Toon et al., 1997). It is well known

(and well-illustrated by the aftermath of the impact of comet

Shoemaker-Levy 9 on Jupiter, in July 1994), both from obser-

vations and computer modeling, that large impacts can inject

hot, melted or vaporized debris from the projectile and target

well above the tropopause and beyond the stratosphere, and

ballistically distribute it globally over the Earth.

The “darkness-at-noon” scenario (which incidentally led to the

“nuclear winter” idea) was first explored using simple numeric

models of the atmosphere’s radiative balance. The results initially

indicated that even a rapidly coagulating dust layer caused sub-

freezing temperatures in continental interiors for at least 2 months

after the impact and a global loss of photosynthesis for about half

that time. Subsequently it was realized that most of the impact-

generated dust (large particles) would have a very short residence

time in the atmosphere. Furthermore, only the stratospheric

portion of the fine (sub-micron) dust, which corresponds to a very

small fraction of the impact-produced dust (Toon et al., 1997),

would affect climate globally and over a significant period of time.

Therefore, the initial “darkness-at-noon” hypothesis was scaled

down to a less catastrophic view. There is still much uncertainty

on the amount and size distribution of dust injected into the atmo-

sphere by the impact, thus causing equal uncertainty on the extent

of the climatic change. An early three-dimensional (3D) atmo-

spheric general circulationmodel was used to investigate the effect

of a uniform, thick stratospheric dust-layer that would decay in

about a year (Covey et al., 1994). The results indicate a strong

and “patchy” cooling on land, with temperature declining by up

to 12



C, and a mild cooling (few degrees) over the oceans,

accompanied by a collapse of the hydrologic cycle. Although the

climate model was extended to only one year after the impact, it

suggested that the climate would probably recover in about a dec-

ade. Recently, the role of the oceans in the long-term response of

the radiative perturbation following a Chicxulub-type dust loading

of the atmosphere was explored using a two-dimensional (2D),

zonally-averaged dynamic ocean circulation model (Luder et al.,

2003). The model results indicate that in the first year after

the impact the sea surface temperature dropped by several degrees,

with the strongest effects in the equatorial regions, but this change

remained confined to the upper 200 m of the oceans. Deep-sea

temperatures started to change only after 100 years, but never

exceeded few tenths of



C. Overall, the dust-related climatic

change did not affect the structure of the ocean circulation. These

studies confirm the deep ocean’s crucial role as moderator of

the Earth’s climate.

The presence of a thick sequence of carbonates and evaporites

at the impact location brings into the discussion the atmospheric

release of large amounts of carbon dioxide (CO

), sulfur-bearing

gases, and water vapor from the sedimentary layer (e.g., Ivanov

et al., 1996; Pierazzo et al., 1998). Large amounts of greenhouse

gases could be responsible for an abrupt and long lasting climate

shift. Various computer simulations have been carried out to assess

the influence of the sedimentary layer in the climatic effects of the

Chicxulub impact. Early models suggested the release of CO

amounts large enough to affect the global climate significantly.

However, more detailed modeling studies (Ivanov et al., 1996;

Pierazzo et al., 1998) strongly revised the amount of CO

released

in the impact, and consequently its climatic effects (even these

lower estimates may have to be further downsized in light of

new laboratory and field studies on the behavior of carbonates in

impacts). Considering that the estimated atmospheric inventory

CRETACEOUS/ TERTIARY (K-T) BOUNDARY IMPACT, CLIMATE EFFECTS 219

of CO

toward the end of the Cretaceous was about 2–10 times

pre-industrial values, it is doubtful that the CO

directly released

by the Chicxulub impact event would have increased the end-

Cretaceous atmospheric inventory of CO

by more than 50%.

The radiative forcing associated with the impact-released CO

is in the range of 1.2–3.4 W m

2

, comparable to the estimated

forcing from greenhouse gases due to industrialization (Pierazzo

et al., 2003). This forcing is not considered to be large enough

to cause a climate change capable of a major mass extinction.

