| 5 |
viscous-plastic (VP) dynamic-thermodynamic sea ice model |
viscous-plastic (VP) dynamic-thermodynamic sea ice model |
| 6 |
\citep{zhang97} first introduced by \citet{hibler79, hibler80}. In |
\citep{zhang97} first introduced by \citet{hibler79, hibler80}. In |
| 7 |
order to adapt this model to the requirements of coupled |
order to adapt this model to the requirements of coupled |
| 8 |
ice-ocean simulations, many important aspects of the original code have |
ice-ocean state estimation, many important aspects of the original code have |
| 9 |
been modified and improved: |
been modified and improved: |
| 10 |
\begin{itemize} |
\begin{itemize} |
| 11 |
\item the code has been rewritten for an Arakawa C-grid, both B- and |
\item the code has been rewritten for an Arakawa C-grid, both B- and |
| 12 |
C-grid variants are available; the C-grid code allows for no-slip |
C-grid variants are available; the C-grid code allows for no-slip |
| 13 |
and free-slip lateral boundary conditions; |
and free-slip lateral boundary conditions; |
| 14 |
\item two different solution methods for solving the nonlinear |
\item two different solution methods for solving the nonlinear |
| 15 |
momentum equations have been adopted: LSOR \citep{zhang97}, EVP |
momentum equations have been adopted: LSOR \citep{zhang97}, and EVP |
| 16 |
\citep{hunke97}; |
\citep{hunke97}; |
| 17 |
\item ice-ocean stress can be formulated as in \citet{hibler87}; |
\item ice-ocean stress can be formulated as in \citet{hibler87} or as in \citet{cam08}; |
| 18 |
\item ice variables are advected by sophisticated advection schemes; |
\item ice variables \ml{can be} advected by sophisticated, \ml{conservative} |
| 19 |
\item growth and melt parameterizaion have been refined and extended |
advection schemes \ml{with flux limiting}; |
| 20 |
in order to allow for automatic differentiation of the code. |
\item growth and melt parameterizations have been refined and extended |
| 21 |
|
in order to allow for more stable automatic differentiation of the code. |
| 22 |
\end{itemize} |
\end{itemize} |
|
The model equations and their numerical realization are summarized |
|
|
below. |
|
| 23 |
|
|
| 24 |
\subsection{Dynamics} |
The sea ice model is tightly coupled to the ocean compontent of the MITgcm |
| 25 |
\label{sec:dynamics} |
\citep{mar97a}. Heat, fresh water fluxes and surface stresses are computed |
| 26 |
|
from the atmospheric state and modified by the ice model at every time step. |
| 27 |
|
The model equations and details of their numerical realization are summarized |
| 28 |
|
in the appendix. Further documentation and model code can be found at |
| 29 |
|
\url{http://mitgcm.org}. |
| 30 |
|
|
| 31 |
The momentum equation of the sea-ice model is |
%\subsection{C-grid} |
|
\begin{equation} |
|
|
\label{eq:momseaice} |
|
|
m \frac{D\vek{u}}{Dt} = -mf\vek{k}\times\vek{u} + \vtau_{air} + |
|
|
\vtau_{ocean} - m \nabla{\phi(0)} + \vek{F}, |
|
|
\end{equation} |
|
|
where $m=m_{i}+m_{s}$ is the ice and snow mass per unit area; |
|
|
$\vek{u}=u\vek{i}+v\vek{j}$ is the ice velocity vector; |
|
|
$\vek{i}$, $\vek{j}$, and $\vek{k}$ are unit vectors in the $x$, $y$, and $z$ |
|
|
directions, respectively; |
|
|
$f$ is the Coriolis parameter; |
|
|
$\vtau_{air}$ and $\vtau_{ocean}$ are the wind-ice and ocean-ice stresses, |
|
|
respectively; |
|
|
$g$ is the gravity accelation; |
|
|
$\nabla\phi(0)$ is the gradient (or tilt) of the sea surface height; |
|
|
$\phi(0) = g\eta + p_{a}/\rho_{0}$ is the sea surface height potential |
|
|
in response to ocean dynamics ($g\eta$) and to atmospheric pressure |
|
|
loading ($p_{a}/\rho_{0}$, where $\rho_{0}$ is a reference density); |
|
|
and $\vek{F}=\nabla\cdot\sigma$ is the divergence of the internal ice stress |
