Numerical Modeling of Weather and Climate

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– 3.1 –

Numerical Modeling of Weather and Climate

Christoph Schär, Institute for Atmospheric and Climate Science ETH, 8092 Zürich http://www.iac.ethz.ch/staff/schaer

Chapter 3:

Adiabatic Formulation of Atmospheric Models

April 18, 2007

3.1 Shallow water dynamics ... 3.3

3.1.1 Background ... 3.3

3.1.2 Integration in advective form using centered differences ... 3.5

3.1.3 Integration in flux form using the Euler time step ... 3.6

3.2 Quasi-geostrophic dynamics... 3.8

3.3 Coordinate transformation... 3.10

3.3.1 Motivation... 3.10

3.3.2 Wind vector and advection operator in generalized coordinates 3.12

3.3.3 Transformation of derivatives with generalized coordinates... 3.13

3.3.4 Transformation of the horizontal momentum equation... 3.14

3.4 Hydrostatic dynamics in pressure coordinates ... 3.16

3.5 Isentropic coordinates... 3.18

3.5.1 Background ... 3.18

3.5.2 Isentropic form of governing equations... 3.19

3.5.3 Numerical integration... 3.21

3.6 Sigma coordinates ... 3.25

3.6.1 Transformation of hydrostatic equations in σ-coordinates... 3.25

3.6.2 Summary of equations and boundary conditions... 3.26

3.6.3 Numerical integration... 3.28

3.6.4 Additional aspects of terrain-following coordinates ... 3.30

3.7 Spectral methods... 3.32

3.7.1 Classical spectral methods ... 3.32

3.7.2 Pseudo-spectral method ... 3.35

3.7.3 Global spectral models ... 3.36

3.8 References ... 3.39

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– 3.2 –

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– 3.3 –

3. Adiabatic formulation of atmospheric models

3.1 Shallow water dynamics

3.1.1 Background

The shallow water equations represent the simplest model problem for stratified fluids. They are employed in atmospheric and oceanic sciences to investigate elementary properties of stratified fluid dynamics.

The shallow water equations represent a homogeneous layer of incompressible fluid of constant density, which is confined above by a free surface. The free surface represents a discontinuous density interface, and it implies stratification effects. The one-dimensional shallow-water system is sketched in Fig.3.1.1. We make two important assumptions: First, the interaction with the overlying layer of fluid (e.g. air) is neglected. Second, it is assumed that the horizontal velocity is independent of height, i.e.

∂u/∂z=0. The latter assumption restricts

the validity of the system to shallow fluid layers.

h

H !

! u < ! u

Fig.3.1.1: Variables in the one-dimensional shallow water equations:

horizontal velocity u, layer depth H, height of topography h.

In the one-dimensional case and in the absence of background rotation, the shallow-water equations encompass a horizontal momentum equation

Du

Dt + g

*

! ( h + H )

!x = 0 with D Dt = !

!t + u !

! x , (3.1.1)

and a continuity equation

!H

!t + !(uH )

!x = 0 . (3.1.2)

Here

!

g* =g"

# # with

!

"

#

=

#

$

#

u

denotes the reduced gravity. For a lake or river (where ρ and ρ

u

represent water and air, respectively), we have g*=g. However, the shallow water equations may be used to represent a wide range of other two-layer systems. For instance, an atmospheric two-layer structure may be represented as

!

g*=g"

# # , where

Δ

θ denotes the

potential temperature contrast across the interface between the two layers.

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– 3.4 –

The shallow water equations support propagating shallow-water waves. In the approximation considered, they are non-dispersive and have phase (and group) velocities given by

c = g

*

H (3.1.3)

An associated non-dimensional parameter is the Froude number Fr = u

g

*

H , (3.1.4)

which is defined as the ratio between advective and phase velocities. The Froude number is used to distinguish between two major flow regimes. Flows with Fr < 1 are referred to as subcritical. In these flows, the group velocity exceeds the advective velocity. Thus, waves are able to propagate against the mean flow. Flows with Fr > 1 are referred to as supercritical. In supercritical flow, upstream propagating waves are washed downstream by the advective velocity. Thus, there is a kind of communication problem, in the sense that the presence of an obstacle in a supercritical flow cannot be communicated into the upstream direction.

Supercritical flows can thus lead to the formation of a hydraulic jump (or shock). In analytical theory, these are discontinuities in layer depth and flow velocity. Examples of such features in the shallow water systems include hydraulic jumps in rivers and kitchen sinks (kayak sportsmen and cooks should know about these!). Related phenomena occur in supersonic booms (aviation), foehn winds and downslope windstorms (meteorology), or traffic jams (transportation). In all these examples, the key issue is the inability of wave propagation against the mean flow.

For further analysis, we cast (3.1.1-2) into a dimensionless form by choosing the following scales:

• a horizontal scale L (e.g. defined by the half width of an obstacle),

• a vertical scale H

o

(usually defined by the mean water depth), and

• a velocity scale (defined by the phase speed g

*

H

o

).

The resulting dimensionless system reads Du

Dt + ! ( h + H )

!x = 0 with D Dt = !

!t + u !

! x , (3.1.5)

!H

!t + !(uH )

!x = 0 , (3.1.6)

and the Froude number simplifies to Fr = u H .

