this d
18 O shift. Following the boundary event,
planktonic foraminifera d
18 O record during the
Oligocene and early Miocene mirrored the benthic
record in many respects. For much of the Oligocene
and early Miocene, the absolute values are close to
À 0.5%, which translates into a temperature estimate of 211C (Figure 7B). By 15 Ma, tropical surface waters had warmed to 261C.
The middle Miocene d
18 O shift represents an increase of 1.5% in the benthic record between 15 and
13 Ma. This transition is composed of two sharp
increases around 14 and 13 Ma (Figure 8B). These
d
18 O steps occurred in less than 200 000 years with
each recording an increase of B1% followed by a
small decrease. During these two shifts, deep waters
cooled from 9 to 51C. The planktonic foraminiferal
d
18 O record from 15 to 13 Ma shows two increases
as recorded in the benthic foraminiferal record
(Figure 7). However, it does not show the large permanent shift recorded by the benthic foraminifera,
indicating a small cooling from 26 to 241C. From 13
to 3 Ma, the deep ocean cooled slightly from 5 to
31C while the surface waters warmed from 24 to
261C (Figure 7B).
The last of the large d
18
O steps in the Cenozoic
was recorded during the late Pliocene from 3.2 to
2.6 Ma. This ‘step’ is better characterized as a series
of d
18 O cycles with increasing amplitudes and values
over this interval (Figures 6 and 8C). The cycles
have been subsequently determined to be 40 000
year cycles related to variations in the solar radiation received in the high latitudes. This
interval ushered in the large-scale Northern Hemisphere ice sheets that have since dominated Earth’s
climate. At 2.6 Ma, the first IRD was deposited in
the open North Atlantic and was coeval with the
d
18 O maximum. Prior to 2.6 Ma, IRD was confined
to the marginal basins to the north, Greenland’s
and Iceland’s continental margins. Subsequent
d
18 O maxima were associated with IRD. Between
2.6 and 1 Ma, large Northern Hemisphere ice sheets
waxed and waned on the 40 000 year beat. Beginning around 1 Ma, the ice sheets increased in size
and switched to a 100 000 year beat (Figure 6).
During this interval, deep-water temperatures remained similar to those in the modern ocean (0 to
31C).
The planktonic foraminiferal d
18
O response during the late Pliocene event shows the cyclic behavior,
but not the overall increase recorded by the benthic
foraminifera. As with the middle Miocene d
18 O shift,
the late Pliocene increase represents the cyclic buildup of ice sheets accompanied by deep water cooling.
The tropical surface water temperatures, however,
varied between 26 and 281C.
Mechanisms for Climate Change
Most climate change hypotheses for the Cenozoic
focus on either oceanic heat transport and/or greenhouse gas concentrations. Each mechanism produces
different responses in the equatorial-to-pole and
surface-to-deep temperature gradients. An increase
in the meridional heat transport generally cools the
tropics and warms the poles. If poleward transport of
heat decreases, then the tropics will warm and the
poles will cool. Variations in greenhouse gas concentrations should produce similar changes in both
the tropical and polar regions.
Tropical surface water and deep-ocean records covaried for the first part of the Cenozoic. The warming and subsequent cooling between 65 and 34 Ma
are most often ascribed to changing greenhouse gas
concentrations. The interval of warming that began
around 58 Ma and peaked at 52 Ma coincided with
the release of large amounts of CO 2 into the atmosphere as a consequence of tectonic processes.
The eruption of the Thulean basalts in the northeastern Atlantic Ocean began during the Paleocene
and peaked around 54 Ma. It is also recognized that
there was a large-scale reorganization of the midocean ridge hydrothermal system which began during the late Paleocene and extended into the Eocene.
Both tectonic processes accelerate mantle degassing
which raises atmospheric levels of CO 2 . Recently,
evidence for another potentially large CO 2 reservoir
was found along the eastern continental margin of
North America. Methane hydrates frozen within the
sediments appear to have released catastrophically at
least once and possibly multiple times during the
latest Paleocene and early Eocene (B58–52 Ma).
One or all of these sources could have contributed to
the build-up of greenhouse gases in the atmosphere
between 58 and 52 Ma. Following the thermal
maximum, the long-term cooling in both the surface
and deep waters implies that greenhouse gas concentrations slowly decreased. Proxies for estimating
pCO 2 concentrations (d
13
C fractionation within organic carbon and boron isotopes) are still being developed and refined. However, preliminary
indications are that atmospheric pCO 2 levels were
high (>1000 p.p.m.) during the early Eocene, dropped to B400–500 p.p.m. during the middle to late
Eocene, and reached late Pleistocene concentrations
(200–300 p.p.m.) by the early Oligocene.
The deep-water temperature cooling across the
Eocene/Oligocene boundary (Figure 7B) was not
accompanied by tropical cooling, and resulted from
the first step in the thermal isolation of Antarctica. In
modern ocean, the Antarctic Circumpolar Current is
a vigorous surface-to-bottom current that provides
420 CENOZOIC CLIMATE – OXYGEN ISOTOPE EVIDENCE
18 O shift. Following the boundary event,
planktonic foraminifera d
18 O record during the
Oligocene and early Miocene mirrored the benthic
record in many respects. For much of the Oligocene
and early Miocene, the absolute values are close to
À 0.5%, which translates into a temperature estimate of 211C (Figure 7B). By 15 Ma, tropical surface waters had warmed to 261C.