|
|
tensor $\sigma_{ij}$. |
|
|
When using the rescaled vertical coordinate system, z$^\ast$, of |
|
|
\citet{cam08}, $\phi(0)$ also includes a term due to snow and ice |
|
|
loading, $mg/\rho_{0}$. |
|
|
Advection of sea-ice momentum is neglected. The wind and ice-ocean stress |
|
|
terms are given by |
|
|
\begin{align*} |
|
|
\vtau_{air} = & \rho_{air} C_{air} |\vek{U}_{air} -\vek{u}| |
|
|
R_{air} (\vek{U}_{air} -\vek{u}), \\ |
|
|
\vtau_{ocean} = & \rho_{ocean}C_{ocean} |\vek{U}_{ocean}-\vek{u}| |
|
|
R_{ocean}(\vek{U}_{ocean}-\vek{u}), \\ |
|
|
\end{align*} |
|
|
where $\vek{U}_{air/ocean}$ are the surface winds of the atmosphere |
|
|
and surface currents of the ocean, respectively; $C_{air/ocean}$ are |
|
|
air and ocean drag coefficients; $\rho_{air/ocean}$ are reference |
|
|
densities; and $R_{air/ocean}$ are rotation matrices that act on the |
|
|
wind/current vectors. |
|
|
|
|
|
For an isotropic system the stress tensor $\sigma_{ij}$ ($i,j=1,2$) can |
|
|
be related to the ice strain rate and strength by a nonlinear |
|
|
viscous-plastic (VP) constitutive law \citep{hibler79, zhang97}: |
|
|
\begin{equation} |
|
|
\label{eq:vpequation} |
|
|
\sigma_{ij}=2\eta(\dot{\epsilon}_{ij},P)\dot{\epsilon}_{ij} |
|
|
+ \left[\zeta(\dot{\epsilon}_{ij},P) - |
|
|
\eta(\dot{\epsilon}_{ij},P)\right]\dot{\epsilon}_{kk}\delta_{ij} |
|
|
- \frac{P}{2}\delta_{ij}. |
|
|
\end{equation} |
|
|
The ice strain rate is given by |
|
|
\begin{equation*} |
|
|
\dot{\epsilon}_{ij} = \frac{1}{2}\left( |
|
|
\frac{\partial{u_{i}}}{\partial{x_{j}}} + |
|
|
\frac{\partial{u_{j}}}{\partial{x_{i}}}\right). |
|
|
\end{equation*} |
|
|
The maximum ice pressure $P_{\max}$, a measure of ice strength, depends on |
|
|
both thickness $h$ and compactness (concentration) $c$: |
|
|
\begin{equation} |
|
|
P_{\max} = P^{*}c\,h\,e^{[C^{*}\cdot(1-c)]}, |
|
|
\label{eq:icestrength} |
|
|
\end{equation} |
|
|
with the constants $P^{*}$ and $C^{*}$. The nonlinear bulk and shear |
|
|
viscosities $\eta$ and $\zeta$ are functions of ice strain rate |
|
|
invariants and ice strength such that the principal components of the |
|
|
stress lie on an elliptical yield curve with the ratio of major to |
|
|
minor axis $e$ equal to $2$; they are given by: |
|
|
\begin{align*} |
|
|
\zeta =& \min\left(\frac{P_{\max}}{2\max(\Delta,\Delta_{\min})}, |
|
|
\zeta_{\max}\right) \\ |
|
|
\eta =& \frac{\zeta}{e^2} \\ |
|
|
\intertext{with the abbreviation} |
|
|
\Delta = & \left[ |
|
|
\left(\dot{\epsilon}_{11}^2+\dot{\epsilon}_{22}^2\right) |
|
|
(1+e^{-2}) + 4e^{-2}\dot{\epsilon}_{12}^2 + |
|
|
2\dot{\epsilon}_{11}\dot{\epsilon}_{22} (1-e^{-2}) |
|
|
\right]^{\frac{1}{2}} |
|
|
\end{align*} |
|
|
The bulk viscosities are bounded above by imposing both a minimum |
|
|
$\Delta_{\min}=10^{-11}\text{\,s}^{-1}$ (for numerical reasons) and a |
|
|
maximum $\zeta_{\max} = P_{\max}/\Delta^*$, where |
|
|
$\Delta^*=(5\times10^{12}/2\times10^4)\text{\,s}^{-1}$. For stress |
|
|
tensor computation the replacement pressure $P = 2\,\Delta\zeta$ |
|
|
\citep{hibler95} is used so that the stress state always lies on the |
|
|
elliptic yield curve by definition. |
|
|
|
|
|
In the so-called truncated ellipse method the shear viscosity $\eta$ |
|
|
is capped to suppress any tensile stress \citep{hibler97, geiger98}: |
|
|
\begin{equation} |
|
|
\label{eq:etatem} |
|
|
\eta = \min\left(\frac{\zeta}{e^2}, |
|
|
\frac{\frac{P}{2}-\zeta(\dot{\epsilon}_{11}+\dot{\epsilon}_{22})} |
|
|
{\sqrt{(\dot{\epsilon}_{11}+\dot{\epsilon}_{22})^2 |
|
|
+4\dot{\epsilon}_{12}^2}}\right). |
|
|
\end{equation} |
|
|
|
|
|
In the current implementation, the VP-model is integrated with the |
|
|
semi-implicit line successive over relaxation (LSOR)-solver of |
