-
Earlier IOBC approaches treated
$ {f}_{\mathrm{w}} $ analogous to momentum fluxwhere
$\,{\rho }_{w}$ is the seawater density,${c}_{{\rm{w}}}$ is the specific heat capacity of liquid water,$ {T}_{\mathrm{f}\mathrm{m}\mathrm{i}\mathrm{x}} $ is the ocean freezing point,$ {h}_{\mathrm{m}\mathrm{i}\mathrm{x}} $ is the mixed layer (ML) depth, and$ {T}_{\mathrm{m}\mathrm{i}\mathrm{x}} $ is the ML temperature. Here,$ {T}_{\mathrm{f}\mathrm{m}\mathrm{i}\mathrm{x}} $ is taken as a linear function of the ML salinity$ {S}_{\mathrm{m}\mathrm{i}\mathrm{x}} $ :where
$\, \mu =0.054 $ is an empirical coefficient for salinity and ocean freezing temperature. The momentum model assumes that any excess heat absorbed in seawater of an ice-covered ML is almost instantaneously used by sea ice melt. Such an IO system is often referred to as an “ice bath” (e.g., Josberger, 1983; Mellor et al., 1986). The ice-bath paradigm leads to rapid ocean heat loss from the ML to the sea ice above (e.g., Røed, 1984, Fichefet et al., 1997); thus, little heat energy can be stored in the ML below sea ice for extended periods.However, direct measurements from the Marginal Ice Zone Experiments (MIZE) in 1984 (McPhee et al., 1987) and observation-based studies (Krishfield and Perovich, 2005) have reported that the ice-covered upper Arctic Ocean stores significant amounts of heat for extended periods during the summer months when
$ {T}_{\mathrm{m}\mathrm{i}\mathrm{x}} $ is greater than$ {T}_{\mathrm{f}\mathrm{m}\mathrm{i}\mathrm{x}} $ . Such heat storage has also been observed in earlier laboratory studies (e.g., Yaglom and Kader, 1974). This evidence demonstrated that, IO heat exchanges are more likely governed by the turbulent heat flux or heat/salt molecular diffusion occurring in a very thin viscous sublayer below the sea ice bottom, within which the turbulent heat process or heat/salt molecular diffusion played a significant role. This is different from the momentum flux at the ocean ML. This view was further bolstered from the one-year-long Arctic Ice Dynamic Joint Experiment (AIDJEX, Maykut and McPhee, 1995), the Surface Heat Budget of the Arctic-SHEBA (Uttal et al., 2002), and short summer projects (McPhee, 2002). Motivated by these studies, turbulent heat flux and molecular diffusion models have been proposed (e.g., McPhee, 1992; Maykut and McPhee, 1995; Notz et al., 2003; Schmidt et al., 2004) to parametrize the heat budget at the IO interface. -
In the turbulent model,
$ {f}_{\mathrm{w}} $ is parametrized based on the Reynolds averaged turbulent heat flux:${f}_{\mathrm{w}}={\rho }_{\mathrm{w}}{c}_{\mathrm{w}}⟨{w}{{'}}{T}{{'}}⟩$ . Using dimensional analysis, the basic formulation of the turbulent model can be written as follows:where
$ {\alpha }_{\mathrm{t}} $ is the turbulent heat transfer coefficient;$ {u}_{*} $ is the friction velocity;$ {T}_{\mathrm{f}\mathrm{i}\mathrm{o}} $ is the freezing temperature of seawater.$ {T}_{\mathrm{f}\mathrm{i}\mathrm{o}}-{T}_{\mathrm{m}\mathrm{i}\mathrm{x}} $ is the temperature difference between the IO interface and the ML ocean. Compared with the momentum “ice bath” model, the amount of heat exchange at the IO is reduced because the value is scaled by$ {u}_{*} $ and$ {\alpha }_{{\rm{t}}} $ . Applying the turbulent heat flux model given by Eq. (4) to the Stefan condition [Eq. (1)], we obtain the following equation:Depending on the definition of the seawater freezing point,
$ {T}_{\mathrm{f}\mathrm{i}\mathrm{o}} $ , the turbulent model can appear as a one-, two- or three-equation approach.In the 1-equation approach,
$ {T}_{\mathrm{f}\mathrm{i}\mathrm{o}} $ is simply set to a constant value (i.e., –1.8°C). In the two-equation approach,$ {T}_{\mathrm{f}\mathrm{i}\mathrm{o}} $ is defined by Eq. (3), which forms the two-equation approach together with Eq. (5). The two-equation approach has been used to estimate the oceanic heat flux in observational studies (e.g., McPhee, 1992; Maykut and McPhee, 1995; Krishfield and Perovich, 2005), and because of its simplicity for computation, it is the default IO parametrization scheme in several sea ice models, i.e., CICE6, the Global Sea Ice Component (GSI8.1) or the Louvain-la-Neuve Sea Ice Model (LIM2, LIM3). Using Eq. (3) to define$ {T}_{\mathrm{f}\mathrm{i}\mathrm{o}} $ results in fast salt flux exchanges, identical to the turbulent heat exchanges between the IO interface and the ML. However, both theoretical studies and observations have shown that the opposite is true: the transfer rate of salt is much slower than that of heat (e.g., Woods, 1992; Notz et al., 2003; McPhee et al., 2008). To realistically represent the salinity effects in the models,$ {T}_{\mathrm{f}\mathrm{m}\mathrm{i}\mathrm{x}} $ and$ {S}_{\mathrm{m}\mathrm{i}\mathrm{x}} $ in Eq. (3) are replaced with$ {T}_{\mathrm{f}\mathrm{i}\mathrm{o}} $ and actual interfacial salinity$ {S}_{\mathrm{i}\mathrm{o}} $ , respectively:To solve for
$ {S}_{\mathrm{i}\mathrm{o}} $ , an equation describing the IO boundary salt flux balance is added to the two-equation scheme:where
$ {S}_{\mathrm{i}} $ is the salinity of the bottom sea ice, and$ {\alpha }_{\mathrm{s}} $ is the turbulent salt transfer coefficient. The above Eqs. (5), (6), and (7) constitute the three-equation boundary approach (e.g., Holland and Jenkins, 1999). In the three-equation approach,$ {T}_{\mathrm{f}\mathrm{i}\mathrm{o}} $ can be greater than the$ {T}_{\mathrm{m}\mathrm{i}\mathrm{x}} $ under a relatively small$ {S}_{\mathrm{i}\mathrm{o}} $ , in this case a downward heat flux from the sea ice bottom to the ocean ML can happen (e.g., Notz et al., 2003; McPhee et al., 2008; Shi et al., 2021). Considering the IO brine energy balance during sea ice melt or freezing, Schmidt et al. (2004) replaced the latent heat of fusion,$ L $ in Eq. (5), with energy changes$\Delta {E}'$ during the sea ice melt or growth process:where,