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– 3.5 –

3.1.2 Integration in advective form using centered differences

A simple scheme to integrate the shallow-water equations uses centered differences in space and time. The accuracy of the scheme can be improved by using a staggered grid defined by

H

nj

= H x ( = j !x, t = n !t ) , (3.1.10)

u

nj+1/ 2

= u x ( = ( j + 1 / 2) !x, t = n !t ) . (3.1.11)

Thus, the grid has the following structure:

ui!1/ 2 ui+1/ 2 ui+3/ 2

Hi!1 Hi Hi+1 Hi+2

H[i-1] H[i] H[i+1] H[i+2]

u[i] u[i+1] u[i+2]

In this staggered grid, the location of the u-grid is shifted by half a grid increment in the x- direction. The two grids points are commonly referred to as mass and velocity grids, respectively. When implementing the code on a staggered grid, care is required regarding the staggered indices. In analytical terminology, we use the notation with half integers (

!

ui+1 / 2

), while in the code itself only integer array indices are allowed (u[i+1]=

!

ui+1 / 2

).

The discretization of (3.1.5-6) on this grid, using centered differencing, yields

!

1

2"t

[

un+1j+1/ 2#un#1j+1/ 2

]

+ u2"xnj+1/ 2

[

unj+3 / 2#unj#1/ 2

]

+ 1

"x

[ (

Hnj+1+hj+1

)

#

(

Hnj +hj

) ]

= 0

, (3.1.12)

!

1

2"t [ H

n+1j

# H

n#1j

] + 2 1 "x [ u

nj+1

H

nj+1

# u

nj#1

H

nj#1

] = 0

with u

nj

= 1

2 ( u

nj#1/ 2

+ u

nj+1/ 2

)

, (3.1.13)

This is an explicit scheme, i.e. (3.1.12-13) can be solved for u

n+1

and H

n+1

.

The scheme (3.1.12-13) is stable, but in complex circumstances (e.g. hydraulic jumps),

additional steps to preserve stability are needed. This may include an Asselin filter to suppress

the computational mode of centered time step, or some digital diffusion to suppress nonlinear

instability (see Schär, 2006, chapters 4.3 and 5).

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– 3.6 –

3.1.3 Integration in flux form using the Euler time step

The scheme (3.1.12-13) is well suited for a range of purposes (e.g. wave propagation) but it has some serious deficiencies that become important in nonlinear flows. In particular, the scheme conserves mass, but the momentum equation is not in conservative flux form, and this implies that momentum can unphysically be created or destroyed.

To suppress such artifacts, we write (3.1.5) into its conservative flux form. To this end, we start with the product rule from differential calculus, i.e. ! ( Hu ) !t = H ( !u !t ) + u ( !H !t ) ,

and substitute from (3.1.5-6) to obtain

!

(

Hu

)

!t + !

[

u Hu

( ) ]

!x + H!

(

h+H

)

!x =0

. (3.1.20)

In absence of topography ( h ! 0 ), the third term may be written as

!

"(H2/2) "x

, and this shows that (3.1.20) is in momentum conservative flux form. In particular, integration of (3.1.20) over an interval [a,b] yields

!

!t Hudx

a b

" = # [ Hu

2

+ H

2

2 ]

ab

, (3.1.21)

This shows that the total momentum in this interval can only change in response to momentum fluxes at the boundaries. In the case of topography ( h ! 0 ) however, there is some momentum exchange with the underlying surface by pressure forces. This momentum exchange is the shallow-water analog of the mountain drag (or orographic drag), which is an important term in the atmospheric momentum balance. However, even in this case, the conservative flux form (3.1.20) is preferable, as it describes the physics of the system more realistically than its advective counterpart (3.1.5).

I

NTEGRATION OF THE CONTINUITY EQUATION

The continuity equation (3.1.6) is integrated using the first-order flux form of the upstream scheme (see Schär 2006, chapter 6.3).

!

Hin+1 " Hin

#t + Fi+1/ 2n+1/ 2 " Fi"1/ 2n+1/ 2

#x = 0

(3.1.23)

with the upstream fluxes estimated as

!

Fi+1 / 2n+1 / 2 =

ui+1/ 2n Hin+1 for ui+1/ 2n " 0 ui+1/ 2n Hin for ui+1/ 2n # 0

$ %

&

. (3.1.24)

On the computer, this is implemented as

!

Fi+1/ 2n+1/ 2 = Max

(

0,ui+1/ 2n

)

Hin + Min

(

0,ui+1/ 2n

)

Hi+1n

(3.1.25)

in order to avoid a case distinction (if statement). The main advantage of the staggered grid is

that it allows one to employ finite differences over Δ x (rather than 2 Δ x) in (3.1.24).

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– 3.7 – I

NTEGRATION OF THE

M

OMENTUM EQUATION

With the momentum variable q:=Hu, (3.1.20) is written as

!

"q

"t + "

( )

uq

"x + H"

(

h+H

)

"x =0

. (3.1.26)

As a first step, we need to diagnose the momentum q at the mass points from the other variables as

!

q

in

:= H

in

u

in"1/ 2

+ u

i+1/ 2n

2 (3.1.27)

The simplest discretization of (3.1.26) using the flux-form of the upstream scheme then reads

!

q

in+1

" q

in

#t + Q

i+1/ 2n+1/ 2

" Q

i"1/ 2n+1/ 2

#x = H

in

h

i+1

" h

i"1

( ) + ( H

i+1n

" H

i"1n

)

2#x (3.1.28)

where, analogous to (3.1.25), the upstream fluxes

!

Qi+1/ 2n+1/ 2 = Max

(

0,ui+1 / 2n

)

qin + Min

(

0,uin+1 / 2

)

qi+1n

(3.1.29)

are employed. Once

!

qin+1

has been computed, one diagnoses the “advective velocities” on the staggered grid as

!

uin+1+1 / 2 = qin+1+qin+1+1

Hin+1+Hi+1n+1

. (3.1.30)

This procedure is referred to as momentum averaging.

