376
DYNAMICAL OCEANOGRAPHY
In this chapter the focus will be on the different roles of the transport of
heat and salt in the component of the ocean circulation associated with
density differences, the thermohaline circulation. It will appear that the
stability properties of the large-scale ocean circulation are highly sensitive to the representation of these transports. Note that we have totally
neglected these aspects in chapter 13, where only the density field was
considered and not the separate effects of heat and salt. The motivation
to study the thermohaline circulation in isolation is that variability in this
flow may have been responsible for rapid climate change in the past (section 16.1). In section 16.2 we present an idealized model to investigate
the stability of the thermohaline circulation, the Stommel two-box model.
In the following sections (16.3 and 16.4) we analyze the equilibrium solutions of this model and the feedbacks affecting the stability of the thermohaline circulation (section 16.5).
16.1. Past climate variability
Much information on past climates has been obtained through measurement
of isotope content, such as oxygen and carbon isotopes in material derived from
ocean sediments and from ice cores, combined with accurate dating techniques.
For example, the carbonate in shells of marine organisms (e.g. foraminifera) and
water in ice caps contain two isotopes of oxygen, 18 O and 16 O. The normalized
isotope ratio δ 18 O is calculated as a deviation from a reference sample as
δ
18 O =
(
18 O
16 O
) sample − (
18 O
16 O
) reference
(
18 O
16 O
) reference
,
(16.1)
where the reference sample is different for ice cores (i.e. standard mean ocean
water) than for carbonate shells (i.e. a specific fossil Cretaceous species). The
isotope 16 O is lighter than 18 O so that water containing 16 O is preferentially evaporated and a temperature-dependent fractionation occurs. Under cold conditions,
less water containing 18 O is able to evaporate into the atmosphere.
Changes in δ 18 O reflect the combined effect of changes in global ice volume
and temperature at the time of deposition of the sampled material. During very
cold conditions, global ice volume is relatively large and hence sea level is low,
which enriches water in the ocean with 18 O. Also because of the colder temperatures, more 18 O remains in the ocean and less 18 O becomes locked in the ice.
Hence, in ocean sediments the ratio δ 18 O will increase under cold conditions,
whereas in ice cores it will decrease.
When corrections for global ice volume (with respect to the reference sample)
are made, δ 18 O can be used as an indicator for the temperature at the time of
DYNAMICAL OCEANOGRAPHY
In this chapter the focus will be on the different roles of the transport of
heat and salt in the component of the ocean circulation associated with
density differences, the thermohaline circulation. It will appear that the
stability properties of the large-scale ocean circulation are highly sensitive to the representation of these transports. Note that we have totally
neglected these aspects in chapter 13, where only the density field was
considered and not the separate effects of heat and salt. The motivation
to study the thermohaline circulation in isolation is that variability in this
flow may have been responsible for rapid climate change in the past (section 16.1). In section 16.2 we present an idealized model to investigate
the stability of the thermohaline circulation, the Stommel two-box model.
In the following sections (16.3 and 16.4) we analyze the equilibrium solutions of this model and the feedbacks affecting the stability of the thermohaline circulation (section 16.5).
16.1. Past climate variability
Much information on past climates has been obtained through measurement
of isotope content, such as oxygen and carbon isotopes in material derived from
ocean sediments and from ice cores, combined with accurate dating techniques.
For example, the carbonate in shells of marine organisms (e.g. foraminifera) and
water in ice caps contain two isotopes of oxygen, 18 O and 16 O. The normalized
isotope ratio δ 18 O is calculated as a deviation from a reference sample as
δ
18 O =
(
18 O
16 O
) sample − (
18 O
16 O
) reference
(
18 O
16 O
) reference
,
(16.1)
where the reference sample is different for ice cores (i.e. standard mean ocean
water) than for carbonate shells (i.e. a specific fossil Cretaceous species). The
isotope 16 O is lighter than 18 O so that water containing 16 O is preferentially evaporated and a temperature-dependent fractionation occurs. Under cold conditions,
less water containing 18 O is able to evaporate into the atmosphere.
Changes in δ 18 O reflect the combined effect of changes in global ice volume
and temperature at the time of deposition of the sampled material. During very
cold conditions, global ice volume is relatively large and hence sea level is low,
which enriches water in the ocean with 18 O. Also because of the colder temperatures, more 18 O remains in the ocean and less 18 O becomes locked in the ice.
Hence, in ocean sediments the ratio δ 18 O will increase under cold conditions,
whereas in ice cores it will decrease.
When corrections for global ice volume (with respect to the reference sample)
are made, δ 18 O can be used as an indicator for the temperature at the time of