$\Delta {E}'={E}_{0}'({T}_{\mathrm{f}\mathrm{i}\mathrm{o}},{S}_{\mathrm{i}\mathrm{o}})-{E}_{\mathrm{i}}'({T}_{\mathrm{i}\mathrm{b}},{S}_{\mathrm{i}\mathrm{b}})$ ,${E}_{0}'$ is the energy of the seawater,${E}_{\mathrm{i}}'$ is the internal energy of sea ice, and$ {T}_{\mathrm{i}\mathrm{b}} $ and$ {S}_{\mathrm{i}\mathrm{b}} $ are the temperature and salinity of the volume transferred at the IO interface. Replacing the$ {S}_{\mathrm{i}} $ in Eq. (7) with$ {S}_{\mathrm{i}\mathrm{b}} $ , we obtain:and the values of
$ {T}_{{\rm{ib}}} $ and$ {S}_{{\rm{ib}}} $ depend on ice melt ($ \dot{h} > 0 $ ) or new ice formation ($ \dot{h} < 0 $ ):where
$ {f}_{\mathrm{s}} $ is the fraction of the boundary salinity initially retained within the ice. In this study, we used the three-equation system by Eqs. (6), (8), (9), and (10) to parametrize the IO boundary condition to examine the impacts of a three-equation approach on IO heat budgets in CICE. Note that the energy is defined by per mass (J kg–1) and melt rate is in units of kg m–2 s–1 as in Schmidt et al. (2004), while in CICE, the energy is defined by per volume (J m–3) and the unit of melt rate$ \dot{h} $ has unit of m s–1.To our knowledge, in the CMIP6 (Coupled Model Intercomparison Project Phase 6) models, the three-equation approach is the IO boundary condition parametrization in the Goddard Institute for Space Studies (GISS) ModelE and the two-equation approach is used in models that treat the sea ice component with CICE5 (i.e., CESM2-CAM, the CESM2-WACCAM, the NorESM-LL, and the NorESM-MM), the GSI8.1 (i.e., UKESM1-0-LL, HadGEM3-GC31-LL, HadGEM3-GC31-MM, ACCESS-CM2,) or the LIM (i.e., AEMET_EC-Earth3, SMHI_EC-Earth3-Veg, IPSL-CM6A-LR, and CanESM5).
-
Finally, we provide a short description of double molecular diffusion. As demonstrated above, the three-equation boundary condition induces a much fresher IO interface; because of this, double-diffusive convection of heat and salt occurs between the fresh IO water and underlying saltwater. Previous studies (i.e., Notz et al., 2003; McPhee et al., 2008; Tsamados et al., 2015; Shi et al., 2021) showed that the choice of the ratio R (=
$ {\alpha }_{\mathrm{t}}/{\alpha }_{\mathrm{s}}) $ , where$ {\alpha }_{\mathrm{t}} $ and$ {\alpha }_{\mathrm{s}} $ are the heat and salt flux diffusion rate, respectively, has important effects on the sea ice states. Considering that$ {\alpha }_{\mathrm{t}} $ is faster than$ {\alpha }_{\mathrm{s}} $ in the sea ice melting process, Notz et al. (2003) used an R$ =35 $ to model the observed summer false bottom persistence (Jeffries et al., 1995; Eicken et al., 2002; Perovich et al., 2003). McPhee et al. (2008) pointed out that an R used under melting conditions is inappropriate for the ice growth process. In some previous studies based on CICE, R was set to 35 (Shi et al., 2021) or 50 (Tsamados et al., 2015) for the melt condition and was set to 1 ($ {\alpha }_{\mathrm{t}} $ =$ {\alpha }_{\mathrm{s}} $ ) under ice growth conditions. In this study, following Tsamados et al. (2015), we set the ratio R = 50 for melt and R = 1 for growth conditions. -
The sea ice model used in this study is the CICE6. There are many physical parameterization schemes in CICE6 for users to turn on and off according to specific configurations, and detailed descriptions can be found in Hunke et al. (2013); Craig et al. (2018). A brief description follows. We use mushy thermodynamics in both the two-equation and the three-equation runs in this study because the three-equation scheme needs to operate with a mushy sea ice layer (Feltham et al., 2006), where the mush enthalpy is the combination of the brine and pure ice enthalpy according to the brine fraction. The model uses multiple ice thickness categories compatible with the ice thickness redistribution scheme (ITD) of Lipscomb et al. (2007), and we chose five ice thickness categories in this study. The level-ice formulation pond scheme is used for melt pond parametrizations, by which the ponds are carried as tracers on the level ice area of each thickness category. For the dynamics option, we chose the elastic-viscous-plastic (EVP) rheology in all model runs, as described in Hunke and Dukowicz (2002).
We ran the CICE6 in stand-alone mode and coupled the ML that forms part of the CICE6 package. The ML depth is fixed at 10 m in summer and 20 m in winter, the ML salinity had a fixed seasonal circle obtained from climatology data, and the SST evolved through the IO heat exchange. We apply a slow (15 days) temperature restoration of the ML temperature toward the monthly climatology of the reanalysis SSTs in our model runs. The climatology of the monthly atmospheric dataset from the Modern-Era Retrospective analysis for Research and Applications version 2 (Gelaro et al., 2017) and monthly reanalysis ocean data from ECMWF-ORAS5 (Zuo et al., 2019) are used to force the model. The input field consists of the 10 m wind velocity, 10 m specific humidity, incoming shortwave/long radiation, snowfall/rainfall rate, air density, sea level pressure (SLP), and SST. The model integrations reach equilibrium after ~20 model years and run continuously for 100 years. The results presented in this paper are the mean values over the last 50 model years.