A bigger contribution to the atmospheric CO

inventory may

come from the impact-related wildfires. Estimates from the identi-

fied soot layer at the K/T boundary suggest at most a doubling

of the preimpact CO

inventory (Toon et al., 1997).

The release of sulfur-bearing gases and water vapor in the

stratosphere results in the production of sulfate aerosols, as

documented by volcanic eruptions. Sulfate aerosols are strong

absorbers of long-wave (LW) radiation, while they mainly scat-

ter short-wave (SW) radiation, resulting in a net cooling of the

Earth’s surface. The effect of injecting sulfur-bearing gases and

water vapor into the stratosphere has been investigated with

simple atmospheric models. Using a general planetary radiative

transfer model, Pope et al. (1997) modeled the solar flux

through a sulfate aerosol cloud. Sulfate formation from over

2  10

g of sulfur dioxide (SO

) and of water vapor (average

K/T boundary impact scenario) was modeled with a simple

coagulation model. Their results indicate that solar transmis-

sion would be significantly reduced (below 50% of normal)

for about 8–13 years after the impact and cause a negative for-

cing (300 W m

2

) about two orders of magnitude larger than

the estimated forcing from CO

. The presence of dust and soot

grains could speed up the coagulation process, significantly

decreasing the overall duration of the effect, but increasing its

magnitude by further decreasing the solar transmission. Pope

et al. (1997) estimated that forcing from the sulfates would

cause the continental surface temperature to drop by several

degrees, approaching freezing for several years. Using a similar

approach but assuming that S would be released as a mixture of

and SO

Pierazzo et al. (2003) suggest a slightly shorter

duration of the sulfate effect, with a 50% reduction in solar

transmission for 4 –5 years after the impact, but a stronger over-

all forcing (over 500 W m

2

). They also found no significant

change in the maximum forcing for sulfur loadings above about

3  10

g of S (or 6  10

gofSO

; minimum K/T boundary

impact scenario), suggesting a saturation of the sulfur-related

forcing. In particular, in investigating the immediate (few days)

post-impact climate responses using a single column version of

an atmospheric general circulation model, Pierazzo and collea-

gues found strong temperature drops over the summer hemi-

sphere continents, with milder effects over winter hemisphere

continents and oceans. Adding impact-loads of CO

to the sul-

fate effects mitigates the overall cooling only slightly (0.5



Cat

most). More comprehensive and detailed 3D climate system

models of the impact-related climate change at the end-Cretac-

eous are still not available.

Study of the responses of terrestrial biota to impact-related cli-

matic effects is still in its infancy. Lomax et al. (2000, 2001)useda

dynamic global vegetation-biogeochemistry model to investigate

the potential global-scale response of terrestrial ecosystems to an

increase in atmospheric CO

, and to a combined initial cooling

plusgreenhouse warmingresulting from a Chicxulub-type impact.

They found that a four to tenfold increase of the atmospheric end-

Cretaceous CO

inventory produces spatially heterogeneous

increases in the distribution of canopy structure, net primary

productivity, and soil carbon concentrations, possibly causing a

biotic feedback mechanism that would ultimately help long-term

climate stabilization (Lomax et al., 2000). They then investigated

the overall combined effect of dust/sulfates and CO

loading on

the terrestrial ecosystem by artificially reducing the mean annual

temperature in the model by 6



C and the solar irradiance by

30% for 100 years, and simulating the effects of wildfires by burn-

ing 25% of the vegetation carbon and setting the CO

concentra-

tion to ten times pre-impact levels (Lomax et al., 2001). The

model results indicate a collapse of the terrestrial net primary pro-

ductivity that recovered to pre-impact values within a decade. In

particular, they found that the overall change in productivity

and vegetation biomass is stronger at low latitudes than at high

latitudes, and is consistent with ecological extinction gradients

derived from terrestrial paleobotanical data.

