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-revision 1.7 by cnh,
Thu Oct 25 12:06:56 2001 UTC
+revision 1.23 by jmc,
Mon Jul 11 13:49:28 2005 UTC
 Line 37
  % $Header$
  % $Name$
- \section{Introduction}
+ This document provides the reader with the information necessary to
- This documentation provides the reader with the information necessary to
  carry out numerical experiments using MITgcm. It gives a comprehensive
  description of the continuous equations on which the model is based, the
  numerical algorithms the model employs and a description of the associated
-Line 49 
 are available. A number of examples illu
+Line 47 
 are available. A number of examples illu
  both process and general circulation studies of the atmosphere and ocean are
  also presented.
+ \section{Introduction}
+ \begin{rawhtml}
+ <!-- CMIREDIR:innovations: -->
+ \end{rawhtml}
  MITgcm has a number of novel aspects:
  \begin{itemize}
-Line 84 
 computational platforms.
+Line 88 
 computational platforms.
  \end{itemize}
  Key publications reporting on and charting the development of the model are
- listed in an Appendix.
+ \cite{hill:95,marshall:97a,marshall:97b,adcroft:97,marshall:98,adcroft:99,hill:99,maro-eta:99,adcroft:04a,adcroft:04b,marshall:04}:
+ \begin{verbatim}
+ Hill, C. and J. Marshall, (1995)
+ Application of a Parallel Navier-Stokes Model to Ocean Circulation in
+ Parallel Computational Fluid Dynamics
+ In Proceedings of Parallel Computational Fluid Dynamics: Implementations
+ and Results Using Parallel Computers, 545-552.
+ Elsevier Science B.V.: New York
+ Marshall, J., C. Hill, L. Perelman, and A. Adcroft, (1997)
+ Hydrostatic, quasi-hydrostatic, and nonhydrostatic ocean modeling
+ J. Geophysical Res., 102(C3), 5733-5752.
+ Marshall, J., A. Adcroft, C. Hill, L. Perelman, and C. Heisey, (1997)
+ A finite-volume, incompressible Navier Stokes model for studies of the ocean
+ on parallel computers,
+ J. Geophysical Res., 102(C3), 5753-5766.
+ Adcroft, A.J., Hill, C.N. and J. Marshall, (1997)
+ Representation of topography by shaved cells in a height coordinate ocean
+ model
+ Mon Wea Rev, vol 125, 2293-2315
+ Marshall, J., Jones, H. and C. Hill, (1998)
+ Efficient ocean modeling using non-hydrostatic algorithms
+ Journal of Marine Systems, 18, 115-134
+ Adcroft, A., Hill C. and J. Marshall: (1999)
+ A new treatment of the Coriolis terms in C-grid models at both high and low
+ resolutions,
+ Mon. Wea. Rev. Vol 127, pages 1928-1936
+ Hill, C, Adcroft,A., Jamous,D., and J. Marshall, (1999)
+ A Strategy for Terascale Climate Modeling.
+ In Proceedings of the Eighth ECMWF Workshop on the Use of Parallel Processors
+ in Meteorology, pages 406-425
+ World Scientific Publishing Co: UK
+ Marotzke, J, Giering,R., Zhang, K.Q., Stammer,D., Hill,C., and T.Lee, (1999)
+ Construction of the adjoint MIT ocean general circulation model and
+ application to Atlantic heat transport variability
+ J. Geophysical Res., 104(C12), 29,529-29,547.
+ \end{verbatim}
  We begin by briefly showing some of the results of the model in action to
  give a feel for the wide range of problems that can be addressed using it.
-Line 102 
 of them here. A more detailed descriptio
+Line 150 
 of them here. A more detailed descriptio
  numerical algorithm and implementation that lie behind these calculations is
  given later. Indeed many of the illustrative examples shown below can be
  easily reproduced: simply download the model (the minimum you need is a PC
- running linux, together with a FORTRAN\ 77 compiler) and follow the examples
+ running Linux, together with a FORTRAN\ 77 compiler) and follow the examples
  described in detail in the documentation.
  \subsection{Global atmosphere: `Held-Suarez' benchmark}
+ \begin{rawhtml}
+ <!-- CMIREDIR:atmospheric_example: -->
+ \end{rawhtml}
  A novel feature of MITgcm is its ability to simulate, using one basic algorithm,
  both atmospheric and oceanographic flows at both small and large scales.