-
In our reference run (REF), the IO boundary condition is parameterized by the two-equation scheme [Eqs. (2) and (3)]. We set
$ {\alpha }_{\mathrm{t}} $ to a constant value of 0.006 and parametrized the thermal conductivity according to Maykut and Untersteiner [1971, hereafter, we refer to this conductivity method as (MU71)].We take the three-equation runs as our sensitivity runs, in which the IO boundary condition was parametrized using the three-equation scheme [Eqs. (6) to (10)] with several combinations of coefficient
$ {\alpha }_{t} $ and ice conductivity$ \lambda $ . For the$ {\alpha }_{\mathrm{t}} $ , we used the constant (= 0.006) and form-drag-based$ {\alpha }_{t} $ by switching the form drag parametrization off/on (Tsamados et al., 2014) in CICE6, where the form-drag-based$ {\alpha }_{\mathrm{t}} $ is accounted for by half of the ocean form drag coefficient under melt conditions (Tsamados et al., 2015). Then we combined the two kinds of$ {\alpha }_{\mathrm{t}} $ to the MU71 and the conductivity of Pringle et al. (2007, hereafter, we refer to this conductivity method as P07) in our three-equation sensitivity runs. Table 1 lists the name and coefficient combination for each model run used in this study. The oceanic heat flux$ {f}_{\mathrm{w}} $ has a linear dependency on the friction velocity$ {u}_{*} $ , it can be calculated by$ {u}_{*}=\sqrt{{\tau }_{\mathrm{w}}/{\rho }_{\mathrm{w}}} $ (where$ {\tau }_{\mathrm{w}} $ is the IO drag) in CICE6. To exclude the effects of constant and form drag parametrization on$ {u}_{*} $ , we fixed$ {u}_{*} $ at 0.002 (Notz et al., 2003) in all our model runs. Finally, we ran our model experiments using a grid resolution of 1o on the global domain and focused our analysis on the Arctic region.Name IO boundary Heat/salt transfer coef. Ice conductivity REF two-equation Constant drag, $ {\alpha }_{\mathrm{t}}=0.006 $, $ {\alpha }_{\mathrm{s}}={\alpha }_{\mathrm{t}}/50 $ Maykut and Untersteiner CST-DRAG-MU71 three-equation Constant drag, $ {\alpha }_{\mathrm{t}}=0.006 $, $ {\alpha }_{\mathrm{s}}={\alpha }_{\mathrm{t}}/50 $ Maykut and Untersteiner FORM-DRAG-MU71 three-equation Form drag (${ {\rm{Cdn} } }\_{\mathrm{o}\mathrm{c}\mathrm{n} }$), ${\alpha }_{\mathrm{t} }={{\rm{Cdn}}}\_{\mathrm{o}\mathrm{c}\mathrm{n} }/2$, $ {\alpha }_{\mathrm{s}}={\alpha }_{\mathrm{t}}/50 $ Maykut and Untersteiner CST-DRAG-P07 three-equation Constant drag, $ {\alpha }_{\mathrm{t}}=0.006 $, $ {\alpha }_{\mathrm{s}}={\alpha }_{\mathrm{t}}/50 $ Pringle et al.(2007) FORM-DRAG-P07 three-equation Form drag (${ {\rm{Cdn} } }\_{\mathrm{o}\mathrm{c}\mathrm{n} }$), ${\alpha }_{\mathrm{t} }={{\rm{Cdn}}}\_{\mathrm{o}\mathrm{c}\mathrm{n} }/2$, $ {\alpha }_{\mathrm{s}}={\alpha }_{\mathrm{t}}/50 $ Pringle et al.(2007) Table 1. The name, IO boundary condition approaches, heat (
$ {\alpha }_{\mathrm{t}} $ ), salt ($ {a}_{\mathrm{s}} $ ) transfer coefficient, and ice conductivity ($ \lambda ) $ for each model run. -
This section describes the impact of the two-equation and three-equation approaches. In CICE6, all heat fluxes are positive (negative) downwards (upwards). To avoid confusing “increase” and “decrease” of negative values, we refer to the upward oceanic heat flux as
$-{f}_{\mathrm{w}}$ (positive, Figs. 1–4, 6–9). Accordingly, we use the temperature difference$ \Delta {T}_{\mathrm{o}} $ between the ocean ML (sea surface temperature, SST) and the freezing temperature ($ {T}_{\mathrm{f}} $ ,$ \Delta {T}_{\mathrm{o}} $ ${={\rm{SST}}-T}_{\mathrm{f}}$ ) to explain the simulated upward oceanic heat flux$-{f}_{\mathrm{w}}$ . Note that$ {T}_{\mathrm{f}} $ is taken as the$ {T}_{\mathrm{f}\mathrm{m}\mathrm{i}\mathrm{x}} $ in the REF, while it is defined as$ {T}_{\mathrm{f}\mathrm{i}\mathrm{o}} $ in the three-equation runs. We look first at the general responses of the main results to the two approaches. Figure 1 shows the seasonal cycles of several thermodynamic processes at the IO interface in each run, including the basal melt, interfacial salinity$ {S}_{\mathrm{i}\mathrm{o}} $ and freezing temperature, temperature difference$ \Delta {T}_{\mathrm{o}} $ , upward oceanic heat flux ($-{f}_{\mathrm{w}}$ ), temperature difference$ \Delta {T}_{\mathrm{b}} $ (between the sea ice temperature of the bottom layer ($ {T}_{\mathrm{i}\mathrm{b}} $ ) and the freezing temperature (Tf ):$ \Delta {T}_{\mathrm{b}} $ =$ {T}_{\mathrm{i}\mathrm{b}}-{T}_{\mathrm{f}} $ , downward bottom heat conduction$ {f}_{\mathrm{c}\mathrm{o}\mathrm{n}\mathrm{d}\mathrm{b}\mathrm{o}\mathrm{t}} $ , the net IO heat exchange by$ {f}_{\mathrm{c}\mathrm{o}\mathrm{n}\mathrm{d}\mathrm{b}\mathrm{o}\mathrm{t}} $ –$ {f}_{\mathrm{w}} $ , congelation, and sea ice thickness. We start with the responses of basal melt (Fig. 1a) and interfacial salinity (Fig. 1b). In all our model runs, the basal melt speeds up beginning in June, reaches a maximum in July, and remains high until August. The acceleration of the basal melt leads to a large release of freshwater, freshening the IO interface directly below the sea ice bottom in each three-equation run. From Fig. 1a, we can see the significant reduction in the interfacial salinity$ {S}_{\mathrm{i}\mathrm{o}} $ during the summer months (June−July−August, JJA). Note that the maximum melt rate is in July, but the minimum$ {S}_{\mathrm{i}\mathrm{o}} $ occurs in August. This reflects an accumulated freshening process at the IO interface from June until August. For the two-equation REF run, there is no interfacial freshening process since the interfacial salinity$ {S}_{\mathrm{i}\mathrm{o}} $ is used as the salinity of the ML ($ {S}_{\mathrm{i}\mathrm{o}} $ =$ {S}_{\mathrm{m}\mathrm{i}\mathrm{x}} $ ) in the two-equation approach.Figure 1. Impacts of the two-equation (REF) and three-equation approach on (a) basal melt, (b) interfacial salinity (