|
|
\citet{zhang98}, which allows for long time steps that, in our case, |
|
|
are limited by the explicit treatment of the Coriolis term. The |
|
|
explicit treatment of the Coriolis term does not represent a severe |
|
|
limitation because it restricts the time step to approximately the |
|
|
same length as in the ocean model where the Coriolis term is also |
|
|
treated explicitly. |
|
|
|
|
|
\citet{hunke97}'s introduced an elastic contribution to the strain |
|
|
rate in order to regularize Eq.\refeq{vpequation} in such a way that |
|
|
the resulting elastic-viscous-plastic (EVP) and VP models are |
|
|
identical at steady state, |
|
|
\begin{equation} |
|
|
\label{eq:evpequation} |
|
|
\frac{1}{E}\frac{\partial\sigma_{ij}}{\partial{t}} + |
|
|
\frac{1}{2\eta}\sigma_{ij} |
|
|
+ \frac{\eta - \zeta}{4\zeta\eta}\sigma_{kk}\delta_{ij} |
|
|
+ \frac{P}{4\zeta}\delta_{ij} |
|
|
= \dot{\epsilon}_{ij}. |
|
|
\end{equation} |
|
|
%In the EVP model, equations for the components of the stress tensor |
|
|
%$\sigma_{ij}$ are solved explicitly. Both model formulations will be |
|
|
%used and compared the present sea-ice model study. |
|
|
The EVP-model uses an explicit time stepping scheme with a short |
|
|
timestep. According to the recommendation of \citet{hunke97}, the |
|
|
EVP-model is stepped forward in time 120 times within the physical |
|
|
ocean model time step (although this parameter is under debate), to |
|
|
allow for elastic waves to disappear. Because the scheme does not |
|
|
require a matrix inversion it is fast in spite of the small timestep |
|
|
\citep{hunke97}. For completeness, we repeat the equations for the |
|
|
components of the stress tensor $\sigma_{1} = |
|
|
\sigma_{11}+\sigma_{22}$, $\sigma_{2}= \sigma_{11}-\sigma_{22}$, and |
|
|
$\sigma_{12}$. Introducing the divergence $D_D = |
|
|
\dot{\epsilon}_{11}+\dot{\epsilon}_{22}$, and the horizontal tension |
|
|
and shearing strain rates, $D_T = |
|
|
\dot{\epsilon}_{11}-\dot{\epsilon}_{22}$ and $D_S = |
|
|
2\dot{\epsilon}_{12}$, respectively, and using the above abbreviations, |
|
|
the equations can be written as: |
|
|
\begin{align} |
|
|
\label{eq:evpstresstensor1} |
|
|
\frac{\partial\sigma_{1}}{\partial{t}} + \frac{\sigma_{1}}{2T} + |
|
|
\frac{P}{2T} &= \frac{P}{2T\Delta} D_D \\ |
|
|
\label{eq:evpstresstensor2} |
|
|
\frac{\partial\sigma_{2}}{\partial{t}} + \frac{\sigma_{2} e^{2}}{2T} |
|
|
&= \frac{P}{2T\Delta} D_T \\ |
|
|
\label{eq:evpstresstensor12} |
|
|
\frac{\partial\sigma_{12}}{\partial{t}} + \frac{\sigma_{12} e^{2}}{2T} |
|
|
&= \frac{P}{4T\Delta} D_S |
|
|
\end{align} |
|
|
Here, the elastic parameter $E$ is redefined in terms of a damping timescale |
|
|
$T$ for elastic waves \[E=\frac{\zeta}{T}.\] |
|
|
$T=E_{0}\Delta{t}$ with the tunable parameter $E_0<1$ and |
|
|
the external (long) timestep $\Delta{t}$. \citet{hunke97} recommend |
|
|
$E_{0} = \frac{1}{3}$. |
|
|
|
|
|
For details of the spatial discretization, the reader is referred to |
|
|
\citet{zhang98, zhang03}. Our discretization differs only (but |
|
|
importantly) in the underlying grid, namely the Arakawa C-grid, but is |
|
|
otherwise straightforward. The EVP model, in particular, is discretized |
|
|
naturally on the C-grid with $\sigma_{1}$ and $\sigma_{2}$ on the |
|
|
center points and $\sigma_{12}$ on the corner (or vorticity) points of |
|
|
the grid. With this choice all derivatives are discretized as central |
|
|
differences and averaging is only involved in computing $\Delta$ and |
|
|
$P$ at vorticity points. |
|
|
|
|
|
For a general curvilinear grid, one needs in principle to take metric |
|
|
terms into account that arise in the transformation of a curvilinear |
|
|