The procedure outlined above has some attractive properties. In particular, we have employed

the conservative flux form, which is well suited to treat hydraulic jumps. However, the

procedure also has some disadvantages. First, the use of the upstream fluxes yields a shallow-

water scheme merely of first-order accuracy. Second, the treatment of the pressure gradient

term is not optimal. As it employs differences over 2 Δ x, spurious small-scale oscillations may

result, in particular near hydraulic jumps. Both these disadvantages can be addressed using

more sophisticated approaches.

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– 3.8 –

3.2 Quasi-geostrophic dynamics

As a second example we consider the quasi-geostrophic set of equations in the Boussinesq approximation and in "pseudo-height" coordinates. This set of equations represents a simplified idealized dynamical system, which is – due to its conceptual elegance – commonly used for research purposes in large-scale geophysical dynamics. For an introduction into the quasi-geostrophic dynamics, see Holton (2004).

The system is described by few quantitites which are all closely related to the geopotential φ : the geostrophic wind

(u

g

,v

g

)= f

!1

(!"

y

,"

x

) , (3.2.1)

the quasi-geostrophic potential vorticity q = !

xx

+ !

yy

+ f

2

N

2

!

z

"

#

$ %

&

'

z

(3.2.2) and the potential temperature

!

o

+ !(z) + " (x, y , z, t) . (3.2.3)

The latter is split into 3 terms: The first term represents a constant value that is usually associated with the Earth’s surface temperature. The second term defines the vertical temperature profile and the Brunt-Väisälla frequency

N (z )= g

!

o

"!

"z

#

$

% &

' (

1/2

. (3.2.4)

The third term finally defines the spatial and temporal variations around the background profile. It is linked to the geopotential by the hydrostatic relation

!

= "o g

#

$

#z

. (3.2.5)

The quasi-geostrophic dynamics in a domain 0 ! z ! z

t

is governed by the conservation of quasi-geostrophic potential vorticity inside the domain, i.e.

!

D

g

q

D t = 0 for 0 < z < z

t

, (3.2.6)

and the conservation of potential temperature on the bounding surfaces, i.e.

!

D

g

"

D t =0 for z =0,z

t

. (3.2.7)

In the latter two equations, the advection operator accounts for the geostrophic wind

Dg

D t = !

!t + ug !

!x + vg !

!y

. (3.2.8)

while neglecting vertical advection.

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– 3.9 –

A simple numerical scheme to integrate the quasi-geostrophic system involves the following steps:

(i) Prognosis of q inside the domain using (3.2.6).

(ii) Prognosis of θ on the bounding surfaces z = 0, z

t

using (3.2.7).

(iii) Quasi-geostrophic inversion: This involves solving the boundary value problem formed by (3.2.2) and (3.2.5), i.e.

!

"

xx

+ "

yy

+ f

2

N

2

"

z

#

$

% %

&

' ( (

z

= q for 0 ) z ) z

t

"

z

= g*

+

o

for z =0,z

t

(3.2.9)

with the quantities on the right-hand sides provided by the prognostic steps (i) and (ii).

Equation (3.2.9) is an elliptic problem, and its solution involves iterative procedures.

(iv) Computation of the geostrophic wind from the geopotential φ using (3.2.1).

In distinction to the shallow-water equations, solving the quasi-geostrophic systems involves

both prognostic (i,ii) and diagnostic steps (iii, iv). The latter do not include any time stepping

but rather aim at the consistent diagnosis of φ , u

g

and v

g

.

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– 3.10 –

3.3 Coordinate transformation

3.3.1 Motivation

The proper choice of vertical and horizontal coordinates in atmospheric and oceanic models is an important topic. There are a large number of choices, and the optimal approach depends upon a number of factors (dynamical approximation, computational domain, model resolution, etc). Of particular importance is the choice of the vertical coordinate system, due to the presence of underlying topography. Some vertical coordinate systems are sketched in Fig.3.3.1. In absence of topography (or bathymetry in the case of the ocean), the height coordinate (z-coordinate system) is commonly used in non-hydrostatic models, while the pressure coordinate (p-coordinate) is optimal for idealized hydrostatic models.

There are two major principles that guide the choice of the vertical coordinate:

(1) A clever choice of the vertical coordinate allows the governing equations to be greatly simplified. The choice depends upon the dynamical approximation. For the hydrostatic dynamics, the pressure coordinate (see Fig.3.3.1b) is optimal, as it leads to a considerable simplification of the governing equations. For instance, the continuity equation becomes linear in pressure coordinates, i.e.

!u

!x

"

# $

$

%

&

' '

p

+ !v

!y

"

# $

$

%

&

' '

p

+!

(

!p=0

. (3.3.1)

In pressure coordinates, vertical motions are described by the vertical velocity

!

= Dp Dt

(measured in hPa/s) and horizontal derivatives are taken on surfaces of constant pressure (see later). In comparison, the corresponding equation in z- coordinates (Fig.3.3.1a) reads

!

"

!t +!

( ) "

u

!x +!

( ) "

v

!y +!

( "

w

)

!z =0

. (3.3.2)

The latter equation has an addition term (∂ ρ /∂t) and is highly nonlinear.

(2) The coordinate system should allow for the definition of a simple numerical grid. The aforementioned pressure coordinate has an inconvenient disadvantage. The atmospheric variables are defined in the domain

p

s

(t, x, y) > p > 0 ,

where p

s

refers to surface pressure. In (x,y,p) space, the computational domain has an

irregular and time-dependent lower boundary condition, as pressure surfaces intersect

the ground, in particular in the presence of topography (Fig.3.3.1b). A classical

approach to this problem is the terrain-following sigma-coordinate (Fig.3.3.1c) It is

defined as

(11)

– 3.11 –

!