$ {S}_{\mathrm{i}\mathrm{o}} $ ), (c) interfacial temperature ($ {T}_{\mathrm{f}\mathrm{i}\mathrm{o}} $ ), (d) temperature difference between SST and IO interface ($ \Delta {T}_{\mathrm{o}} $ ), (e) upward heat flux ($-{f}_{\mathrm{w}}$ ), (f) temperature difference between ice bottom and IO ($ \Delta {T}_{\mathrm{b}} $ ), (g) bottom heat conduction ($ {f}_{\mathrm{c}\mathrm{o}\mathrm{n}\mathrm{d}\mathrm{b}\mathrm{o}\mathrm{t}} $ ), (h) net IO heat flux ($ {f}_{\mathrm{c}\mathrm{o}\mathrm{n}\mathrm{d}\mathrm{b}\mathrm{o}\mathrm{t}} $ –$ {f}_{\mathrm{w}} $ ), (i) congelation, (j) top melt, (k) lateral melt, and (l) sea ice thickness. Units are shown on the y-axis.Figure 2. The simulated bottom thermal process in August in the two-equation run (REF). (a) ML salinity (
$ {S}_{\mathrm{m}\mathrm{i}\mathrm{x}} $ , PSU), (b) ocean freezing temperature ($ {T}_{\mathrm{f}\mathrm{m}\mathrm{i}\mathrm{x}} $ , °C ), (c) temperature difference ($\Delta {T}_{{\rm{b}}}$ ,°C) between ice bottom and$ {T}_{\mathrm{f}\mathrm{m}\mathrm{i}\mathrm{x}} $ , (d) temperature difference ($ \Delta {T}_{\mathrm{o}} $ , °C) between SST and$ {T}_{\mathrm{f}\mathrm{m}\mathrm{i}\mathrm{x}} $ , (e) bottom heat conduction ($ {f}_{\mathrm{c}\mathrm{o}\mathrm{n}\mathrm{d}\mathrm{b}\mathrm{o}\mathrm{t}} $ , W m–2), (f) upward oceanic heat flux($-{f}_{\mathrm{w}}$ , W m–2), (g) net IO heat exchanges ($ {f}_{\mathrm{c}\mathrm{o}\mathrm{n}\mathrm{d}\mathrm{b}\mathrm{o}\mathrm{t}} $ –$ {f}_{\mathrm{w}} $ , W m–2), (h) congelation (cm d–1) and (i) basal melt rate (cm d–1). The black line in panel (d) denotes the 15% SIC contour.Figure 3. The simulated bottom thermal process in August in the three-equation sensitivity run (CST-DRAG-MU71). (a) Interfacial salinity (
$ {S}_{\mathrm{i}\mathrm{o}} $ , PSU), (b) interfacial temperature ($ {T}_{\mathrm{f}\mathrm{i}\mathrm{o}} $ ,°C), (c) temperature difference ($ \Delta {T}_{\mathrm{b}} $ , °C) between ice bottom and$ {T}_{\mathrm{f}\mathrm{i}\mathrm{o}} $ , (d) temperature difference ($ \Delta {T}_{\mathrm{o}} $ , °C) between SST and$ {T}_{\mathrm{f}\mathrm{i}\mathrm{o}} $ , (e) bottom heat conduction ($ {f}_{\mathrm{c}\mathrm{o}\mathrm{n}\mathrm{d}\mathrm{b}\mathrm{o}\mathrm{t}} $ , W m–2), (f) upward oceanic heat flux ($-{f}_{\mathrm{w}}$ , W m–2), (g) net IO heat exchanges ($ {f}_{\mathrm{c}\mathrm{o}\mathrm{n}\mathrm{d}\mathrm{b}\mathrm{o}\mathrm{t}} $ $-{f}_{\mathrm{w}}$ , W m–2), (h) congelation (cm d–1) and (i) basal melt rate (cm d–1). The black line in panels (d), (f), and (h) denotes the 15% SIC contour.Figure 4. The difference in (a) upward oceanic heat flux (
$-{f}_{\mathrm{w}}$ , W m–2), (b) basal melt rate (cm d–1), (c) sea ice thickness (m), and (d) SST (°C) in August between each of the three-equation sensitivity runs and the REF in August. In each panel, (1) denotes CST-DRAG-MU71 minus REF, (2) FORM-DRAG-MU71 minus REF (3) CST-DRAG-P07 minus REF, and (4) FORM-DRAG-P07 minus REF.In each three-equation run, a common lower
$ {S}_{\mathrm{i}\mathrm{o}} $ results in a common greater IO freezing temperature$ {T}_{\mathrm{f}\mathrm{i}\mathrm{o}} $ above the oceanic freezing temperature$ {T}_{\mathrm{f}\mathrm{m}\mathrm{i}\mathrm{x}} $ (Fig. 1c). This greater$ {T}_{\mathrm{f}\mathrm{i}\mathrm{o}} $ impacts the heat balance by changing the temperature differences at the IO interface. First, the increase in$ {T}_{\mathrm{f}\mathrm{i}\mathrm{o}} $ reduces the temperature differences ($ \Delta {T}_{\mathrm{o}} $ ; Fig. 1d), and thus reduces the upward oceanic heat flux ($-{f}_{\mathrm{w}}$ ; Fig. 1e). Meanwhile, it leads to a reduced temperature difference between the ice bottom and the$ {T}_{{\rm{fio}}} $ ($ \Delta {T}_{\mathrm{b}} $ ; Fig. 1f), and hence reduces the downward heat conduction ($ {f}_{\mathrm{c}\mathrm{o}\mathrm{n}\mathrm{d}\mathrm{b}\mathrm{o}\mathrm{t}} $ ; Fig. 1g). Both the decrease in downward$ {f}_{\mathrm{c}\mathrm{o}\mathrm{n}\mathrm{d}\mathrm{b}\mathrm{o}\mathrm{t}} $ and upward heat flux ($-{f}_{\mathrm{w}}$ ) then resulted in a weakened heat energy convergence (${f}_{\mathrm{c}\mathrm{o}\mathrm{n}\mathrm{d}\mathrm{b}\mathrm{o}\mathrm{t}}{-f}_{\mathrm{w}}$ ; Fig. 1h) at the IO interface and consequently, slowed the basal melt rate (Fig. 1a). Furthermore, ice congelation occurs in August in CST_DRAG_MU71 and CST_DRAG_P07 (Fig. 1i), indicating a heat divergence at the IO interface (Fig. 3g), while no ice congelation is found in other experiments.Finally, the sea ice thickness (Fig. 1l) increased in each three-equation run compared with the REF in summer (JJA). Comparison among the bottom (Fig. 1a), top (Fig. 1j), and lateral (Fig. 1k) melt rates show that the decrease in the basal melt is the major contributor to the increase in the sea ice thickness: the mean anomaly of bottom melt is ~ –0.5 cm d–1 (1.2–1.5 cm d–1 in three-equation runs against ~1.8 cm d–1 in REF), while that for top melt is ~ –0.2 cm d–1 (1.3–1.4 cm d–1 against ~1.5 cm d–1), and no pronounced differences are observed for the lateral melt.