Conclusions

The Chicxulub structure, 65 million years in age, is the young-

est of the known terrestrial impact structures larger than

100 km in diameter. Structures of this size are associated with

impact events that are capable of perturbing the environment

and climate worldwide. Likely global scale effects include

depletion of stratospheric ozone, ignition of global wildfires,

release of climatically active gases, production of toxic oxides

and acid rain, and atmospheric loading of dust and smoke.

When looking at the geologic record, it appears that any cli-

matic change introduced by the Chicxulub impact event did

not destabilize the paleoclimate, which was able to recover to

the pre-impact state over a geologically short timescale

(<10,000 years). On the other hand, a sudden, large climate

change lasting a few hundred years would still have enormous

repercussions on the evolution of the biosphere.

Various investigations on the climatic effects associated with

the K/T boundary impact indicate that an abrupt cooling epi-

sode occurred immediately after the impact and could have

extended for a few years. The atmospheric loading of impact-

produced CO

could then have caused a mild warming that

may have helped climate recovery. The impact apparently did

not affect the deep ocean. The slow response of the deep ocean

and the quick reaction of the terrestrial ecosystem contribute to

a fast recovery of the climate system.

To this day, the magnitude of the end-Cretaceous climate

change is still an open question; answering it may shed new light

on our understanding of the end-Cretaceous mass extinction.

Elisabetta Pierazzo

Bibliography

Alvarez, L.W., Alvarez, W., Asaro, F., and Michel, H.V., 1980. Extraterres-

trial causefor the Cretaceous-Tertiary extinction. Science, 208,1085–1108.

Barron, E.J., Fawcett, P.J., Peterson, W.H., and Pollard, D., 1995. A “simu-

lation” of mid-Cretaceous climate. Paleoceanography, 10, 953–962.

Covey, C., Thompson, S.L., Weissman, P.R., and MacCracken, M.C., 1994.

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impact on Earth. Glob. Planet. Change, 9, 263–273.

Ivanov, B.A., Badjukov, D.D., Yakovlev, O.I., Gerasimov, M.V., Dikov, Y.P.,

Pope, K.O., and Ocampo, A.C., 1996. Degassing of sedimentary

rocks due to Chicxulub impact: Hydrocode and physical simulations. In

Ryder, G., Fastovsky, D., and Gartner, S. (eds.), The Cretaceous-Tertiary

Event and Other Catastrophes in Earth History. Boulder, CO: Geological

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2000. Terrestrial ecosystem responses to global environmental change

220 CRETACEOUS/ TERTIARY (K-T) BOUNDARY IMPACT, CLIMATE EFFECTS

across the Cretaceous-Tertiary boundary. Geophys. Res. Lett., 27,

2149–2152.

Lomax, B., Beerling, D., Upchurch, G. Jr., and Otto-Bliesner, B., 2001.

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of the terminal Cretaceous impact event. Earth Planet. Sci. Lett., 192,

137–144.

Luder, T., Benz, W., and Stocker, T.F., 2003. A model for long-term climatic

effects of impacts. J. Geophys. Res., 108, DOI: 10.1029/2002JE001894.

Pierazzo, E., Kring, D.A., and Melosh, H.J., 1998. Hydrocode simulations

of the Chicxulub impact event and the production of climatically active

gases. J. Geophys. Res., 103, 28, 607–28, 626.

Pierazzo, E., Hahmann, A.N., and Sloan, L.C., 2003. Chicxulub and

Climate: Effects of stratospheric injections of impact-produced

S-bearing gases. Astrobiology, 3,99–118.

Pope, K.O., Baines, K.H., Ocampo, A.C., and Ivanov, B.A., 1997. Energy,

volatile production, and climatic effects of the Chicxulub Cretaceous/

Tertiary impact. J. Geophys. Res., 102(E9), 21, 645–21, 664.