grid on the sphere. For now, however, we can neglect these metric |
|
|
terms because they are very small on the \ml{[modify following |
|
|
section3:] cubed sphere grids used in this paper; in particular, |
|
|
only near the edges of the cubed sphere grid, we expect them to be |
|
|
non-zero, but these edges are at approximately 35\degS\ or 35\degN\ |
|
|
which are never covered by sea-ice in our simulations. Everywhere |
|
|
else the coordinate system is locally nearly cartesian.} However, for |
|
|
last-glacial-maximum or snowball-earth-like simulations the question |
|
|
of metric terms needs to be reconsidered. Either, one includes these |
|
|
terms as in \citet{zhang03}, or one finds a vector-invariant |
|
|
formulation for the sea-ice internal stress term that does not require |
|
|
any metric terms, as it is done in the ocean dynamics of the MITgcm |
|
|
\citep{adcroft04:_cubed_sphere}. |
|
|
|
|
|
Lateral boundary conditions are naturally ``no-slip'' for B-grids, as |
|
|
the tangential velocities points lie on the boundary. For C-grids, the |
|
|
lateral boundary condition for tangential velocities is realized via |
|
|
``ghost points'', allowing alternatively no-slip or free-slip |
|
|
conditions. In ocean models free-slip boundary conditions in |
|
|
conjunction with piecewise-constant (``castellated'') coastlines have |
|
|
been shown to reduce in effect to no-slip boundary conditions |
|
|
\citep{adcroft98:_slippery_coast}; for sea-ice models the effects of |
|
|
lateral boundary conditions have not yet been studied. |
|
|
|
|
|
Moving sea ice exerts a stress on the ocean which is the opposite of |
|
|
the stress $\vtau_{ocean}$ in Eq.\refeq{momseaice}. This stess is |
|
|
applied directly to the surface layer of the ocean model. An |
|
|
alternative ocean stress formulation is given by \citet{hibler87}. |
|
|
Rather than applying $\vtau_{ocean}$ directly, the stress is derived |
|
|
from integrating over the ice thickness to the bottom of the oceanic |
|
|
surface layer. In the resulting equation for the \emph{combined} |
|
|
ocean-ice momentum, the interfacial stress cancels and the total |
|
|
stress appears as the sum of windstress and divergence of internal ice |
|
|
stresses: $\delta(z) (\vtau_{air} + \vek{F})/\rho_0$, \citep[see also |
|
|
Eq.\,2 of][]{hibler87}. The disadvantage of this formulation is that |
|
|
now the velocity in the surface layer of the ocean that is used to |
|
|
advect tracers, is really an average over the ocean surface |
|
|
velocity and the ice velocity leading to an inconsistency as the ice |
|
|
temperature and salinity are different from the oceanic variables. |
|
|
|
|
|
%\subparagraph{boundary conditions: no-slip, free-slip, half-slip} |
|
| 32 |
%\begin{itemize} |
%\begin{itemize} |
| 33 |
%\item transition from existing B-Grid to C-Grid |
%\item no-slip vs. free-slip for both lsr and evp; |
| 34 |
%\item boundary conditions: no-slip, free-slip, half-slip |
% "diagnostics" to look at and use for comparison |
| 35 |
%\item fancy (multi dimensional) advection schemes |
% \begin{itemize} |
| 36 |
%\item VP vs.\ EVP \citep{hunke97} |
% \item ice transport through Fram Strait/Denmark Strait/Davis |
| 37 |
%\item ocean stress formulation \citep{hibler87} |
% Strait/Bering strait (these are general) |
| 38 |
%\end{itemize} |
% \item ice transport through narrow passages, e.g.\ Nares-Strait |
| 39 |
|
% \end{itemize} |
| 40 |
\subsection{Thermodynamics} |
%\item compare different advection schemes (if lsr turns out to be more |
| 41 |
\label{sec:thermodynamics} |
% effective, then with lsr otherwise I prefer evp), eg. |
| 42 |
|
% \begin{itemize} |
| 43 |
In the original formulation the sea ice model \citep{menemenlis05} |