The above results show a common lower
$ {S}_{\mathrm{i}\mathrm{o}} $ in all the three-equation runs in comparison to that in the REF run. This increases the IO freezing temperature and thus reduces oceanic turbulent heat flux and slows basal melt rate. Finally, we find a common increase in the sea ice thickness during summer compared with the two-equation REF run.After analyzing the general differences among the seasonal cycles between the three-equation and two-equation approaches, we now compare the differences among the three-equation runs. The results show that the most significant differences among the three-equation runs occur in the summer months (JJA) since the heat absorption from the incoming solar shortwave radiation in the ocean ML is significantly increased, which, in turn, causes more significant changes in the upward oceanic heat flux (Fig. 1e). First, for the differences between the constant and form-drag-based
$ {\alpha }_{\mathrm{t}} $ runs (CST-DRAG-MU71 vs. FORM-DRAG-MU71, CST-DRAG-P07 vs. FORM-DRAG-P07), we can see that the basal melt rate is faster in the constant$ {\alpha }_{\mathrm{t}} $ runs than that in the form-drag-based$ {\alpha }_{\mathrm{t}} $ model runs (Fig. 1a). The mean basal melt rates during JJA are ~1.5 cm d–1 and ~1.3 cm d–1 in the CST-DRAG-MU71 and CST-DRAG-P07, faster than the melt rate of ~1.25 cm d–1 and ~1.20 cm d–1 in the FORM_DRAG_MU71 and FORM_DRAG_P07, respectively. This faster basal melt rate can be attributed to the fact that the constant$ {\alpha }_{\mathrm{t}} $ (0.006) is larger than the form-drag-based$ {\alpha }_{\mathrm{t}} $ in most regions of the Arctic basin except for the heavily ridged regions north of Greenland. It amplifies the upward oceanic heat flux ($-{f}_{\mathrm{w}}$ ; Fig. 1e) in the constant$ {\alpha }_{\mathrm{t}} $ run in comparison to that in the form-drag-based$ {\alpha }_{\mathrm{t}} $ and, therefore, causes a faster basal melt rate. As a result, the$ {S}_{\mathrm{i}\mathrm{o}} $ is lower, and hence the$ {T}_{\mathrm{f}\mathrm{i}\mathrm{o}} $ is greater in the constant$ {\alpha }_{\mathrm{t}} $ runs as compared with those in the form-drag-based$ {\alpha }_{\mathrm{t}} $ model runs (Fig. 1b). The relative faster bottom rate and greater upward oceanic heat flux result in the thinner sea ice thickness in the constant$ {\alpha }_{\mathrm{t}} $ runs (Fig. 1l). Again, from Fig. 1i, the ice congelation in August occurs only in the constant$ {\alpha }_{\mathrm{t}} $ runs (CST-DRAG-MU71 and CST-DRAG-P07) are most notable. Further analysis in section 3.5 shows that this ice congelation in August lies in the region of downward heat flux from the ice to the ocean (Fig. 8), which causes a heat divergence at the IO interface and leads to the congelation as a consequence. This important difference will be addressed by detailed analysis based on the spatial pattern maps in section 3.5. For the differences between the conductivity in the MU71 and P07 runs, we can see that the upward heat flux ($-{f}_{\mathrm{w}}$ ) is less in the P07 runs (Fig. 1e) than that in the conductivity MU71 runs, it is associated with relatively slower basal melt rates (Fig. 1a), greater$ {S}_{\mathrm{i}\mathrm{o}} $ (Fig. 1b), and lower$ {T}_{\mathrm{f}\mathrm{i}\mathrm{o}} $ (Fig. 1c), and hence leads to a greater thickness (Fig.1l) in each P07 run (CST-DRAG-P07 vs. CST-DRAG-MU71, FORM-DRAG-P07 vs. CST-DRAG-MU71). A previous study (Hunke, 2010), based on the two-equation approach, reported that the conductivity P07 decreases the e-folding scale of the ice thickness redistribution function, reduces the ocean heating on the sea ice, and adjusts the albedo upward, each of which tends to increase the sea ice thickness. In this study, the increased sea ice thickness is also be found in our three-equation runs by the conductivity P07, as discussed above. Further analysis will be given in section 3.5.The above results indicate that August has the largest differences between the two approaches, such as the interfacial salinity, interfacial freezing temperature, and IO heat exchanges. This can be attributed to the insolation effect in that more background insolation leads to more pronounced changes in the IO heat flux between the two approaches. During the summer months (JJA), less surface ablation due to reduced sea ice concentration (SIC) and thinner sea ice thickness allow more incoming shortwave radiation to penetrate through the sea ice. Meanwhile, the ocean ML is also absorbing more shortwave radiation. This insolation effect increases both the upward oceanic heat flux (
$ -{f}_{\mathrm{w}} $ ) and the downward heat conduction ($ {f}_{\mathrm{c}\mathrm{o}\mathrm{n}\mathrm{d}\mathrm{b}\mathrm{o}\mathrm{t}} $ ), where the upward oceanic heat flux$ -{f}_{\mathrm{w}} $ determines the net heat flux at the IO interface (~50 W m−2 of the$ -{f}_{\mathrm{w}} $ vs. ~4 W m−2 of the$ {f}_{\mathrm{c}\mathrm{o}\mathrm{n}\mathrm{d}\mathrm{b}\mathrm{o}\mathrm{t}} $ on average, see Fig. 1e, Fig. 1g, and Fig. 1h). In August, the$ -{f}_{\mathrm{w}} $ reaches the greatest magnitude (Fig. 1e), causing the largest changes in the related bottom process among our model runs for this month. In the next section, we focus our study on the August spatial maps of the absolute values from the REF and CST-DRAG-MU71 (the run of the lowest$ {S}_{\mathrm{i}\mathrm{o}} $ , the greatest$ {T}_{\mathrm{f}\mathrm{i}\mathrm{o}} $ and the summer bottom congelation) to study how the two approaches impact the IO heat exchanges and associated basal melt/growth. Then, we will compare the differences among our three-equation runs to assess how different results were obtained by different combinations of$ {\alpha }_{\mathrm{t}} $ and thermal conductivities. -
Figure 2 shows the spatial maps of the August ML salinity, thermal differences, basal melt, and congelation in the REF. From Figs. 2a and 2b, we see an opposite linear dependency of