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Cross-references

Bolide Impacts and Climate

Cretaceous Warm Climates

Sea Level Change, Last 250 Million Years

CRYOSPHERE

Introduction

The cryosphere (derived from the Greek kryos for ‘‘cold’’)

comprises all of the frozen water and soil on the surface of

the Earth. While this term is often used by climatologists, a

more accurate name for the study of all aspects of ice and snow

is glaciology. Ice in clouds is also an important element of cli-

mate but is not considered part of the cryosphere. Because the

freezing point of water lies in the center of the range of terrest-

rially achieved temperatures, the cryosphere is a very change-

able feature of the Earth’s surface. Its six main elements, in

diminishing order of area for today’s climate, are: (a) seasonal

snow cover, (b) sea ice, (c) permafrost, (d) ice sheets, (e) river

and lake ice, and (f) mountain glaciers and small ice caps.

Seasonal snow cover is the most dynamic element of the

cryosphere. It responds rapidly to atmospheric conditions on

timescales of days to weeks and, in fact, exhibits the greatest

seasonal variation of any geophysical element on the Earth’s

surface. Although it has a small seasonal heat storage capacity,

its main effect on climate comes from its high albedo (fraction

of incident solar radiation reflected back to space). Fresh snow

has an albedo of 0.9. Sea ice forms when sea water is cooled

below its freezing point. This freezing is a complex process

because the freezing point of water depends on the salt concen-

tration of the ocean. A sea ice field usually consists of numer-

ous individual floes, typically 0.5 –4 m thick, with diameters

between 1 meter and many kilometers. Sea ice has a high

albedo and therefore the effect its seasonal cycle exerts on the

surface heat balance of the Earth is similar to, though smaller

than, that exerted by seasonal snow cover. Sea ice also acts

as a barrier to the exchange of heat, moisture, and momentum

between the atmosphere and the ocean, and furthermore plays a

role in the formation of deep water masses in some oceanic areas.

Sea ice is covered by snow for much of its lifetime. River and

lake ice form in a similar manner to sea ice but without

the complications of high salt concentrations. Their climatic

behavior resembles that of sea ice. Ice sheets are quasi-permanent

features, responding to environmental changes on time scales

of millennia. T oday, only the two polar ice sheets on Antarctica

and Greenland remain. Together , they contain by far the greatest

proportion of mass of the cryosphere but are least sensitive to cli-

mate change on the shorter timescales. Ice sheets have a high ele-

vation and high albedo, and so act as elevated cooling surfaces

for atmospheric heat balance. They also play an important role

in the freshwater balance of the oceans. Changes in their volume,

though slow, are the main cause for sea-level changes on geolo-

gical timescales. Mountain glaciers and small ice caps are a small

component of the cryosphere. They are distinguished from the

large ice sheets by their size and by the fact that they are topogra-

phically controlled. They are too small to have a direct influence

on climate; however, their high rates of mass turnover allow rapid

changes in response to climatic change (timescales of the order of

decades to centuries). Permafrost is ground material that remains

below freezing for at least two consecutive years, whether or not

it contains ice. It may be continuous, underlying the whole of a

region, or discontinuous, occurring in patches. In many regions

the uppermost, so-called active layer, thaws annually to a depth

of 1 to a few meters. Permafrost is a product of heat exchange

between the land surface and the atmosphere, changing on time-

scales of centuries. It affects surface ecosystems and river dis-

charge, and changes in its extent can control rates of trace gas

emissions, especially of methane. Permafrost records temperature

changes, hence its role as geo-indicator for monitoring and asses-

sing environmental change.

Current distribution of ice in the cryosphere

Today, nearly 80% of the Earth’s freshwater is stored on the con-

tinents in the form of ice, almost all of it in the polar regions and

in high mountain areas, cf. Table C9 and Figure C78.Perennial

ice covers 10% of the Earth’s land surface and 4% of the oceans.

The seasonal distribution of snow, lake and riverice, and to a lesser

extent, sea ice, is closely linked with atmospheric conditions, in

particular with temperature.

Seasonal snow

Seasonal snow reaches its maximum extent in late winter,

when it covers almost 50% of the land surface of the Northern

Hemisphere. It is almost absent in the Southern Hemisphere,

where it is limited to the mountainous areas of New Zealand,

southeastern Australia, and South Africa; it is more widespread

only in Patagonia and ice-free regions of Antarctica and on

sub-Antarctic islands.