-Line 126 
 there are no mountains or land-sea contr
+Line 179 
 there are no mountains or land-sea contr
  %% CNHend
  As described in Adcroft (2001), a `cubed sphere' is used to discretize the
- globe permitting a uniform gridding and obviated the need to Fourier filter.
+ globe permitting a uniform griding and obviated the need to Fourier filter.
  The `vector-invariant' form of MITgcm supports any orthogonal curvilinear
  grid, of which the cubed sphere is just one of many choices.
-Line 142 
 latitude-longitude grid. Both grids are
+Line 195 
 latitude-longitude grid. Both grids are
  %% CNHend
  \subsection{Ocean gyres}
+ \begin{rawhtml}
+ <!-- CMIREDIR:oceanic_example: -->
+ \end{rawhtml}
+ \begin{rawhtml}
+ <!-- CMIREDIR:ocean_gyres: -->
+ \end{rawhtml}
  Baroclinic instability is a ubiquitous process in the ocean, as well as the
  atmosphere. Ocean eddies play an important role in modifying the
-Line 163 
 warm water northward by the mean flow of
+Line 222 
 warm water northward by the mean flow of
  visible.
  %% CNHbegin
- \input{part1/ocean_gyres_figure}
+ \input{part1/atl6_figure}
  %% CNHend
  \subsection{Global ocean circulation}
+ \begin{rawhtml}
+ <!-- CMIREDIR:global_ocean_circulation: -->
+ \end{rawhtml}
  Figure \ref{fig:large-scale-circ} (top) shows the pattern of ocean currents at
  the surface of a 4$^{\circ }$
-Line 186 
 circulation of the global ocean in Sverd
+Line 248 
 circulation of the global ocean in Sverd
  %%CNHend
  \subsection{Convection and mixing over topography}
+ \begin{rawhtml}
+ <!-- CMIREDIR:mixing_over_topography: -->
+ \end{rawhtml}
  Dense plumes generated by localized cooling on the continental shelf of the
  ocean may be influenced by rotation when the deformation radius is smaller
  than the width of the cooling region. Rather than gravity plumes, the
  mechanism for moving dense fluid down the shelf is then through geostrophic
- eddies. The simulation shown in the figure \ref{fig::convect-and-topo}
+ eddies. The simulation shown in the figure \ref{fig:convect-and-topo}
  (blue is cold dense fluid, red is
  warmer, lighter fluid) employs the non-hydrostatic capability of MITgcm to
  trigger convection by surface cooling. The cold, dense water falls down the
-Line 205 
 instability of the along-slope current.
+Line 271 
 instability of the along-slope current.
  %%CNHend
  \subsection{Boundary forced internal waves}
+ \begin{rawhtml}
+ <!-- CMIREDIR:boundary_forced_internal_waves: -->
+ \end{rawhtml}
  The unique ability of MITgcm to treat non-hydrostatic dynamics in the
  presence of complex geometry makes it an ideal tool to study internal wave
-Line 224 
 nonhydrostatic dynamics.
+Line 293 
 nonhydrostatic dynamics.
  %%CNHend
  \subsection{Parameter sensitivity using the adjoint of MITgcm}
+ \begin{rawhtml}
+ <!-- CMIREDIR:parameter_sensitivity: -->
+ \end{rawhtml}
  Forward and tangent linear counterparts of MITgcm are supported using an
  `automatic adjoint compiler'. These can be used in parameter sensitivity and
-Line 231 
 data assimilation studies.
+Line 303 
 data assimilation studies.