$ {T}_{\mathrm{f}\mathrm{m}\mathrm{i}\mathrm{x}} $ on$ {S}_{\mathrm{m}\mathrm{i}\mathrm{x}} $ : a greater (lower) salinity results in a lower (greater) ocean$ {T}_{\mathrm{f}\mathrm{m}\mathrm{i}\mathrm{x}} $ . In the REF, the minimal$ {T}_{\mathrm{f}\mathrm{m}\mathrm{i}\mathrm{x}} $ is ~ –1.8°C and located in the eastern Arctic Ocean. It increases to ~ –1.6°C to –1.4°C in the central Arctic Ocean and to ~ –1°C in the coastline regions of the East Siberian Sea and the Beaufort Sea (ESB). This freezing temperature$ {T}_{\mathrm{f}\mathrm{m}\mathrm{i}\mathrm{x}} $ results in an overall positive temperature difference between the ocean ML layer and IO interface ($ \Delta {T}_{\mathrm{o}} $ ; Fig. 2d), which is relatively greater (~1.0°C) along the sea ice edge (defined as the 15% SIC contour) regions than that in the central Arctic Ocean, the heavily ridged regions north of Greenland, and in some parts of the Canadian Archipelago (hereafter, we refer to this region as NGCA). Consequently, the positive$ \Delta {T}_{\mathrm{o}} $ between the ocean ML and the IO interface leads to an overall upward heat flux (–$ {f}_{\mathrm{w}} $ ; Fig. 2f) to the ice bottom, with magnitudes of 20–60 W m–2 and 2–10 W m–2 in the ice edge regions and the Arctic basin, respectively. Meanwhile, the$ {T}_{\mathrm{f}\mathrm{m}\mathrm{i}\mathrm{x}} $ results in an overall positive temperature difference between the ice bottom and IO interface ($ \Delta {T}_{\mathrm{b}} $ ; Fig. 2c), causing downward heat conduction$ {f}_{\mathrm{c}\mathrm{o}\mathrm{n}\mathrm{d}\mathrm{b}\mathrm{o}\mathrm{t}} $ in most of the Arctic Ocean (Fig. 2e).The upward oceanic heat flux (
$ -{f}_{\mathrm{w}} $ ) and downward$ {f}_{\mathrm{c}\mathrm{o}\mathrm{n}\mathrm{d}\mathrm{b}\mathrm{o}\mathrm{t}} $ result in heat convergence ($ {f}_{\mathrm{c}\mathrm{o}\mathrm{n}\mathrm{d}\mathrm{b}\mathrm{o}\mathrm{t}}{-f}_{\mathrm{w}} $ > 0; Fig. 2g) in the whole Arctic basin during summer in the REF. Accordingly, the basal melt is faster than 2.0 cm d−1 in the ice edge regions (Fig. 2i). Over the Arctic basin, the basal melt is ~ 0.7 cm d−1, on average, and less than 0.2 cm d−1 in the NGCA. No bottom growth occurred in August in the REF (Fig. 2h). Comparison between the magnitude of the upward oceanic heat flux ($ -{f}_{\mathrm{w}} $ ) and downward$ {f}_{\mathrm{c}\mathrm{o}\mathrm{n}\mathrm{d}\mathrm{b}\mathrm{o}\mathrm{t}} $ shows that the upward oceanic heat flux ($ -{f}_{\mathrm{w}} $ ) is the primary heat contributor to the basal melt in August. -
Section 3.2 described how the two-equation boundary condition influences the IO heat exchanges and the associated basal melt/growth in CICE6. Different from the two-equation approach, the three-equation approach treats the freezing temperature as the interfacial temperature
$ {T}_{\mathrm{f}\mathrm{i}\mathrm{o}} $ (rather than the$ {T}_{\mathrm{f}\mathrm{m}\mathrm{i}\mathrm{x}} $ ), which is dependent on the interfacial salinity$ {S}_{\mathrm{i}\mathrm{o}} $ (Eq. 6). Following the algorithm by Schmidt et al. (2004), the$ {S}_{\mathrm{i}\mathrm{o}} $ values were solved by the Newton-Raphson iteration method. We calculated those physical variables in Eqs. (6)–(10) for each ice thickness category and used their mean values at the IO boundary for our analysis. We have shown in Fig. 1 that the three-equation boundary condition results, on average, in a much lower$ {S}_{\mathrm{i}\mathrm{o}} $ compared with$ {S}_{\mathrm{m}\mathrm{i}\mathrm{x}} $ , indicating an IO freshening process induced by the three-equation approach. Here, the spatial map (Fig. 3a) shows that IO freshening mainly occurs in the ice edge regions where fast basal melt occurs (Fig. 3i), while in the central Arctic Ocean,$ {S}_{\mathrm{i}\mathrm{o}} $ is nearly identical to$ {S}_{\mathrm{m}\mathrm{i}\mathrm{x}} $ (~30 PSU), showing a relatively weaker IO freshening process in this region. This can be attributed to the fact that the much higher SIC in the center Arctic Ocean reduces the change in the radiative heat flux absorbed by the seawater. Hence, the changes in both oceanic heat flux and the basal melt rate are relatively smaller than the changes in the sea ice edge regions.Compared with the REF, the lower
$ {S}_{\mathrm{i}\mathrm{o}} $ in the ice edge regions results in a greater$ {T}_{\mathrm{f}\mathrm{i}\mathrm{o}} $ (Fig. 3b), which changes the heat exchanges at the IO interface. (1) For the upward oceanic turbulent heat flux ($ -{f}_{\mathrm{w}} $ ; Fig. 3f), the increased$ {T}_{\mathrm{f}\mathrm{i}\mathrm{o}} $ reduces the$ \Delta {T}_{\mathrm{o}} $ (Fig. 3d), causes an average decrease of ~25 W m−2 in upward oceanic heat flux ($ -{f}_{\mathrm{w}} $ ) in the ice edge regions (Fig. 4a), which demonstrates the heat-reducing role of the three-equation condition in the IO heat exchange. Furthermore, the increased$ {T}_{\mathrm{f}\mathrm{i}\mathrm{o}} $ results in a negative$ \Delta {T}_{\mathrm{o}} $ (SST <$ {T}_{\mathrm{f}\mathrm{i}\mathrm{o}} $ ) in areas inside the sea ice edges in the Chukchi Sea and East Siberian Sea (hereafter, we refer to this region as CES), leading to ~5 W m−2 of downward oceanic heat flux to the ocean ML in the CES region; (2) For the heat conduction$ {f}_{\mathrm{c}\mathrm{o}\mathrm{n}\mathrm{d}\mathrm{b}\mathrm{o}\mathrm{t}} $ (Fig. 3e), the increased$ {T}_{{\rm{fio}}} $ weakens the$ \Delta {T}_{\mathrm{b}} $ (Fig. 3c) and then decreases the$ {f}_{\mathrm{c}\mathrm{o}\mathrm{n}\mathrm{d}\mathrm{b}\mathrm{o}\mathrm{t}} $ mainly along the ice edge regions. Finally, the decreased upward oceanic heat flux and downward$ {f}_{\mathrm{c}\mathrm{o}\mathrm{n}\mathrm{d}\mathrm{b}\mathrm{o}\mathrm{t}} $ result in reduced heat convergence in this three-equation run, which slows the basal melt (Fig. 3i and Fig. 4b) and contributes to the increase in the sea ice thickness along the ice edge regions (Fig. 4c). Additionally, in the CES region, there is a heat divergence ($ {f}_{\mathrm{c}\mathrm{o}\mathrm{n}\mathrm{d}\mathrm{b}\mathrm{o}\mathrm{t}}-{f}_{\mathrm{w}} < 0 $ ), causing approximately 0.1~0.2 cm d−1 of bottom congelation in August (Fig. 3h).The above studies detected a heat-reducing impact of the three-equation boundary condition on the IO heat exchanges in the CST-DARG-MU71. The differences between the other sensitivity runs and the REF run show also a common decrease in the upward oceanic heat flux along the ice edge regions (Fig. 4a), which is associated with a decrease in basal melt (Fig. 4b) and an increase in sea ice thickness in each three-equation run (Fig. 4c).