The middle Miocene d
18 O shift represents an increase of 1.5% in the benthic record between 15 and
13 Ma. This transition is composed of two sharp
increases around 14 and 13 Ma (Figure 8B). These
d
18 O steps occurred in less than 200 000 years with
each recording an increase of B1% followed by a
small decrease. During these two shifts, deep waters
cooled from 9 to 51C. The planktonic foraminiferal
d
18 O record from 15 to 13 Ma shows two increases
as recorded in the benthic foraminiferal record
(Figure 7). However, it does not show the large permanent shift recorded by the benthic foraminifera,
indicating a small cooling from 26 to 241C. From 13
to 3 Ma, the deep ocean cooled slightly from 5 to
31C while the surface waters warmed from 24 to
261C (Figure 7B).
The last of the large d
18
O steps in the Cenozoic
was recorded during the late Pliocene from 3.2 to
2.6 Ma. This ‘step’ is better characterized as a series
of d
18 O cycles with increasing amplitudes and values
over this interval (Figures 6 and 8C). The cycles
have been subsequently determined to be 40 000
year cycles related to variations in the solar radiation received in the high latitudes. This
interval ushered in the large-scale Northern Hemisphere ice sheets that have since dominated Earth’s
climate. At 2.6 Ma, the first IRD was deposited in
the open North Atlantic and was coeval with the
d
18 O maximum. Prior to 2.6 Ma, IRD was confined
to the marginal basins to the north, Greenland’s
and Iceland’s continental margins. Subsequent
d
18 O maxima were associated with IRD. Between
2.6 and 1 Ma, large Northern Hemisphere ice sheets
waxed and waned on the 40 000 year beat. Beginning around 1 Ma, the ice sheets increased in size
and switched to a 100 000 year beat (Figure 6).
During this interval, deep-water temperatures remained similar to those in the modern ocean (0 to
31C).
The planktonic foraminiferal d
18
O response during the late Pliocene event shows the cyclic behavior,
but not the overall increase recorded by the benthic
foraminifera. As with the middle Miocene d
18 O shift,
the late Pliocene increase represents the cyclic buildup of ice sheets accompanied by deep water cooling.
The tropical surface water temperatures, however,
varied between 26 and 281C.
Mechanisms for Climate Change
Most climate change hypotheses for the Cenozoic
focus on either oceanic heat transport and/or greenhouse gas concentrations. Each mechanism produces
different responses in the equatorial-to-pole and
surface-to-deep temperature gradients. An increase
in the meridional heat transport generally cools the
tropics and warms the poles. If poleward transport of
heat decreases, then the tropics will warm and the
poles will cool. Variations in greenhouse gas concentrations should produce similar changes in both
the tropical and polar regions.
Tropical surface water and deep-ocean records covaried for the first part of the Cenozoic. The warming and subsequent cooling between 65 and 34 Ma
are most often ascribed to changing greenhouse gas
concentrations. The interval of warming that began
around 58 Ma and peaked at 52 Ma coincided with
the release of large amounts of CO 2 into the atmosphere as a consequence of tectonic processes.
The eruption of the Thulean basalts in the northeastern Atlantic Ocean began during the Paleocene
and peaked around 54 Ma. It is also recognized that
there was a large-scale reorganization of the midocean ridge hydrothermal system which began during the late Paleocene and extended into the Eocene.
Both tectonic processes accelerate mantle degassing
which raises atmospheric levels of CO 2 . Recently,
evidence for another potentially large CO 2 reservoir
was found along the eastern continental margin of
North America. Methane hydrates frozen within the
sediments appear to have released catastrophically at
least once and possibly multiple times during the
latest Paleocene and early Eocene (B58–52 Ma).
One or all of these sources could have contributed to
the build-up of greenhouse gases in the atmosphere
between 58 and 52 Ma. Following the thermal
maximum, the long-term cooling in both the surface
and deep waters implies that greenhouse gas concentrations slowly decreased. Proxies for estimating
pCO 2 concentrations (d
13
C fractionation within organic carbon and boron isotopes) are still being developed and refined. However, preliminary
indications are that atmospheric pCO 2 levels were
high (>1000 p.p.m.) during the early Eocene, dropped to B400–500 p.p.m. during the middle to late
Eocene, and reached late Pleistocene concentrations
(200–300 p.p.m.) by the early Oligocene.
The deep-water temperature cooling across the
Eocene/Oligocene boundary (Figure 7B) was not
accompanied by tropical cooling, and resulted from
the first step in the thermal isolation of Antarctica. In
modern ocean, the Antarctic Circumpolar Current is
a vigorous surface-to-bottom current that provides
420 CENOZOIC CLIMATE – OXYGEN ISOTOPE EVIDENCE