  As one example of application of the MITgcm adjoint, Figure \ref{fig:hf-sensitivity}
  maps the gradient $\frac{\partial J}{\partial \mathcal{H}}$where $J$ is the magnitude
- of the overturning streamfunction shown in figure \ref{fig:large-scale-circ}
+ of the overturning stream-function shown in figure \ref{fig:large-scale-circ}
  at 60$^{\circ }$N and $
  \mathcal{H}(\lambda,\varphi)$ is the mean, local air-sea heat flux over
  a 100 year period. We see that $J$ is
-Line 244 
 yields sensitivities to all other model
+Line 316 
 yields sensitivities to all other model
  %%CNHend
  \subsection{Global state estimation of the ocean}
+ \begin{rawhtml}
+ <!-- CMIREDIR:global_state_estimation: -->
+ \end{rawhtml}
  An important application of MITgcm is in state estimation of the global
  ocean circulation. An appropriately defined `cost function', which measures
  the departure of the model from observations (both remotely sensed and
- insitu) over an interval of time, is minimized by adjusting `control
+ in-situ) over an interval of time, is minimized by adjusting `control
  parameters' such as air-sea fluxes, the wind field, the initial conditions
- etc. Figure \ref{fig:assimilated-globes} shows an estimate of the time-mean
+ etc. Figure \ref{fig:assimilated-globes} shows the large scale planetary
- surface elevation of the ocean obtained by bringing the model in to
+ circulation and a Hopf-Muller plot of Equatorial sea-surface height.
+ Both are obtained from assimilation bringing the model in to
  consistency with altimetric and in-situ observations over the period
--1997. {\bf CHANGE THIS TEXT - FIG FROM PATRICK/CARL/DETLEF}
+-1997.
  %% CNHbegin
- \input{part1/globes_figure}
+ \input{part1/assim_figure}
  %% CNHend
  \subsection{Ocean biogeochemical cycles}
+ \begin{rawhtml}
+ <!-- CMIREDIR:ocean_biogeo_cycles: -->
+ \end{rawhtml}
  MITgcm is being used to study global biogeochemical cycles in the ocean. For
  example one can study the effects of interannual changes in meteorological
-Line 275 
 telescoping to $\frac{1}{3}^{\circ}\time
+Line 355 
 telescoping to $\frac{1}{3}^{\circ}\time
  %%CNHend
  \subsection{Simulations of laboratory experiments}
+ \begin{rawhtml}
+ <!-- CMIREDIR:classroom_exp: -->
+ \end{rawhtml}
  Figure \ref{fig:lab-simulation} shows MITgcm being used to simulate a
- laboratory experiment enquiring in to the dynamics of the Antarctic Circumpolar Current (ACC). An
+ laboratory experiment inquiring into the dynamics of the Antarctic Circumpolar Current (ACC). An
  initially homogeneous tank of water ($1m$ in diameter) is driven from its
  free surface by a rotating heated disk. The combined action of mechanical
  and thermal forcing creates a lens of fluid which becomes baroclinically
-Line 293 
 stratification of the ACC.
+Line 376 
 stratification of the ACC.
  % $Name$
  \section{Continuous equations in `r' coordinates}
+ \begin{rawhtml}
+ <!-- CMIREDIR:z-p_isomorphism: -->
+ \end{rawhtml}
  To render atmosphere and ocean models from one dynamical core we exploit
  `isomorphisms' between equation sets that govern the evolution of the
-Line 301 
 One system of hydrodynamical equations i
+Line 387 
 One system of hydrodynamical equations i
  and encoded. The model variables have different interpretations depending on
  whether the atmosphere or ocean is being studied. Thus, for example, the
  vertical coordinate `$r$' is interpreted as pressure, $p$, if we are
- modeling the atmosphere (left hand side of figure \ref{fig:isomorphic-equations})
+ modeling the atmosphere (right hand side of figure \ref{fig:isomorphic-equations})
- and height, $z$, if we are modeling the ocean (right hand side of figure
+ and height, $z$, if we are modeling the ocean (left hand side of figure
  \ref{fig:isomorphic-equations}).
  %%CNHbegin
-Line 323 
 see figure \ref{fig:zandp-vert-coord}.
+Line 409 
 see figure \ref{fig:zandp-vert-coord}.