:= p

ps

. (3.3.3)

With this choice, grid points are located on surfaces defined by σ = const. The coordinate σ runs from σ =0 (at the top of the atmosphere defined by p=0) to σ =1 (at the Earth’s surface defined by p= p

s

). This yields a regular computational domain defined by

1 > ! > 0 .

The deformation of σ -coordinates that results from underlying topography decays with height. Sigma-coordinates are also suited for oceanic models. However, to avoid collapsing computational surfaces near coast lines, vertical walls must be introduced (Fig.3.3.1d). This implies neglecting the continental shelf.

! = 0

! = 1 p = ps p = 0

! = const

p = const z = const

(a)

(c)

(b)

(d)

Fig.3.3.1: Examples of vertical coordinate systems: (a) z-coordinates, (b) p-coordinates, (c) σ-coordinates in atmospheric models, (d) σ-coordinates in oceanic models.

In addition to the coordinate systems considered above, there is a large variety of additional coordinates:

• Isentropic coordinate systems adopt potential temperature θ as vertical coordinate (see section 3.5). This system is particularly attractive for idealized dynamical studies, and closely related to isosteric coordinates (where density ρ serves as vertical coordinate).

• Hybrid coordinates typically use terrain-following σ -coordinates at low levels, with a smooth transition to p-coordinates and/or θ -coordinates at upper levels.

• In the context of finite-volume atmospheric models, there is some recent research

about variants of z-coordinates (e.g. the shaved-cell approach).

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– 3.12 –

Coordinate transformations are also important in order to represent the spherical geometry of the planet. On a sphere, a simple Cartesian coordinate system does not exist. Common meteorological convention defines a local Cartesian (x,y,z), where the z-axis points upwards, and the x-axis towards east. This coordinate system is closely related to the spherical ( λ , φ ,z)- system. In these lecture notes we will – except for section 3.7 – restrict attention to the f-plane approximation and thus mostly work with a Cartesian

!

(x,y, ˜ z )

-system where

!

z ˜ denotes a generalized vertical coordinate. However, the proper representation of the spherical geometry is an important topic, both in global and regional atmospheric and oceanic models.

The implementation of an atmospheric or oceanic model in generalized coordinates

!

( ˜ x , ˜ y , ˜ z )

involves two steps. In a first step, the governing equations need to be transformed into the new coordinate system. In a second step, a discretization is introduced, for instance by using grid increments

!

("x ˜ ,"˜ y ,"˜ z ) in the new coordinate directions.

3.3.2 Wind vector and advection operator in generalized coordinates

We consider the transformation of a Cartesian coordinate system (x,y,z) into a (generally non- Cartesian) coordinate system

!

( ˜ x , ˜ y , ˜ z )

. To this end, we first need a precise definition of the wind vector. In Cartesian coordinates (x,y,z) the definition reads

!

u := Dx

Dt , v := Dy

Dt , w := Dz

Dt , (3.3.10a)

and it can readily be generalized to an arbitrary (curvilinear) coordinate system

!

( ˜ x , ˜ y , ˜ z )

:

!

u ˜ := x

Dt , ˜ v := y

Dt , ˜ w := z

Dt . (3.3.10b).

The wind vector is thus defined by the rate of change with time of the air parcel’s location.

The advection operator

D Dt

appearing in (2.2.10) is the usual total derivative used in fluid dynamics.

According to (3.3.10b), the vertical wind in

!

( ˜ x , ˜ y , ˜ z )=(x,y,p)

coordinates is defined as

! : = D p

D t , (3.3.11)

and measured in [hPa/s]. Note that it is negative (positive) for ascending (descending) motions. Similarly, the vertical wind in isentropic coordinates

!

( ˜ x , ˜ y , ˜ z )=(x,y,

"

)

is defined as

!

. := D

!

D t

. (3.3.12)

and measured in [K/s]. Note that for this system !

. "0

implies the presence of diabatic

processes, as for the adiabatic dynamics the potential temperature is conserved following the

motion of the flow, i.e.

D

!

D t"0

. For adiabatic flows, the vertical velocity component thus

vanishes, i.e. v=(u,v,0).

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– 3.13 –

The advection operator can be formed in generalized coordinates using the total derivative. In the

(x,y,z)

-system this reads

!

D

D t =

"

"

t + D x

D t

"

"

x

#

$ % &

' (

y,z

+ D y

D t

"

"

y

#

$ % &

' (

x,z

+ D z D t

"

"

z

#

$ % &

' (

x,y

=

"

"

t + u

"

"

x

#

$ % &

' (

y,z

+ v

"

"

y

#

$ % &

' (

x,z

+ w

"

"

z

#

$ % &

' (

x,y

(3.3.13a)

and in the generalized

!

( ˜ x , ˜ y , ˜ z )

-system

!

D

D t =

"

"

t + Dx ˜

D t

"

"

x ˜

#

$ % &

' (

y ˜ ,˜ z

+ Dy ˜ D t

"

"

y ˜

#

$ % &

' (

x ˜ ,˜ z

+ Dz ˜ D t

"

"

z ˜

#

$ % &

' (

x ˜ , ˜ y

=

"

"

t + ˜ u

"

"

x ˜

#

$ % &

' (

y ˜ ,˜ z

+ ˜ v

"

"

y ˜

#

$ % &

' (

x ˜ ,˜ z

+ ˜ w

"

"

z ˜

#

$ % &

' (

x ˜ , ˜ y

(3.3.13b)

Here we have used the usual convention of partial differential calculus and employ subscripts to define the coordinates that are held constant when forming derivatives. In general, all subscripts are suppressed unless they are needed to avoid confusion. For instance, when considering the two systems

(x,y,z)

and

!