% \item default 2nd-order with diff1=0.002 |
| 44 |
uses simple thermodynamics following the appendix of |
% \item 1st-order upwind with diff1=0. |
| 45 |
\citet{semtner76}. This formulation does not allow storage of heat |
% \item DST3FL (SEAICEadvScheme=33 with diff1=0., should work, works for me) |
| 46 |
(heat capacity of ice is zero, and this type of model is often refered |
% \item 2nd-order wit flux limiter (SEAICEadvScheme=77 with diff1=0.) |
| 47 |
to as a ``zero-layer'' model). Upward conductive heat flux is parameterized |
% \end{itemize} |
| 48 |
assuming a linear temperature profile and together with a constant ice |
% That should be enough. Here, total ice mass and location of ice edge |
| 49 |
conductivity. It is expressed as $(K/h)(T_{w}-T_{0})$, where $K$ is |
% is interesting. However, this comparison can be done in an idealized |
| 50 |
the ice conductivity, $h$ the ice thickness, and $T_{w}-T_{0}$ the |
% domain, may not require full Arctic Domain? |
| 51 |
difference between water and ice surface temperatures. The surface |
%\item |
| 52 |
heat flux is computed in a similar way to that of \citet{parkinson79} |
%Do a little study on the parameters of LSR and EVP |
| 53 |
and \citet{manabe79}. |
%\begin{enumerate} |
| 54 |
|
%\item convergence of LSR, how many iterations do you need to get a |
| 55 |
|
% true elliptic yield curve |
| 56 |
|
%\item vary deltaTevp and the relaxation parameter for EVP and see when |
| 57 |
|
% the EVP solution breaks down (relative to the forcing time scale). |
| 58 |
|
% For this, it is essential that the evp solver gives use "stripeless" |
| 59 |
|
% solutions, that is your dtevp = 1sec solutions/or 10sec solutions |
| 60 |
|
% with SEAICE\_evpDampC = 615. |
| 61 |
|
%\end{enumerate} |
| 62 |
|
|
| 63 |
The conductive heat flux depends strongly on the ice thickness $h$. |
%\end{itemize} |
|
However, the ice thickness in the model represents a mean over a |
|
|
potentially very heterogeneous thickness distribution. In order to |
|
|
parameterize this sub-grid scale distribution for heat flux |
|
|
computations, the mean ice thickness $h$ is split into seven thickness |
|
|
categories $H_{n}$ that are equally distributed between $2h$ and |
|
|
minimum imposed ice thickness of $5\text{\,cm}$ by $H_n= |
|
|
\frac{2n-1}{7}\,h$ for $n\in[1,7]$. The heat fluxes computed for each |
|
|
thickness category area averaged to give the total heat flux. \ml{[I |
|
|
don't have citation for that, anyone?]} |
|
|
|
|
|
The atmospheric heat flux is balanced by an oceanic heat flux from |
|
|
below. The oceanic flux is proportional to |
|
|
$\rho\,c_{p}\left(T_{w}-T_{fr}\right)$ where $\rho$ and $c_{p}$ are |
|
|
the density and heat capacity of sea water and $T_{fr}$ is the local |
|
|
freezing point temperature that is a function of salinity. Contrary to |
|
|
\citet{menemenlis05}, this flux is not assumed to instantaneously melt |
|
|
or create ice, but a time scale of three days is used to relax $T_{w}$ |
|
|
to the freezing point. |
|
|
|
|
|
The parameterization of lateral and vertical growth of sea ice follows |
|
|
that of \citet{hibler79, hibler80}. |
|
|
|
|
|
On top of the ice there is a layer of snow that modifies the heat flux |
|
|
and the albedo \citep{zhang98}. If enough snow accumulates so that its |
|
|
weight submerges the ice and the snow is flooded, a simple mass |
|
|
conserving parameterization of snowice formation (a flood-freeze |
|
|
algorithm following Archimedes' principle) turns snow into ice until |
|
|
the ice surface is back at $z=0$ \citep{leppaeranta83}. |
|
|
|
|
|
Effective ice thickness (ice volume per unit area, |
|
|
$c\cdot{h}$), concentration $c$ and effective snow thickness |
|
|
($c\cdot{h}_{s}$) are advected by ice velocities: |