  \input{part1/vertcoord_figure.tex}
  %%CNHend
- \begin{equation*}
+ \begin{equation}
  \frac{D\vec{\mathbf{v}_{h}}}{Dt}+\left( 2\vec{\Omega}\times \vec{\mathbf{v}}
  \right) _{h}+\mathbf{\nabla }_{h}\phi =\mathcal{F}_{\vec{\mathbf{v}_{h}}}
- \text{ horizontal mtm}
+ \text{ horizontal mtm} \label{eq:horizontal_mtm}
- \end{equation*}
+ \end{equation}
- \begin{equation*}
+ \begin{equation}
  \frac{D\dot{r}}{Dt}+\widehat{k}\cdot \left( 2\vec{\Omega}\times \vec{\mathbf{
  v}}\right) +\frac{\partial \phi }{\partial r}+b=\mathcal{F}_{\dot{r}}\text{
- vertical mtm}
+ vertical mtm} \label{eq:vertical_mtm}
- \end{equation*}
+ \end{equation}
  \begin{equation}
  \mathbf{\nabla }_{h}\cdot \vec{\mathbf{v}}_{h}+\frac{\partial \dot{r}}{
- \partial r}=0\text{ continuity}  \label{eq:continuous}
+ \partial r}=0\text{ continuity}  \label{eq:continuity}
  \end{equation}
- \begin{equation*}
+ \begin{equation}
- b=b(\theta ,S,r)\text{ equation of state}
+ b=b(\theta ,S,r)\text{ equation of state} \label{eq:equation_of_state}
- \end{equation*}
+ \end{equation}
- \begin{equation*}
+ \begin{equation}
  \frac{D\theta }{Dt}=\mathcal{Q}_{\theta }\text{ potential temperature}
- \end{equation*}
+ \label{eq:potential_temperature}
+ \end{equation}
- \begin{equation*}
+ \begin{equation}
  \frac{DS}{Dt}=\mathcal{Q}_{S}\text{ humidity/salinity}
- \end{equation*}
+ \label{eq:humidity_salt}
+ \end{equation}
  Here:
-Line 424 
 in later chapters.
+Line 512 
 in later chapters.
  at fixed and moving $r$ surfaces we set (see figure \ref{fig:zandp-vert-coord}):
  \begin{equation}
- \dot{r}=0atr=R_{fixed}(x,y)\text{ (ocean bottom, top of the atmosphere)}
+ \dot{r}=0 \text{\ at\ } r=R_{fixed}(x,y)\text{ (ocean bottom, top of the atmosphere)}
  \label{eq:fixedbc}
  \end{equation}
  \begin{equation}
- \dot{r}=\frac{Dr}{Dt}atr=R_{moving}\text{ \
+ \dot{r}=\frac{Dr}{Dt} \text{\ at\ } r=R_{moving}\text{ \
- (oceansurface,bottomoftheatmosphere)}  \label{eq:movingbc}
+ (ocean surface,bottom of the atmosphere)}  \label{eq:movingbc}
  \end{equation}
  Here
-Line 523 
 The boundary conditions at top and botto
+Line 611 
 The boundary conditions at top and botto
  atmosphere)}  \label{eq:moving-bc-atmos}
  \end{eqnarray}
- Then the (hydrostatic form of) eq(\ref{eq:continuous}) yields a consistent
+ Then the (hydrostatic form of) equations
- set of atmospheric equations which, for convenience, are written out in $p$
+ (\ref{eq:horizontal_mtm}-\ref{eq:humidity_salt}) yields a consistent
- coordinates in Appendix Atmosphere - see eqs(\ref{eq:atmos-prime}).
+ set of atmospheric equations which, for convenience, are written out
+ in $p$ coordinates in Appendix Atmosphere - see
+ eqs(\ref{eq:atmos-prime}).
  \subsection{Ocean}
-Line 560 
 w &=&\frac{D\eta }{Dt}\text{ at }r=R_{mo
+Line 650 
 w &=&\frac{D\eta }{Dt}\text{ at }r=R_{mo
  \end{eqnarray}
  where $\eta $ is the elevation of the free surface.
- Then eq(\ref{eq:continuous}) yields a consistent set of oceanic equations
+ Then equations (\ref{eq:horizontal_mtm}-\ref{eq:humidity_salt}) yield a consistent set
+ of oceanic equations
  which, for convenience, are written out in $z$ coordinates in Appendix Ocean
  - see eqs(\ref{eq:ocean-mom}) to (\ref{eq:ocean-salt}).