( ˜ x , ˜ y , ˜ z )=(x,y,p)

, we simplify the notation as follows:

!

" "x

( )

p:=

(

" "x

)

y,p

and

!

" "p:=

(

" "p

)

x,y

.

3.3.3 Transformation of derivatives with generalized coordinates

Most coordinate systems of relevance to atmospheric and oceanic models belong to the class of curvilinear coordinates. Simple examples of curvilinear coordinates include polar or spherical coordinates. The definition of “curvilinear” requires that these coordinates are locally invertible and differentiable, but the coordinates surfaces may be curved and non- Cartesian.

When transforming the governing system of equations from Cartesian to some curvilinear coordinates, we need to eliminate the coordinates (x,y,z) and to express the derivatives in terms of the new coordinates. For instance, when considering the generalized three- dimensional coordinate transformation, we need to express the derivative

!

(" "x)

y,z

in terms of

!

(" " x ˜ )

y ˜ z

,

!

(" " y ˜ )

x ˜ z

and

!

(" " z ˜ )

x ˜ , ˜ y

. In the general case, this task requires tensor calculus (e.g. Fletcher 1991).

However, here we restrict our attention to transformations of the vertical coordinate assuming horizontal Cartesian coordinates (x,y) in f-plane geometry. In this case, the coordinate transformations are much simpler and tensor calculus is not needed. In this simplified framework, a new vertical coordinates may be defined by a function

!

s=s(x,y,z)

.

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– 3.14 –

At each location (x,y), the function s(z) must be strictly monotonic in z to guarantee an invertible coordinate system. To address the associated transformation

!

(x,y,z)"(x,y,s), we

consider derivatives of some quantity χ = χ (x,y,z) = χ (x,y,s).

For vertical derivatives, the transformation from the z- to the s-system can be accomplished using the chain rule i.e.

! "

!s = ! "

!z

! z

!s , (3.3.20)

where the factor

!z !s

can be determined from s=s(x,y,z).

For horizontal derivatives, we need to represent the fact that horizontal derivatives are taken on z- and s-surfaces, respectively. To this end, consider variations of χ (x,y,s) in an environment around (x

o

,y

o

,s

o

). Wandering around in this environment with

!

(

"

x#0,

"

y=0,

"

s)

, implies variations of χ given by

!" = # "

# x

$

% &

&

' ( ) )

s

!x + # "

# s

$

% &

&

' ( ) ) !s

.

Here the subscript s denotes a derivative on an s-surface. Division by δ x yields

!"

!x = # "

# x

$

% &

&

' ( ) )

s

+ # "

# s

! s

!x

$

% & ' ( ) .

Next we assume that δ x and δ s depend upon each other such that δ z = 0; this means we restrict our attention to walk tours on z-surfaces. It follows that δ / δ x may be replaced by

!

(" /"x)

z

, i.e.

! "

! x

#

$

% &

'

z

= ! "

! x

#

$

% &

'

s

+ ! "

! s

! s

! x

#

$

% &

'

z

(3.3.21)

This equation allows converting a horizontal derivative on a z-surface by a horizontal derivative on an s-surface.

Equations (3.3.20) and (3.3.21) are completely general. We have not made any assumptions, not even that (x,y,z) is a Cartesian system. It follows that s and z can be arbitrarily exchanged with each other, or replaced by other symbols (e.g. by p or θ ). The only important assumption is that the transformation

!

s=s(x,y,z) is invertible and continuous.

3.3.4 Transformation of the horizontal momentum equation

As example we next consider the transformation of the hydrostatic momentum equations in pressure coordinates on an f-plane. The respective equations in z-coordinates read

Du

Dt ! fv = !1

"

#p

#x

(3.3.22a)

(15)

– 3.15 –

Dv

Dt + fu = !1

"

#p

#y

(3.3.22b)

Here p denotes pressure, ρ is density, and

!

f =2"sin(

#

) denotes the Coriolis-parameter with

!

"=2

#

24h

the angular velocity of the Earth’s rotation and φ the geographical latitude.

The full form of the advection operator reads

!

D

Dt = "

"t+u "

"x +v "

"y +w "

"z

(3.3.23)

with

u=(u,v,w)

representing the three-dimensional velocity vector. As the transformation considered does not affect (x,y), the horizontal velocity components (u,v) are unchanged.

However, in place of the Cartesian vertical wind w [m/s], we use the vertical wind “omega”

!

"

=Dp Dt

[hPa/s] as discussed in section 3.3.2. Thus, the left-hand side of equation (3.3.22) remains unchanged, except for a modified form of the advection operator

!

D

Dt = "

"t+u "

"x

#

$ % &

' (

p

+v "

"y

#

$ % &

' (

p

+

)

"

"p

(3.3.24)

In order to transform the pressure gradient terms on the right-hand side of (3.3.22), we next use (3.3.21), where we replace z by p and s by z, i.e.