-
Figure 4d shows the impact of the three-equation condition on the SST. Because of the absence of the changes in feedback between the ML and deep ocean in our stand-alone sensitivity runs with a prescribed ML depth, the SST changes are dependent mainly on the heat balance among the net atmospheric heat flux (
$ {f}_{\mathrm{s}\mathrm{u}\mathrm{r}\mathrm{f}} $ ) on the sea surface, the solar shortwave radiation that penetrates through the sea ice to the ocean ML ($ {f}_{\mathrm{s}\mathrm{w}\mathrm{t}\mathrm{h}\mathrm{r}\mathrm{u}} $ ) and the available oceanic heat flux ($ {f}_{\mathrm{o}\mathrm{c}\mathrm{n}} $ ) (accounted for as the available heat flux in the ocean ML after the energy consumption during the melt/sublimation process). From Fig. 4d, our results show that the SST is lower in the sea ice edge regions but close to or warmer in the central Arctic compared with the values observed in the REF (Fig. 4d). According to this SST anomaly (SSTA) spatial pattern, we calculated the mean values of the$ {f}_{\mathrm{s}\mathrm{u}\mathrm{r}\mathrm{f}} $ ,$ {f}_{\mathrm{s}\mathrm{w}\mathrm{t}\mathrm{h}\mathrm{r}\mathrm{u}} $ and$ {f}_{\mathrm{o}\mathrm{c}\mathrm{n}} $ in the ice edge regions (SIC less than 15%) and central Arctic Ocean (SIC greater than 30%). The results are presented in Fig. 5 and Table 2 (Table 3) for the ice edge (central Arctic) regions. Note that all heat fluxes are positive downwards, so the downward (upward) anomalies tend to enhance (reduce) the SST. Now we study how the heat balance anomaly causes these SSTAs. First, we study the absolute values in the REF (Figs. 5a, b). In the ice edge regions, there are an upward (negative)$ {f}_{\mathrm{o}\mathrm{c}\mathrm{n}} $ of ~ –23.6 W m–2 on average from the ocean ML, a downward atmospheric heat flux of$ {f}_{\mathrm{s}\mathrm{u}\mathrm{r}\mathrm{f}} $ (~28.9 W m–2 on average) and$ {f}_{\mathrm{s}\mathrm{w}\mathrm{t}\mathrm{h}\mathrm{r}\mathrm{u}} $ (~9.2 W m–2 on average). Accordingly, there is a downward (positive) net heat flux of 14.5 W m–2, heating the ocean ML in summer. For the central Arctic Ocean, due to the relative higher SIC, greater sea ice thickness, and reduced oceanic heat flux (Fig. 2f), all the three heat fluxes are much smaller (from –23.6 to –6.9 W m–2 in the$ {f}_{\mathrm{o}\mathrm{c}\mathrm{n}} $ , from 28.9 to 6.1 W m–2 in the$ {f}_{\mathrm{s}\mathrm{u}\mathrm{r}\mathrm{f}} $ , and from 9.2 to 4.5 W m–2 in the$ {f}_{\mathrm{s}\mathrm{w}\mathrm{t}\mathrm{h}\mathrm{r}\mathrm{u}} $ ) in comparison to those in the ice edge regions. As a result, the downward (positive) net heat flux is reduced to 3.6 W m–2 on average. The difference in the heat flux balance between the ice edge regions and the central Arctic Ocean shows that the summer heating on the SST by the three heat fluxes is greater in the ice edge regions than in the central Arctic Ocean.Figure 5. The August heat flux balance among the available oceanic heat flux (
$ {f}_{\mathrm{o}\mathrm{c}\mathrm{n}} $ ), the net atmospheric heat flux ($ {f}_{\mathrm{s}\mathrm{u}\mathrm{r}\mathrm{f}} $ ) and the shortwave radiation penetrated to the ocean ML through the sea ice ($ {f}_{\mathrm{s}\mathrm{w}\mathrm{t}\mathrm{h}\mathrm{r}\mathrm{u}} $ ). All heat fluxes are positive downwards and have units of W m–2. Absolute values in the REF in (a) the ice edge regions and (b) the central Arctic, and anomalies in the ice edge regions in (c) CST-DRAG-MU71, (d) FORM-DRAG-MU71, (e) CST-DRAG-P07, and (f) FORM-DRAG-P07. (g)–(j) as (c)–(f) but for the central Arctic.Name SST
(°C)SIT
(m)$ {f}_{\mathrm{o}\mathrm{c}\mathrm{n}} $
(W m–2)$ {f}_{\mathrm{s}\mathrm{u}\mathrm{r}\mathrm{f}} $
(Wm–2)$ {f}_{\mathrm{s}\mathrm{w}\mathrm{t}\mathrm{h}\mathrm{r}\mathrm{u}} $
(W m–2)Net flux
(W m–2)REF 0.25 0.28 –23.58 28.90 9.21 14.53 CST-DRAG-MU71 –0.17 0.50 4.45 –7.49 –5.80 –8.84 FORM-DRAG-MU71 –0.13 0.35 4.27 –6.05 –5.44 –7.22 CST-DRAG-P07 –0.18 0.61 5.43 –8.24 –7.78 –10.58 FORM-DRAG-P07 –0.18 0.43 5.17 –8.03 –7.24 –10.10 Table 2. The absolute mean values in REF and anomalies in the three-equation sensitivity runs of simulated SST, sea ice thickness (SIT), and the heat flux balance among the available oceanic heat flux (
$ {f}_{\mathrm{o}\mathrm{c}\mathrm{n}} $ ), the net atmospheric heat flux ($ {f}_{\mathrm{s}\mathrm{u}\mathrm{r}\mathrm{f}} $ ) and the shortwave radiation that penetrated through the sea ice ($ {f}_{\mathrm{s}\mathrm{w}\mathrm{t}\mathrm{h}\mathrm{r}\mathrm{u}} $ ) in the ice edge regions in August.Name SST
(°C)SIT
(m)$ {f}_{\mathrm{o}\mathrm{c}\mathrm{n}} $
(W m–2)$ {f}_{\mathrm{s}\mathrm{u}\mathrm{r}\mathrm{f}} $
(W m–2)$ {f}_{\mathrm{s}\mathrm{w}\mathrm{t}\mathrm{h}\mathrm{r}\mathrm{u}} $
(W m–2)Net flux
(W m–2)REF −1.86 3.43 –6.96 6.12 4.47 3.62 CST-DRAG-MU71 0.029 0.11 2.77 –1.02 –0.98 0.77 FORM-DRAG-MU71 0.031 0.09 2.85 –1.23 –0.80 0.83 CST-DRAG-P07 0.00 0.00 2.70 –1.35 –1.34 0.01 FORM-DRAG-P07 0.02 0.07 2.64 –1.27 –0.68 0.69 Table 3. Same as Table 2 but for the central Arctic.