Sea ice

The seasonal variation of sea ice is much smaller in the Arctic

than in the Antarctic owing to the different geography of the

two polar regions. The Arctic Ocean is a truly polar ocean sur-

rounded by land, with only a limited range of longitudes within

which sea ice can expand seasonally into lower latitudes. By

contrast, in the Southern Hemisphere the central polar region

is covered by the continent of Antarctica, preventing sea ice

from extending to the highest latitudes in summer but at the

same time allowing winter sea ice to extend in a roughly con-

centric zone around Antarctica at almost all longitudes. A sec-

ond contrast between the two hemispheres is the larger upward

oceanic heat flux in the Antarctic, as less fresh water enters

the ocean to cause a low-salinity layer of polar surface water

CRYOSPHERE 221

hampering vertical heat exchange from below. For these rea-

sons, the Arctic features a thicker permanent sea ice cover of

multiyear ice, augmented in winter by a seasonal ice cover

of first-year ice extending into the surrounding subpolar seas.

Around Antarctica, however, almost all of the winter ice melts

in summer, leaving multiyear ice confined only to the western

Weddell Sea, the Bellingshausen and Amundsen Seas, and

coastal embayments. Arctic sea ice generally reaches its maxi-

mum extent in March and minimum extent in September. By

contrast, Antarctic autumnal ice growth is slower and spring

decay about one month faster.

Polar ice sheets

Present-day continental ice cover is entirely dominated by the

Antarctic Ice Sheet, which is responsible for 90% of the

volume and 85% of the area of all land ice. Most of this ice

is in the East Antarctic Ice Sheet, which constitutes a vast, rela-

tively flat dome with a maximum elevation of 4,030 m and

a maximum known thickness of 4,776 m. The Transantarctic

Mountains divide the East Antarctic Ice Sheet from the

West Antarctic Ice Sheet and the Antarctic Peninsula. The West

Antarctic Ice Sheet rests on bedrock below sea level and lar-

gely drains into two big ice shelves in the Ross and Weddell

Seas. These ice shelves are up to 1,000 m thick and float in

equilibrium with the ocean water. Smaller ice shelves fringe

the Antarctic coastline elsewhere. Ice flow towards the margin

results mainly from internal ice deformation under the action of

gravity. At the margin, the flow is often channeled in outlet gla-

ciers and in ice streams, in which the dominant flow mechan-

ism is basal sliding, and that are responsible for the bulk of

the ice discharge across the grounding line into the ice shelves.

Mass is lost from the ice shelves by bottom melting and by cal-

ving of icebergs from their ice fronts. Current Antarctic surface

temperatures are so low that there is virtually no surface melt-

ing. These low temperatures also limit the amount of water

vapor that can be advected inland. Accumulation rates over

the vast interior of the continent are only a few cm per year

and this makes the central Antarctic Plateau one of the driest

places on Earth.

The only Arctic ice mass of significance is the Greenland

Ice Sheet. It is a relict from the ice ages that overlies a bowl-

shaped continent and has a maximum surface elevation of

3,230 m. In contrast to Antarctica, summer temperatures on

Greenland are high enough to cause widespread summer melt-

ing, resulting in a negative mass balance below elevations of

about 500 m in the north and 1,800 m in the south. Greenland

has no major ice shelves apart from a few small ones along the

north and northeast coast. Ice not lost by ablation is discharged

into the ocean by iceberg calving from glaciers, in roughly the

same amount as runoff. Complete melting of both polar ice

sheets would raise global sea level by about 70 m.

Glaciers and ice caps

The water currently contained in glaciers and ice caps (exclud-

ing Antarctica and Greenland) is equivalent to about 0.5 m of

global sea level. Half of it is in small ice caps, mostly in the

Arctic, and the other half is stored in mountain glaciers else-

where. Most mountain glacier ice is in Alaska, followed by

Central Asia and Patagonia. Ice volume contained in the Alps

is less than 1% of total glacier volume, but its glaciers have

been studied the most. There are more than 160,000 glaciers

worldwide. Glaciers enlarge when the accumulation of snow

and ice exceed the loss by melting, and sometimes calving.