  \subsection{Hydrostatic, Quasi-hydrostatic, Quasi-nonhydrostatic and
  Non-hydrostatic forms}
+ \begin{rawhtml}
+ <!-- CMIREDIR:non_hydrostatic: -->
+ \end{rawhtml}
  Let us separate $\phi $ in to surface, hydrostatic and non-hydrostatic terms:
-Line 573 
 Let us separate $\phi $ in to surface, h
+Line 668 
 Let us separate $\phi $ in to surface, h
  \phi (x,y,r)=\phi _{s}(x,y)+\phi _{hyd}(x,y,r)+\phi _{nh}(x,y,r)
  \label{eq:phi-split}
  \end{equation}
- and write eq(\ref{incompressible}a,b) in the form:
+ %and write eq(\ref{eq:incompressible}) in the form:
+ %                  ^- this eq is missing (jmc) ; replaced with:
+ and write eq( \ref{eq:horizontal_mtm}) in the form:
  \begin{equation}
  \frac{\partial \vec{\mathbf{v}_{h}}}{\partial t}+\mathbf{\nabla }_{h}\phi
-Line 666 
 In the above `${r}$' is the distance fro
+Line 763 
 In the above `${r}$' is the distance fro
  Grad and div operators in spherical coordinates are defined in appendix
  OPERATORS.
- \marginpar{
- Fig.6 Spherical polar coordinate system.}
  %%CNHbegin
  \input{part1/sphere_coord_figure.tex}
-Line 730 
 In the non-hydrostatic ocean model all t
+Line 825 
 In the non-hydrostatic ocean model all t
  three dimensional elliptic equation must be solved subject to Neumann
  boundary conditions (see below). It is important to note that use of the
  full \textbf{NH} does not admit any new `fast' waves in to the system - the
- incompressible condition eq(\ref{eq:continuous})c has already filtered out
+ incompressible condition eq(\ref{eq:continuity}) has already filtered out
  acoustic modes. It does, however, ensure that the gravity waves are treated
  accurately with an exact dispersion relation. The \textbf{NH} set has a
  complete angular momentum principle and consistent energetics - see White
-Line 779 
 coordinates are supported - see eqs(\ref
+Line 874 
 coordinates are supported - see eqs(\ref
  \subsection{Solution strategy}
  The method of solution employed in the \textbf{HPE}, \textbf{QH} and \textbf{
- NH} models is summarized in Fig.7.
+ NH} models is summarized in Figure \ref{fig:solution-strategy}.
- \marginpar{
+ Under all dynamics, a 2-d elliptic equation is
- Fig.7 Solution strategy} Under all dynamics, a 2-d elliptic equation is
  first solved to find the surface pressure and the hydrostatic pressure at
  any level computed from the weight of fluid above. Under \textbf{HPE} and
  \textbf{QH} dynamics, the horizontal momentum equations are then stepped
-Line 838 
 atmospheric pressure pushing down on the
+Line 932 
 atmospheric pressure pushing down on the
  \subsubsection{Surface pressure}
- The surface pressure equation can be obtained by integrating continuity, (
+ The surface pressure equation can be obtained by integrating continuity,
- \ref{eq:continuous})c, vertically from $r=R_{fixed}$ to $r=R_{moving}$
+ (\ref{eq:continuity}), vertically from $r=R_{fixed}$ to $r=R_{moving}$
  \begin{equation*}
  \int_{R_{fixed}}^{R_{moving}}\left( \mathbf{\nabla }_{h}\cdot \vec{\mathbf{v}
-Line 864 
 r $. The above can be rearranged to yiel
+Line 958 
 r $. The above can be rearranged to yiel
  where we have incorporated a source term.
  Whether $\phi $ is pressure (ocean model, $p/\rho _{c}$) or geopotential
- (atmospheric model), in (\ref{mtm-split}), the horizontal gradient term can
+ (atmospheric model), in (\ref{eq:mom-h}), the horizontal gradient term can
  be written
  \begin{equation}
  \mathbf{\nabla }_{h}\phi _{s}=\mathbf{\nabla }_{h}\left( b_{s}\eta \right)
-Line 872 
 be written
+Line 966 
 be written
  \end{equation}
  where $b_{s}$ is the buoyancy at the surface.
