
in geophysical dynamics. From the computational
point of view, however, the Boussinesq equations
are still not accessible.
Owing to the difference of sizes of the vertical and
horizontal dimensions, both in the atmosphere and in
the ocean (10–20 km versus several thousands of
kilometers), the second approximation is based on the
smallness of the vertical length scales with respect to
the horizontal length scales, that is, oceans (and the
atmosphere) compose very thin layers. The scale
analysis ensures that the dominant forces in the
vertical-momentum equation come from the pressure
gradient and the gravity. This leads to the so-called
hydrostatic approximation, which amounts to repla-
cing the vertical component of the momentum equa-
tion by the hydrostatic balance equation, and hence
leading to the well-accepted primitive equations (PEs)
(Washington and Parkinson 1986). As far as we
know, the primitive equations were first considered
by L F Richardson (1922); when it appeared that
they were still too complicated they were left out
and, instead, attention was focused on even simpler
models, the geostrophic and quasigeostrophic mod-
els, considered in the late 1940s by J von Neumann
and his collaborators, in particular J G Charney.
With the increase of computing power, interest
eventually returned to the PEs, which are now the
core of many global circulation models (GCMs) or
ocean global circulation models (OGCMs), avail-
able at the National Center for Atmospheric
Research (NCAR) and elsewhere. GCMs and
OGCMs are very complex models which contain
many components, but still, the PEs are the central
component for the dynamics of the air or the water.
Further approximations based on the fast rotation
of the Earth implying the smallness of the Rossby
number lead to the quasigeostrophic and goes-
trophic equations (Pedlosky 1987).
The mathematical study of the PEs was initiated by
Lions, Temam, and Wang in the early 1990s. They
produced a mathematical formulation of the PEs
which resembles that of the Navier–Stokes due to
Leray, and obtained the existence, for all time, of weak
solutions (see Lions et al. 1992a, b, 1993, 1995).
Further works conducted during the 1990s have
improved and supplemented these early results bring-
ing the mathematical theory of the PEs to that of the
three-dimensional incompressible Navier–Stokes
equations (Constantin and Foias 1998, Teman 2001).
In summary, the following results are now available
which will be presented in this article:
1. existence of weak solutions for all time;
2. existence of strong solutions in space dimension
three, local in time;
3. existence and uniqueness of a strong solution in
space dimension two, for all time; and
4. uniqueness of weak solutions in space dimension
two.
The PEs of the Ocean
The ocean is made up of a slightly compressible
fluid subject to a Coriolis force. The full set of
equations of the large-scale ocean are the following:
the conservation of momentum equation, the con-
tinuity equation (conservation of mass), the thermo-
dynamics equation, the equation of state and the
equation of diffusion for the salinity S:
dV
3
dt
þ 2 W V
3
þr
3
p þ g ¼ D ½1
d
dt
þ div
3
V
3
¼ 0 ½2
dT
dt
¼ Q
T
½3
dS
dt
¼ Q
S
½4
¼ f ðT; S; pÞ½5
Here V
3
is the three-dimensional velocity vector,
V
3
= (u, v, w), , p, T are respectively, the density,
pressure, and temperature, and S is the concentra-
tion of salinity; g = (0, 0, g) is the gravity vector, D
the molecular dissipation, Q
T
and Q
S
are the heat
and salinity diffusions, respectively.
Remark 1 The equation of state for the oceans is
derived on a phenomenological basis. Only empirical
forms of the function f (T, S, ) are known (see
Washington and Parkinson (1986)). It is natural,
however, to expect that decreases if T increases and
that increases if S increases. The simplest law is
¼
0
ð1
T
ðT T
r
Þþ
S
ðS S
r
ÞÞ ½6
corresponding to a linearization around reference
values
0
, T
r
, S
r
of respectively, the density, tem-
perature, and the salinity,
T
and
S
are positive
expansion coefficients.
The Mach number for the flow in the ocean is not
large and, therefore, as a starting point, we can
make the so-called Boussinesq approximation in
which the density is assumed constant, =
0
,
except in the buoyancy term and in the equation of
state. This amounts to replacing [1], [2] by
0
dV
3
dt
þ 2
0
V
3
þr
3
p þ g ¼ D ½7
div
3
V
3
¼ 0 ½8
Furthermore, since for large-scale ocean, the horizon-
tal scale is much larger than the vertical one, a scale
analysis (Pedolsky 1987) shows that @p=@z and g are
Geophysical Dynamics 535