Net accumulation occurs at higher altitude, net ablation at

lower altitude. To compensate for net accumulation and abla-

tion, the ice flows downhill by internal deformation and basal

sliding. The primary controls on glacier mass balance are tem-

perature, especially in summer, and accumulation, especially in

winter. The nature of the response varies from glacier to glacier

depending on mass-balance gradient, hypsometry, and glacier

length.

Permafrost

About 24% of the exposed land area of the Northern

Hemisphere is currently underlain by permafrost, most of it

in the tundra of Canada, Alaska, and Russia (Zhang et al.,

1999). It is also present under shallow polar sea beds, in ice-

free areas in Antarctica, on some sub-Antarctic islands, and in

many mountain areas and high plateaus of the world. Its distri-

bution roughly follows the pattern of air temperature, though

some of it is relict, having formed during colder glacial periods

and having melted only very slowly since then. Sizeable areas

below the polar ice sheets and small ice caps are also below

freezing. This thermal condition of the ground extends to max-

imum depths of about 1,500 m in Siberia, 1,000 m in the Cana-

dian Arctic, and 680 m in Alaska, but often it is also less than

20 m thick. Generally, permafrost with high ice content (>20%

by volume) is found at high latitudes, and permafrost with

low ice content (<10% by volume) mainly in mountainous

regions and high plateaus. Almost half of the total permafrost

area is continuous permafrost. The remaining part is made

up of discontinuous and sporadic permafrost or occurs in

isolated patches.

Table C9 Size of the present-day cryosphere and average mass

turnover times of its main components as given by the ice flux and

average residence time (adapted from Fitzharris et al., 1996; Vaughan

et al., 1999; Zhang et al., 1999; Huybrechts et al., 2000; Church et al.,

2001; Lythe et al., 2001)

Cryospheric element Area

(10

)

Ice volume

(10

)

Ice flux

(km

1

)

Residence

time

(yr)

Seasonal snow cover 47.2 0.01 25,000 0.4

Northern Hemisphere

(max)

46.3

Southern Hemisphere

(max)

0.9

Sea ice

Northern Hemisphere

(max/ min)

16.0/ 9.0 0.05/ 0.03 40,000 1

Southern Hemisphere

(max/ min)

22.0/ 4.0 0.03/

<0.01

20,000 1

Permafrost 25.4 0.02 40 500

Ice sheets

Antarctic Ice Sheet

(grounded ice)

12.37 25.71 2,100 12,500

Antarctic ice shelves 1.49 0.66 540 1,200

East Antarctic Ice

Sheet

10.09 22.59 1,350 17,000

West Antarctic Ice

Sheet

2.28 3.12 750 4,000

Greenland Ice Sheet 1.71 2.85 570 5,000

River and lake ice <1.0 0.3

Mountain glaciers and

small ice caps

0.68 0.18 750 250

The residence time is defined as the ice volume divided by the ice flux.

222 CRYOSPHERE

Current evolution of the cryosphere

The current evolution of many elements of the cryosphere,

especially of those with the shorter response times, appears

well correlated with the recent global warming trend of about

0.6



C since the late nineteenth century.

Snow covers in temperate regions are usually thin and often

close to their melting point, hence they are sensitive indicators

of climate change. Monitoring of seasonal snow is practical

only with satellites, so there are no reliable records prior to

1971. Visible-band and passive microwave data since then

show considerable variability from year to year, but also

reveal a decline in the extent of snow cover in the Northern

Hemisphere by about 10% since the mid-1980s, largely due

to a decrease in spring extent (Robinson et al., 1993). This

reduction in snow cover is highly correlated with increasing air

temperature. Snow-cover observations at ground stations have

been recorded for many decades and suggest that Northern

Hemisphere spring extents are currently at their lowest values in

the past 100 years. Numerous small perennial snow patches have

disappeared over the twentieth century, commensurate with the

general warming trend.