- In the hydrostatic limit ($\epsilon _{nh}=0$), Eqs(\ref{eq:mom-h}), (\ref
+ In the hydrostatic limit ($\epsilon _{nh}=0$), equations (\ref{eq:mom-h}), (\ref
  {eq:free-surface}) and (\ref{eq:phi-surf}) can be solved by inverting a 2-d
  elliptic equation for $\phi _{s}$ as described in Chapter 2. Both `free
  surface' and `rigid lid' approaches are available.
  \subsubsection{Non-hydrostatic pressure}
- Taking the horizontal divergence of (\ref{hor-mtm}) and adding $\frac{
+ Taking the horizontal divergence of (\ref{eq:mom-h}) and adding
- \partial }{\partial r}$ of (\ref{vertmtm}), invoking the continuity equation
+ $\frac{\partial }{\partial r}$ of (\ref{eq:mom-w}), invoking the continuity equation
- (\ref{incompressible}), we deduce that:
+ (\ref{eq:continuity}), we deduce that:
  \begin{equation}
  \nabla _{3}^{2}\phi _{nh}=\nabla .\vec{\mathbf{G}}_{\vec{v}}-\left( \mathbf{
-Line 911 
 tangential component of velocity, $v_{T}
+Line 1005 
 tangential component of velocity, $v_{T}
  depending on the form chosen for the dissipative terms in the momentum
  equations - see below.
- Eq.(\ref{nonormalflow}) implies, making use of (\ref{mtm-split}), that:
+ Eq.(\ref{nonormalflow}) implies, making use of (\ref{eq:mom-h}), that:
  \begin{equation}
  \widehat{n}.\nabla \phi _{nh}=\widehat{n}.\vec{\mathbf{F}}
-Line 951 
 If the flow is `close' to hydrostatic ba
+Line 1045 
 If the flow is `close' to hydrostatic ba
  converges rapidly because $\phi _{nh}\ $is then only a small correction to
  the hydrostatic pressure field (see the discussion in Marshall et al, a,b).
- The solution $\phi _{nh}\ $to (\ref{eq:3d-invert}) and (\ref{homneuman})
+ The solution $\phi _{nh}\ $to (\ref{eq:3d-invert}) and (\ref{eq:inhom-neumann-nh})
  does not vanish at $r=R_{moving}$, and so refines the pressure there.
  \subsection{Forcing/dissipation}
-Line 959 
 does not vanish at $r=R_{moving}$, and s
+Line 1053 
 does not vanish at $r=R_{moving}$, and s
  \subsubsection{Forcing}
  The forcing terms $\mathcal{F}$ on the rhs of the equations are provided by
- `physics packages' described in detail in chapter ??.
+ `physics packages' and forcing packages. These are described later on.
  \subsubsection{Dissipation}
-Line 1006 
 salinity ... ).
+Line 1100 
 salinity ... ).
  \subsection{Vector invariant form}
- For some purposes it is advantageous to write momentum advection in eq(\ref
+ For some purposes it is advantageous to write momentum advection in
- {hor-mtm}) and (\ref{vertmtm}) in the (so-called) `vector invariant' form:
+ eq(\ref {eq:horizontal_mtm}) and (\ref{eq:vertical_mtm}) in the
+ (so-called) `vector invariant' form:
  \begin{equation}
  \frac{D\vec{\mathbf{v}}}{Dt}=\frac{\partial \vec{\mathbf{v}}}{\partial t}
-Line 1025 
 to discretize the model.
+Line 1120 
 to discretize the model.
  \subsection{Adjoint}
- Tangent linear and adjoint counterparts of the forward model and described
+ Tangent linear and adjoint counterparts of the forward model are described
  in Chapter 5.
  % $Header$
-Line 1118 
 In $p$-coordinates, the upper boundary a
+Line 1213 
 In $p$-coordinates, the upper boundary a
  surface ($\phi $ is imposed and $\omega \neq 0$).