! "

!

x

#

$ %

%

&

' ( (

p

=

! "

!

x

#

$ %

%

&

' ( (

z

+

! "

!

z

!

z

!

x

#

$ %

%

&

' ( (

p

. Replacing χ by p implies

!

p

!

x

"

# $

$

%

&

' '

p

=

!

p

!

x

"

# $

$

%

&

' '

z

+

!

p

!

z

!

z

!

x

"

# $

$

%

&

' '

p

(

!

p

!

x

"

# $

$

%

&

' '

z

=)

!

p

!

z

!

z

!

x

"

# $

$

%

&

' '

p

.

After using the hydrostatic relation !

p

!

z="g

# , the transformed version of (3.3.22a) follows

Du

Dt ! fv = ! g "z

"x

#

$ % &

' (

p

Here z=z(x,y,p,t) refers to the height of a pressure surface and replaces p=p(x,y,z,t) in z- coordinates. The preceding equation is usually rewritten using the definition of the geopotential φ =gz, i.e.

Du

Dt ! fv = ! "#

"x

$

% & ' ( )

p

(3.3.25) The latter form is somewhat more general, as it allows for variations of gravity, i.e.

g=g(x,y,z).

Comparison of (3.3.22a) and (3.3.25) shows that the pressure gradient term is much simpler

in the transformed system, since the density factor 1/ ρ has disappeared. This is another

important advantage of pressure coordinates over height coordinates.

(16)

– 3.16 –

3.4 Hydrostatic dynamics in pressure coordinates

We consider the atmospheric equations of motion in their hydrostatic approximation. Most numerical weather prediction models and all climate models are based on this set of equations, although non-hydrostatic atmospheric models are rapidly gaining importance in high-resolution (kilometer-scale) modeling. We consider the equations in their adiabatic / inviscid and dry form. For the moment we also neglect effects associated with the spherical nature of the Earth and consider a form of the equations that is commonly used in regional (limited-area) weather and climate models. We begin by considering the equations in pressure coordinates, and later consider the numerical integration in θ -coordinates (chapter 3.5) und σ - coordinates (chapter 3.6).

The dynamics and thermodynamics of a dry atmosphere are defined by the following three- dimensional time-dependent fields of the form χ=χ(x,y,p,t):

u, v horizontal velocity components, ω vertical velocity ( !

= Dp/ Dt

),

!

=g z

geopotential,

ρ density, and

T temperature.

and by one two-dimensional field:

!

ps(x,y,t)

surface pressure.

The governing system of equations is presented here without derivation (except for the horizontal momentum equations considered in section 3.3). Succinct derivations can be found in Holton (2004) or Gill (1982).

The governing equations include the equation of state

p=

!

RT

, (3.4.1)

the continuity equation

! u

! x

"

#

$ %

&

p

+ !v

!y

"

#

$ %

&

p

+ !'

! p = 0 , (3.4.2)

the thermodynamic equation

DT

Dt =

!

"

cp +Q

cp

, (3.4.3)

the hydrostatic equation

!

"

!p=#1

$ , (3.4.4)

and the horizontal momentum equations (3.3.25)

Du

D t !f v =! "

#

"x

$

% &

&

' ( ) )

p

+Fx

, (3.4.5a)

(17)

– 3.17 –

D v

Dt +f u =! "

#

"y

$

% &

&

' ( ) )

p

+Fy

, (3.4.5b)

All "horizontal" derivatives in (3.4.1-5) are to be taken on surfaces of constant pressure, and the advection operator is defined by

D Dt = !

!t + u !

!x

"

#

$

%

p

+ v ! dy

"

#

& $

%

p

+ ' !

!p , (3.4.6)

The non-conservative force F = ( F

x

, F

y

) and the diabatic heating rate Q relate to processes that are not explicitly represented above (moist dynamics, atmospheric radiation, turbulence, boundary layer processes, etc). For the moment we assume that these terms are externally specified or vanish. Later we will learn how F and Q are represented by parameterizations.

Note also that in the above form the governing equations allow variations of the Earth’s acceleration g (which are particularly important if a deep atmosphere is considered), but for simplicity we will subsequently assume that g=const.

Solving the governing equations (3.4.1-5) requires three boundary conditions. The first is the geometric lower boundary condition

! ( p = p

s

) = g z

s

(3.4.7)

that matches for the height z

s

= z

s

(x,y) of the underlying topography with the geopotential.

The second is the kinematic boundary condition at the top of the atmosphere, i.e.

!

"

(p=0)=0 .

(3.4.8)

It ensures that at the top of the atmosphere (p=0) any vertical motion must vanish. The third is the kinematic boundary condition at the surface, i.e.

!

"

(p= ps)=#ps #t+vs$%hps

(3.4.9)

where

!

vs =(us,vs) denotes the horizontal velocity at the surface. Equation (3.4.9) ensures

that the three-dimensional velocity vector at the surface must always be tangential to the surface. The notion of a tangential flow in the p-system is somewhat obscured by the fact that the surface pressure p=p

s

(x,y,t) may vary with time. The boundary condition can be derived from (3.3.11) after noting that p may be replace by p

s

(as the respective air parcel must remain at the surface) and that

!

"ps "p=0 (as the surface pressure ps

does not depend upon

the vertical coordinate p).

(18)

– 3.18 –

3.5 Isentropic coordinates

3.5.1 Background

In the isentropic system potential temperature

!

=T pref p

"

#

$ %

&

R cp

. (3.5.1)

serves as the vertical coordinate. Here p

ref

= 1000 hPa denotes a reference pressure, R = 287 J/(K kg) the gas constant for dry air, and c

p

= 1004 J/(K kg) the specific heat of dry air at constant pressure.

The vertical wind in isentropic coordinates is defined

!

.

: = D!

Dt (3.5.2)

and is measured in [K/s]. For adiabatic flows !

.