Like seasonal snow, sea ice is also thought to be sensitive to cli-

mate change but the mechanisms are less well understood as addi-

tional processes such as ice transport, ocean circulation, and

convection depths come into play. Warming is nevertheless

thought to cause a reduction in both area and thickness of sea

ice, which sets in motion a positive feedback, as more solar radia-

tion is absorbed and more oceanic heat is transferred to the atmo-

sphere. Satellite passive microwave data show a marked decrease

of perennial Arctic sea ice extent for the period 1978–2000 of 9%

per decade consistent with a summer warming of 1.2



Cperdec-

ade over the same period (Comiso, 2002). Antarctic records of

sea-ice extent, on the other hand, show large variability but a

negligible, albeit slightly positive, overall trend. Knowledge of

sea-ice thickness largely comes from upward sonar profiling from

submarines in the Arctic and more recently, from buoys moored

to the sea floor. Observations of sea ice thickness have also

been derived from satellite altimetry of the freeboard height, but

Figure C78 Geographic extent of main components of the cryosphere in both hemispheres. The left panels are for the present distribution,

the right panels depict the situation during the Last Glacial Maximum between 23,000 and 19,000 years ago (adapted from maps

presented in Cooke and Hays, 1982; Untersteiner, 1984; Frenzel et al., 1992; Fitzharris et al., 1996; Van Vliet-Lanoe

and Lisitsyna, 2001, and the

author’s own ice sheet model results).

CRYOSPHERE 223

geographical coverage and/or record length do not yet allow for

strong conclusions. Summer Arctic sea ice cover and thickness

have decreased in recent years (e.g., Lindsay and Zhang, 2005;

Serreze et al., 2007).

Observations of lake and river ice exist primarily in the

form of freeze-up and break-up dates. Records for the Northern

Hemisphere since 1850 show significant evidence of later

autumn freeze-up and earlier spring break-up of on average

5.8 days and 6.5 days per 100 years, respectively, consistent

with the warming trend from the beginning of the twentieth

century (Magnuson et al., 2000).

The present evolution of the polar ice sheets is still not

known with confidence but a picture is starting to emerge from

a combination of mass budget studies, satellite altimetry, and

modeling (Rignot and Thomas, 2002). Current evidence indi-

cates that the Antarctic Ice Sheet has probably lagged behind

in the last glacial cycle of retreat. Most of the ongoing shrink-

ing probably occurs in West Antarctica as a delayed response to

post-glacial sea-level rise, whereas the East Antarctic Ice Sheet

may be close to a stationary state, or growing slightly in res-

ponse to increased accumulation rates. Superimposed on this

long-term trend is a likely thickening effect of modern accumu-

lation increases. Larger imbalances have recently been detected

in West Antarctica, positive for the Siple Coast ice stream area

and negative for the Pine Island/Thwaites catchment area.

These probably represent local ice-dynamic effects that are

unrelated to climate. The Greenland Ice Sheet is believed to

be close to an overall neutral equilibrium. Current data show

a consistent picture of a small thickening of the accumulation

zone offset by larger thinning rates at lower altitudes. Accord-

ing to the Intergovernmental Panel on Climate Change (IPCC),

both ice sheets together may have contributed between 2cm

and +6 cm to the observed twentieth century sea-level rise

(Church et al., 2001).

A wealth of data exists on the geometry of valley glaciers.

A large amount of information is available from sketches, etch-

ings, paintings, and old photographs. In many cases, informa-

tion from terminal moraines and trimlines is available to

construct the history of a glacier over the last few centuries.

An overwhelming characteristic of global glacier length records

is the almost uniform recession of glaciers, which started about

1850 at the end of the Little Ice Age and is still continuing

today. Rising global temperature is the most likely explanation

and so the worldwide glacier retreat is probably one of the

strongest measures of global warming. The only exceptions

today are Norwegian and New Zealand glaciers, which are

currently advancing, due to local increases in precipitation.