  \subsubsection{Splitting the geo-potential}
+ \label{sec:hpe-p-geo-potential-split}
  For the purposes of initialization and reducing round-off errors, the model
  deals with perturbations from reference (or ``standard'') profiles. For
-Line 1147 
 _{o}(p_{o})=g~Z_{topo}$, defined:
+Line 1243 
 _{o}(p_{o})=g~Z_{topo}$, defined:
  The final form of the HPE's in p coordinates is then:
  \begin{eqnarray}
  \frac{D\vec{\mathbf{v}}_{h}}{Dt}+f\hat{\mathbf{k}}\times \vec{\mathbf{v}}
- _{h}+\mathbf{\nabla }_{p}\phi ^{\prime } &=&\vec{\mathbf{\mathcal{F}}} \\
+ _{h}+\mathbf{\nabla }_{p}\phi ^{\prime } &=&\vec{\mathbf{\mathcal{F}}}
+ \label{eq:atmos-prime} \\
  \frac{\partial \phi ^{\prime }}{\partial p}+\alpha ^{\prime } &=&0 \\
  \mathbf{\nabla }_{p}\cdot \vec{\mathbf{v}}_{h}+\frac{\partial \omega }{
  \partial p} &=&0 \\
  \frac{\partial \Pi }{\partial p}\theta ^{\prime } &=&\alpha ^{\prime } \\
- \frac{D\theta }{Dt} &=&\frac{\mathcal{Q}}{\Pi }  \label{eq:atmos-prime}
+ \frac{D\theta }{Dt} &=&\frac{\mathcal{Q}}{\Pi }
  \end{eqnarray}
  % $Header$
-Line 1171 
 _{h}+\frac{1}{\rho }\mathbf{\nabla }_{z}
+Line 1268 
 _{h}+\frac{1}{\rho }\mathbf{\nabla }_{z}
  \epsilon _{nh}\frac{Dw}{Dt}+g+\frac{1}{\rho }\frac{\partial p}{\partial z}
  &=&\epsilon _{nh}\mathcal{F}_{w} \\
  \frac{1}{\rho }\frac{D\rho }{Dt}+\mathbf{\nabla }_{z}\cdot \vec{\mathbf{v}}
- _{h}+\frac{\partial w}{\partial z} &=&0 \\
+ _{h}+\frac{\partial w}{\partial z} &=&0 \label{eq-zns-cont}\\
- \rho &=&\rho (\theta ,S,p) \\
+ \rho &=&\rho (\theta ,S,p) \label{eq-zns-eos}\\
- \frac{D\theta }{Dt} &=&\mathcal{Q}_{\theta } \\
+ \frac{D\theta }{Dt} &=&\mathcal{Q}_{\theta } \label{eq-zns-heat}\\
- \frac{DS}{Dt} &=&\mathcal{Q}_{s}  \label{eq:non-boussinesq}
+ \frac{DS}{Dt} &=&\mathcal{Q}_{s}  \label{eq-zns-salt}
+ \label{eq:non-boussinesq}
  \end{eqnarray}
  These equations permit acoustics modes, inertia-gravity waves,
- non-hydrostatic motions, a geostrophic (Rossby) mode and a thermo-haline
+ non-hydrostatic motions, a geostrophic (Rossby) mode and a thermohaline
  mode. As written, they cannot be integrated forward consistently - if we
  step $\rho $ forward in (\ref{eq-zns-cont}), the answer will not be
  consistent with that obtained by stepping (\ref{eq-zns-heat}) and (\ref
-Line 1192 
 _{\theta ,p}\frac{DS}{Dt}+\left. \frac{\
+Line 1290 
 _{\theta ,p}\frac{DS}{Dt}+\left. \frac{\
  _{\theta ,S}\frac{Dp}{Dt}  \label{EOSexpansion}
  \end{equation}
- Note that $\frac{\partial \rho }{\partial p}=\frac{1}{c_{s}^{2}}$ is the
+ Note that $\frac{\partial \rho }{\partial p}=\frac{1}{c_{s}^{2}}$ is
- reciprocal of the sound speed ($c_{s}$) squared. Substituting into \ref
+ the reciprocal of the sound speed ($c_{s}$) squared. Substituting into
- {eq-zns-cont} gives:
+ \ref{eq-zns-cont} gives:
  \begin{equation}
  \frac{1}{\rho c_{s}^{2}}\frac{Dp}{Dt}+\mathbf{\nabla }_{z}\cdot \vec{\mathbf{
  v}}+\partial _{z}w\approx 0  \label{eq-zns-pressure}

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changed lines


 
Added in v.1.23
-Removed from v.1.7
+Added in v.1.23

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