Quasigeostrophic equations
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While geostrophic motion refers to the wind that would result from an exact balance between the Coriolis force and horizontal pressuregradient forces,^{[1]} quasigeostrophic (QG) motion refers to flows where the Coriolis force and pressure gradient forces are almost in balance, but with inertia also having an effect. ^{[2]}
Contents
Origin
Atmospheric and oceanographic flows take place over horizontal length scales which are very large compared to their vertical length scale, and so they can be described using the shallow water equations. The Rossby number is a dimensionless number which characterises the strength of inertia compared to the strength of the Coriolis force. The quasigeostrophic equations are approximations to the shallow water equations in the limit of small Rossby number, so that inertial forces are an order of magnitude smaller than the Coriolis and pressure forces. If the Rossby number is equal to zero then we recover geostrophic flow.
The quasigeostrophic equations were first formulated by Jule Charney.^{[3]}
Derivation of the singlelayer QG equations
In Cartesian coordinates, the components of the geostrophic wind are
 (1a)
 (1b)
where is the geopotential.
The geostrophic vorticity
can therefore be expressed in terms of the geopotential as
 (2)
Equation (2) can be used to find from a known field . Alternatively, it can also be used to determine from a known distribution of by inverting the Laplacian operator.
The quasigeostrophic vorticity equation can be obtained from the and components of the quasigeostrophic momentum equation which can then be derived from the horizontal momentum equation
 (3)
The material derivative in (3) is defined by
 (4)
 where is the pressure change following the motion.
The horizontal velocity can be separated into a geostrophic and an ageostrophic part
 (5)
Two important assumptions of the quasigeostrophic approximation are

 1. , or, more precisely .
 2. the betaplane approximation with
The second assumption justifies letting the Coriolis parameter have a constant value in the geostrophic approximation and approximating its variation in the Coriolis force term by .^{[4]} However, because the acceleration following the motion, which is given in (1) as the difference between the Coriolis force and the pressure gradient force, depends on the departure of the actual wind from the geostrophic wind, it is not permissible to simply replace the velocity by its geostrophic velocity in the Coriolis term.^{[4]} The acceleration in (3) can then be rewritten as
 (6)
The approximate horizontal momentum equation thus has the form
 (7)
Expressing equation (7) in terms of its components,
 (8a)
 (8b)
Taking , and noting that geostrophic wind is nondivergent (i.e., ), the vorticity equation is
 (9)
Because depends only on (i.e., ) and that the divergence of the ageostrophic wind can be written in terms of based on the continuity equation
equation (9) can therefore be written as
 (10)
The same identity using the geopotential
Defining the geopotential tendency and noting that partial differentiation may be reversed, equation (10) can be rewritten in terms of as
 (11)
The righthand side of equation (11) depends on variables and . An analogous equation dependent on these two variables can be derived from the thermodynamic energy equation
 (12)
where and is the potential temperature corresponding to the basic state temperature. In the midtroposphere, ≈ .
Multiplying (12) by and differentiating with respect to and using the definition of yields
 (13)
If for simplicity were set to 0, eliminating in equations (11) and (13) yields ^{[5]}
 (14)
Equation (14) is often referred to as the geopotential tendency equation. It relates the local geopotential tendency (term A) to the vorticity advection distribution (term B) and thickness advection (term C).
The same identity using the quasigeostrophic potential vorticity
Using the chain rule of differentiation, term C can be written as
 (15)
But based on the thermal wind relation,
 .
In other words, is perpendicular to and the second term in equation (15) disappears.
The first term can be combined with term B in equation (14) which, upon division by can be expressed in the form of a conservation equation ^{[6]}
 (16)
where is the quasigeostrophic potential vorticity defined by
 (17)
The three terms of equation (17) are, from left to right, the geostrophic relative vorticity, the planetary vorticity and the stretching vorticity.
Implications
As an air parcel moves about in the atmosphere, its relative, planetary and stretching vorticities may change but equation (17) shows that the sum of the three must be conserved following the geostrophic motion.
Equation (17) can be used to find from a known field . Alternatively, it can also be used to predict the evolution of the geopotential field given an initial distribution of and suitable boundary conditions by using an inversion process.
More importantly, the quasigeostrophic system reduces the fivevariable primitive equations to a oneequation system where all variables such as , and can be obtained from or height .
Also, because and are both defined in terms of , the vorticity equation can be used to diagnose vertical motion provided that the fields of both and are known.
References
 ^ Phillips, N.A. (1963). “Geostrophic Motion.” Reviews of Geophysics Volume 1, No. 2., p. 123.
 ^ Kundu, P.K. and Cohen, I.M. (2008). Fluid Mechanics, 4th edition. Elsevier., p. 658.
 ^ Majda, Andrew; Wang, Xiaoming (2006). Nonlinear Dynamics and Statistical Theories for Basic Geophysical Flows. Cambridge University Press. p. 3. Retrieved 24 September 2017.
 ^ ^{a} ^{b} Holton, J.R. (2004). Introduction to Dynamic Meteorology, 4th Edition. Elsevier., p. 149.
 ^ Holton, J.R. (2004). Introduction to Dynamic Meteorology, 4th Edition. Elsevier., p. 157.
 ^ Holton, J.R. (2004). Introduction to Dynamic Meteorology, 4th Edition. Elsevier., p. 160.