|
|
\begin{equation} |
|
|
\label{eq:advection} |
|
|
\frac{\partial{X}}{\partial{t}} = - \nabla\cdot\left(\vek{u}\,X\right) + |
|
|
\Gamma_{X} + D_{X} |
|
|
\end{equation} |
|
|
where $\Gamma_X$ are the thermodynamic source terms and $D_{X}$ the |
|
|
diffusive terms for quantities $X=(c\cdot{h}), c, (c\cdot{h}_{s})$. |
|
|
% |
|
|
From the various advection scheme that are available in the MITgcm |
|
|
\citep{mitgcm02}, we choose flux-limited schemes |
|
|
\citep[multidimensional 2nd and 3rd-order advection scheme with flux |
|
|
limiter][]{roe85, hundsdorfer94} to preserve sharp gradients and edges |
|
|
that are typical of sea ice distributions and to rule out unphysical |
|
|
over- and undershoots (negative thickness or concentration). These |
|
|
scheme conserve volume and horizontal area and are unconditionally |
|
|
stable, so that we can set $D_{X}=0$. \ml{[do we need to proove that? |
|
|
can we proove that? citation?]} |
|
|
|
|
|
There is considerable doubt about the reliability of such a simple |
|
|
thermodynamic model---\citet{semtner84} found significant errors in |
|
|
phase (one month lead) and amplitude ($\approx$50\%\,overestimate) in |
|
|
such models---, so that today many sea ice models employ more complex |
|
|
thermodynamics. A popular thermodynamics model of \citet{winton00} is |
|
|
based on the 3-layer model of \citet{semtner76} and treats brine |
|
|
content by means of enthalphy conservation. This model requires in |
|
|
addition to ice-thickness and compactness (fractional area) additional |
|
|
state variables to be advected by ice velocities, namely enthalphy of |
|
|
the two ice layers and the thickness of the overlying snow layer. |
|
|
\ml{[Jean-Michel, your turf: ]Care must be taken in advecting these |
|
|
quantities in order to ensure conservation of enthalphy. Currently |
|
|
this can only be accomplished with a 2nd-order advection scheme with |
|
|
flux limiter \citep{roe85}.} |
|
|
|
|
|
|
|
|
\subsection{C-grid} |
|
|
\begin{itemize} |
|
|
\item no-slip vs. free-slip for both lsr and evp; |
|
|
"diagnostics" to look at and use for comparison |
|
|
\begin{itemize} |
|
|
\item ice transport through Fram Strait/Denmark Strait/Davis |
|
|
Strait/Bering strait (these are general) |
|
|
\item ice transport through narrow passages, e.g.\ Nares-Strait |
|
|
\end{itemize} |
|
|
\item compare different advection schemes (if lsr turns out to be more |
|
|
effective, then with lsr otherwise I prefer evp), eg. |
|
|
\begin{itemize} |
|
|
\item default 2nd-order with diff1=0.002 |
|
|
\item 1st-order upwind with diff1=0. |
|
|
\item DST3FL (SEAICEadvScheme=33 with diff1=0., should work, works for me) |
|
|
\item 2nd-order wit flux limiter (SEAICEadvScheme=77 with diff1=0.) |
|
|
\end{itemize} |
|
|
That should be enough. Here, total ice mass and location of ice edge |
|
|
is interesting. However, this comparison can be done in an idealized |
|
|
domain, may not require full Arctic Domain? |
|
|
\item |
|
|
Do a little study on the parameters of LSR and EVP |
|
|
\begin{enumerate} |
|
|
\item convergence of LSR, how many iterations do you need to get a |
|
|
true elliptic yield curve |
|
|
\item vary deltaTevp and the relaxation parameter for EVP and see when |
|
|
the EVP solution breaks down (relative to the forcing time scale). |
|
|
For this, it is essential that the evp solver gives use "stripeless" |
|
|
solutions, that is your dtevp = 1sec solutions/or 10sec solutions |
|
|
with SEAICE\_evpDampC = 615. |
|
|
\end{enumerate} |
|
|
|
|
|
\end{itemize} |
|
| 64 |
|
|
| 65 |
%%% Local Variables: |
%%% Local Variables: |
| 66 |
%%% mode: latex |
%%% mode: latex |
Parent Directory
| 
Revision Log
| 
Revision Graph
| 
Patch