However, this is an exceptional state that can probably not be

sustained in the long term. Global glacier melt assessments

indicate that since the late nineteenth century, glaciers and ice

caps have lost between 5 and 10% of their volume and a little

less than that in area. The IPCC reports a glacier contribution

to global sea-level rise of between 2 and 4 cm over the last

100 years, less than the contribution from thermal expansion

but similar to the combined contribution from the Antarctic

and Greenland Ice Sheets. Findings of ancient remains in the

European Alps (e.g., the 5,000-year old Oetztal “ice man”)

indicate that current glacier recession is reaching levels not

seen for probably several millennia. Recent studies suggest that

the twentieth century rate of glacier wastage in Alaska and

the Patagonian Andes doubled after the mid-1990s (Arendt

et al., 2002; Rignot et al., 2003).

Deep permafrost temperatures change by conduction and

therefore are a manifestation of past climatic conditions over

long spans of time. Permafrost is also very sensitive to changes

in its surface energy balance. Very small changes in surface cli-

mate can produce robust changes in permafrost temperatures.

Long-term measurements in deep boreholes in Alaska, Canada,

and elsewhere demonstrate a distinct but spatially heteroge-

neous warming trend in low-land permafrost (Romanovsky

and Osterkamp, 2001). Temperature measurements in northern

Alaska have demonstrated a warming of the permafrost of

2–4



C over the last century. In North America, the southern

extent of the discontinuous permafrost zone has migrated north-

ward in response to warming after the Little Ice Age, and con-

tinues to do so today. In China, both an increase in the lower

altitudinal limit of mountain permafrost and a decrease in areal

extent have been observed. Increases of the thickness of the

active layer in response to climate warming have also been

reported.

The cryosphere during the last glacial cycle

During the last ice age, which was most severe from about

23,000 to 19,000 years ago, an extensive ice complex stretched

across North America, Greenland, the polar seas, and parts

of northern Eurasia. It consisted of huge land-based glaciers,

marine-based ice sheets, and either permanent pack-ice or

shelf ice. Ice increased synchronously over Antarctica and the

Southern Ocean. Sea levels were 125 m lower than today

because water was bound up in continental ice sheets that

increased their total mass by about a factor of three (Clark

and Mix, 2002; Figure C78 and Table C10).

The surface of the Earth during the Last Glacial Maximum

(LGM) was comprehensively mapped by the CLIMAP (Climate:

Long-range Investigation, Mapping, and Prediction) project in

the 1970s from a combination of biological and lithological

evidence from deep-sea cores and from the land (CLIMAP,

1976). Best documented is the chronology of Northern

Hemisphere ice-sheet extent. The largestglacier complex was situ-

ated over North America. It is inferred to have consisted of several

domes centered over the high plateaus of eastern Canada and the

Rocky Mountains. An ice-free corridor separated the Cordilleran

Ice Sheet from the Laurentide Ice Sheet until late in the glacial

Table C10 Size of main cryospheric elements during the Last Glacial

Maximum (adapted from Denton and Hughes, 1981; Cooke and Hays,

1982; Van Vliet-Lanoe

and Lisitsyna, 2001; Huybrechts, 2002; Clark and

Mix, 2002)

Cryospheric element

Area

(10

)

Ice volume

(10

)

Excess ice-

equivalent

sea level (m)

All land ice 41.8 75–88 110–135

Laurentide Ice Sheet 14.6 31–38 75–88

Fennoscandian Ice Sheet 6.2 5.8–7.5 14–18

Greenland Ice Sheet 2.6 4.0–4.5 2–3

Antarctic Ice Sheet (grounded) 15.1 33–36 15–20

Other ice caps and glaciers 3.3 1.4–2.2 4–6

Northern Hemisphere sea ice

(max/ min)

18/ 11 ––

Southern Hemisphere sea ice

(max/ min)

35/ 25 ––

Exposed permafrost (continuous

and discontinuous)

30 ––

224 CRYOSPHERE