= D! Dt =0, i.e. the vertical wind vanishes and the flow becomes quasi-horizontal on the θ -surfaces. This is an important simplification in numerical implementations. Isentropic coordinates are thus well suited for idealized studies of adiabatic flows. However, the invertibility condition of coordinate transformations requires that !

=

!

(z)

and !

=

!

(p)

are strictly monotonic functions, i.e.

!"

!z > 0 and

!

"

!p <0

. (3.5.3)

! = ! s

! = ! t

! = const

! = const ! = const

(a)

(b) (c)

Fig.3.5.1: Vertical sections of isentropic coordinates in (a) an idealized stratified flow over a mountain, and (b,c) in a frontal zone. Panel (c) shows how the lower boundary condition may

be represented by theta-layers of vanishing thickness (see text).

(19)

– 3.19 –

These conditions are often not met in the lower troposphere or in the planetary boundary layer. Thus, isentropic coordinates are not suited for realistic weather prediction or climate models, or only in combination with other coordinates (e.g. σ -p- θ hybrid coordinate, see Zhu et al. 1992).

An additional disadvantage of isentropic coordinates is due to notorious difficulties at the lower boundary. For certain idealized problems (such as adiabatic flows past mountain ridges) the lower boundary may be represented by a surface of constant potential temperature (see Fig.3.5.1a). In general, however, atmospheric flows are baroclinic (e.g. horizontal temperature gradients in frontal zones) and the lower boundary contains some temperature contrast (Fig.3.5.1b). In these cases, near-surface isentropic layers may be represented as collapsed massless layers (see Fig.3.5.1c), which yields a difficult numerical problem.

3.5.2 Isentropic form of governing equations (A) A

DVECTIVE FORM OF MOMENTUM EQUATION

The horizontal momentum equations in pressure coordinates may be expressed as

!

Dv

Dt + f k " v = #$

p

% (3.5.6)

where

v=(u,v)

denotes the horizontal wind vector, !

p

" the pressure gradient on p-surfaces, and k the vertical unit vector. The left-hand side of this equation has the same form in isentropic coordinates, but employs the appropriate advection operator, i.e.

D

Dt = !

!t + u !

!x

"

# $ %

&

'

(

+ v !

!y

"

# $

$

%

&

' '

(

+ (

.

!

!( (3.5.7)

For the transformation of the pressure force (

!

"

!x

)

p

, we use (3.3.21) after replacing z by p, s by θ , and χ by φ . This yields

!"

!x

#

$

%

&

'

= !"

!x

#

$

%

&

p

+ !"

!p

!p

!x

#

$

%

&

'

.

After introducing ∂ φ /∂p=-1/ ρ from the hydrostatic relation (3.4.4), and substituting 1/ ρ

=RT/p

from the equation of state (3.4.1), and after rearranging the terms, one obtains

!"

!x

#

$

%

&

p

= !"

!x

#

$

%

&

'

+ RT p

!p

!x

#

$

%

&

'

.

Finally we use the definition of potential temperature (3.5.1) to eliminate pressure from the equation. Manipulation of the respective term yields

RT p

!p

!x

"

# $ %

&

' (

= RT

pref

(

T

"

# $ %

&

'

cp/R

pref !

!x

"

# $ %

&

' (

T

(

"

# $ %

&

'

cp R

= ! cpT

!x

"

#

$ $

%

&

' ' (

.

and

(20)

– 3.20 – Du

Dt ! fv = ! "M

"x

#

$

%

&

'

(3.5.8a)

Dv

Dt + fu = ! "M

"y

#

$

% &

'

(

(3.5.8b)

where

M = ! + c

p

T = gz + c

p

T (3.5.9)

denotes the Montgomery potential. The Montgomery potential in isentropic coordinates plays the same role as the geopotential in pressure coordinates.

(B) I

SENTROPIC MASS DENSITY AND CONTINUITY RELATION

The mass between two isentropic surfaces, i.e. !

dz, can be derived from the hydrostatic

relation as

!

dz = "1 g

#p

#zdz = "1 g

#p

#z

#z

#

$

d

$

= "1 g

#p

#

$

d

$ .

It is common to define the isentropic mass density (or isentropic density)

!

"

:= #1 g

$p

$

% , (3.5.11)

whereupon

!

dz =

"

d

# . (3.5.12)

Note that the isentropic density should not be mistaken for the σ -coordinate mentioned in section 3.1 and treated in section 3.6.

The isentropic density σ plays the same role as density ρ in z-coordinates. Thus, one can guess the isentropic continuity equation from the z-coordinate version (3.3.2) as

!"

!t + ! "u

! x

#

$ %

%

&

' ( (

)

+ ! "v

! y

#

$ %

%

&

' ( (

)

+ ! " )

.

! ) = 0 . (3.5.13)

The formal derivation of this equation is, however, quite cumbersome.

(C) H

YDROSTATIC RELATION

To derive the isentropic form of the hydrostatic equation, we begin with a logarithmic version of (3.5.1), i.e.

ln

!

= lnT " R

cplnp + R

cplnpref

.

Taking the derivative with respect to θ and rearranging the terms yields

cpT

!

=cp

"T

"

!

# RT

p

"p

"

! .

(21)

– 3.21 –

Finally we use the hydrostatic equation in z-coordinates

!p !z="g

# and the gas equation to obtain

c

p

T

! = "

"! ( c

p

T + gz ) .

Comparison with (3.5.9) shows that this can be expressed as c

p

T

! = "M

"! or c

p

p p

ref

!

"

# $

%

&

R cp

= 'M

'( (3.5.14)

where the second form is obtained after substitution from (3.5.1).

(D) F

LUX FORM OF MOMENTUM EQUATIONS

As with the shallow water system (chapter 3.1.3), equations (3.5.8) and (3.5.13) can be combined to derive a conservative flux from of the momentum equation. In the x-direction this is

!("u)

!t + !(u"u)

!x

#

$ % &

' (

)

+ !(v"u)

!y

#

$ %

%

&

' ( (

)

+ !()

.

"u)

! ) * f"v = * " !M

!x

#

$ % &

' (

)

. (3.5.15)

(E) B

OUNDARY CONDITIONS

As with p-coordinates, integrating the governing equations requires appropriate boundary conditions at the surface and the top of the modeling domain. In the case of an adiabatic flow problem without large-scale temperature gradients (such as in Fig.3.5.1), one can assume that the domain is confined at the surface and the model top by quasi-horizontal surfaces of constant potential temperature θ

s

and θ

t

, respectively. We may assume that the top surface is horizontal (rigid lid) and characterized by constant pressure, i.e.

p( ! = !

t

) = p

t

= const . (3.5.16)

At the lower boundary, the height of the topography determines the geopotential, as in pressure coordinates, i.e.

! ( " = "

s

) = gz

s

. (3.5.17)

3.5.3 Numerical integration

The momentum equation in advective (3.5.8) or flux form (3.5.15), the continuity equation

(3.5.13), the hydrostatic relation (3.5.14), the definition of the isentropic density (3.5.11), and

the boundary conditions (3.5.16-17) together yield a complete prognostic set of governing

equations. Here a simple numerical scheme is presented that employs ideas familiar from the

treatment of the shallow water equations in chapter 3.1. However, in addition to the horizontal

(22)

– 3.22 –

staggering, the variables are now also staggered in the vertical direction. The distribution of grid points in a three-dimensional isentropic model is sketched in Fig.3.5.2.

k = K+1/2

k = 1/2

vertical section

k = K–1/2 k = K–3/2

k = 3/2 k = K k = K–1

k = 1

horizontal section

j–2 j–1/2

j–1 j+1/2

j j+3/2

j+1

i–1 i–1/2

i i+1/2

i+1 i+3/2

i+2

u v v, !, M u

M, !, ", ", p

.

" = "s+(k–1/2)#"

", ", p

.

" = "s

"t = "s+(k–1/2)#"

="s+K#"

k = 5/2 k = 2

Fig.3.5.2: Grid points and variables in a simple three-dimensional isentropic model.

In order to simplify the situation, we will make the following additional assumptions:

• two-dimensional case in (∂/∂y=0),

• adiabatic and inviscid flow ( !

. "0

),

• neglect Earth’s rotation (

f =0

), and

• the lower boundary is an isentropic surface (

!

"

s=const

).

The associated primary model variables are

u

i+1/2, k

horizontal wind velocity,

!

i, k

isentropic mass density,

M

i, k

Montgomery potential, and

p

i, k+1/2

pressure.

Below we assume that all quantities are know at time t = n!t . Below the different phases of a time step are discussed:

(A) P

ROGNOSTIC TIME STEP FOR ISENTROPIC DENSITY AND MOMENTUM

In a first phase we address the approximated form of (3.5.8a) and (3.5.13), i.e.

(23)

– 3.23 –

!u

!t + u !u

!x = " !M

!x

#

$

%

&

'

, (3.5.21)

!"

!t + ! "u

! x

#

$

% &

'

(

= 0 . (3.5.22)

These equations exhibit a remarkable similarity with the shallow-water equations (chapter 3.1), but the isentropic system is fully two-dimensional in the (x,z)-space. Upon introducing a vertical discretization of the form

!

k

= ! " [ = "

s

+ (k # 1 2)$" ]

for all variables appearing in (3.5.21-22), the relationship with the shallow-water equations becomes even more apparent: The equations in each isentropic layer k are formally identical with the shallow-water equations (3.1.5-6). The role of the shallow-water layer depth H is played by the isentropic density σ . The only relevant difference between the two systems concerns the appearance of the Montgomery potential in the pressure term. This couples the different layers to each other, and the isentropic system can be interpreted (and treated) as a vertical stack of shallow-water equations.

The time step for each isentropic layer (3.5.21-22) can employ the same methods as used with the shallow water equations. Both the advective (chapter 3.1.2) or the flux form approach (chapter 3.1.3) may be used. This phase delivers values of σ

n+1

and u

n+1

at time n+1.

(B) D

IAGNOSIS OF PRESSURE

In distinction to the shallow-water equations, the isentropic system supports vertical coupling and wave propagation. The coupling between the different layers is instantaneously communicated vertically by the hydrostatic equation. The remaining equations (3.5.11) and (3.5.14) do not contain any time derivatives: they represent purely diagnostic equations.

Together with the boundary condition, they allow one to compute the Montgomery potential M

n+1

from the results of the previous step (i.e. σ

n+1

and u

n+1

).

To begin with, the pressure is diagnosed with the help of the isentropic density (3.5.11). A vertically discretized form of this equation, using centered differencing, reads

!

i,kn+1

= " p

i,kn+1+1/2

" p

i,kn+1"1/2

g #$ , which can be cast into the form

!

pi,k"1 / 2n+1

=pi,k+1 / 2n+1 +g#

$%

i,kn+1

for (k=K, …, 1). (3.5.24)

This equation can be evaluated from the top to the surface, starting with k=K. The uppermost

pressure value that is needed for the first step is available from the upper boundary condition

(3.5.16), i.e.

Figure

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