Now we move to the anomalies in our three-equation sensitivity runs. We start with the heat balance in the ice edge regions, where we have found a common cooling (Fig. 4d) in all our three-equation sensitivity runs (~ –0.16oC on average, Table 2). First for the
$ {f}_{\mathrm{o}\mathrm{c}\mathrm{n}} $ , we can see a common anomalous downward (positive)$ {f}_{\mathrm{o}\mathrm{c}\mathrm{n}} $ (~ 4.8 W m–2 on average) in these sensitivity runs (Figs. 5c–f). This can be explained by the fact that in our three-equation sensitivity runs, the slower basal melt processes in the ice edge regions (Fig. 4b) consumes less upward oceanic heat flux ($-{f}_{\mathrm{w}}$ ; Fig. 4a), leading to more available heat energy retained in the ocean ML, as indicated by anomalous downward (positive)$ {f}_{\mathrm{o}\mathrm{c}\mathrm{n}} $ to the ocean ML. The anomalous positive$ {f}_{\mathrm{o}\mathrm{c}\mathrm{n}} $ tends to enhance the SST. However, the increased sea ice thickness in the ice edge regions (Fig. 4c) leads to simultaneous decreases in both$ {f}_{\mathrm{s}\mathrm{u}\mathrm{r}\mathrm{f}} $ and$ {f}_{\mathrm{s}\mathrm{w}\mathrm{t}\mathrm{h}\mathrm{r}\mathrm{u}} $ as shown by Fig. 5 and Table 2. As a result, there is a common anomalous upward (negative) atmospheric heat flux (~ –14.0 W m–2 on average), and thus it tends to decrease the SST in each three-equation run. The –14.0 W m–2 anomalous atmospheric heat flux and the 4.8 W m–2 anomaly in the$ {f}_{\mathrm{o}\mathrm{c}\mathrm{n}} $ result in an anomalous upward (negative) net heat flux (~ –9.2 W m–2), explaining the common cooling in the ice edge regions in the sensitivity runs, as shown by Fig. 4d. Further looking at net heat flux and SSTA in each sensitivity run in Table 2, we see that the negative SSTA matches well with the anomalous net upward (negative) heat flux: –0.17°C and –8.8 W m–2 in the CST-DRAG-MU71, –0.13°C and –7.2 W m–2 in the FORM-DRAG-MU71, –0.18°C and –10.6 W m–2 in the CST-DRAG-P07, –0.18°C and –10.1 W m–2 in the FORM-DRAG-P07. On the contrary, in the central Arctic, there is a slight positive SSTA (~0.027°C on average) in three of our sensitivity runs (CST-DRAG-MU71, FORM-DRAG-MU71, and the FORM-DRAG-P07), and there is no clear SSTA (anomaly close to 0°C ) in the CST-DRAG-P07 in this region (Fig. 4d and Table 3). The relatively smaller positive SSTAs in the central Arctic can be explained by the relatively smaller anomalous positive net heat flux among the$ {f}_{\mathrm{o}\mathrm{c}\mathrm{n}} $ ,$ {f}_{\mathrm{s}\mathrm{u}\mathrm{r}\mathrm{f}} $ and$ {f}_{\mathrm{s}\mathrm{w}\mathrm{t}\mathrm{h}\mathrm{r}\mathrm{u}} $ in this region. Compared with the total anomalies (–1.9 W m–2) in the atmospheric heat flux (~ –1.1 W m–2 of the$ {f}_{\mathrm{s}\mathrm{u}\mathrm{r}\mathrm{f}} $ anomalies plus ~ –0.8 W m–2 of the$ {f}_{\mathrm{s}\mathrm{w}\mathrm{t}\mathrm{h}\mathrm{r}\mathrm{u}} $ anomalies, averaged among the three runs of positive SSTA, see Table 3), the anomalous positive$ {f}_{\mathrm{o}\mathrm{c}\mathrm{n}} $ (2.7 W m–2 on average) (Figs. 5g–j) is greater in the three sensitivity runs. Consequently, there is an anomalous positive net heat flux (~ 0.8 W m–2, 2.7 W m–2 vs. –1.9 W m–2), leading to the slight positive SSTA (less than 0.05°C) in the three sensitivity runs.The above results show the influences of heat balance among the
$ {f}_{\mathrm{o}\mathrm{c}\mathrm{n}} $ ,$ {f}_{\mathrm{s}\mathrm{u}\mathrm{r}\mathrm{f}} $ and$ {f}_{\mathrm{s}\mathrm{w}\mathrm{t}\mathrm{h}\mathrm{r}\mathrm{u}} $ on the SSTA and, the influencing magnitude is dependent on the changing in sea ice thickness. From Fig. 4c, the increase in ice thickness is much higher along the ice edge regions than in the central Arctic (0.5 m vs. 0.07 m on average, see Table 2 and Table 3). The 0.5 m increase in sea ice thickness in the ice edge regions leads to a greater reduction in$ {f}_{\mathrm{s}\mathrm{u}\mathrm{r}\mathrm{f}} $ and$ {f}_{\mathrm{s}\mathrm{w}\mathrm{t}\mathrm{h}\mathrm{r}\mathrm{u}} $ than in the increased downward (positive)$ {f}_{\mathrm{o}\mathrm{c}\mathrm{n}} $ (Figs. 5c–f and Table 2), causing the cooling in the ice edge regions, as discussed above. While for the central Arctic (Figs. 5g–j and Table 3), the 0.07 m increase in sea ice thickness is insufficient (Fig. 4c) to cause great decreases in atmospheric heat fluxes, therefore the relatively greater increase in the retained heat flux of the ocean ML ($ {f}_{\mathrm{o}\mathrm{c}\mathrm{n}} $ ) dominates the heat balance among these heat fluxes in the central Arctic. Consequently, the SST is either increased slightly in the CST-DRAG-MU71, FORM-DRAG-MU71, and FORM-DRAG-P07 or comparable in the FORM-DRAG-P07. Finally, the comparison between the ice edge regions and the central Arctic shows that the cooling in the ice edge region is much greater than the warming in the central Arctic Ocean. The degree of the differences is small among the three-equation runs (less than 0.04oC on average); therefore, these will not be addressed further.It should be pointed out that the SSTAs presented here need to be further examined in sea ice-ocean or fully coupled models because the feedbacks between the sea ice and ocean, between the sea ice and atmosphere, and between the ocean ML and deep ocean can influence the heat balance and thus cause SST changes. Nevertheless, the stand-alone model experiments can help us to directly assess the sensitivity of the SST to the three-equation approach.
-
Sections 3.2–3.3 demonstrated how the two-equation and three-equation approaches influence the IO heat exchanges and related basal melt/growth processes in August differently. In this section, we assess the sensitivity of the bottom heat flux, basal melt, and sea ice thickness to different combinations of
$ {\alpha }_{\mathrm{t}} $ ,$ {\alpha }_{\mathrm{s}} $ , and$ \lambda $ used in the three-equation approach (the coefficient combinations are listed in Table 1 and explained in section 2.3).First, by comparing the differences between the form-drag-based and constant
$ {\alpha }_{\mathrm{t}} $ (= 0.006) runs in Fig. 6, we find a general spatially bimodal impact of from-drag-based$ {\alpha }_{\mathrm{t}} $ on the oceanic heat flux (Figs. 6a, d). The upward oceanic heat flux is increased (decreased) in the NGCA (Russian continental shelf) region in each form drag run, which is associated with enhanced (reduced) basal melt (Figs. 6b, e) and decreased (increased) sea ice thickness (Figs. 6c, f). The spatial distributions of these positive/negative differences match the distribution of higher/lower values of the form-drag-based$ {\alpha }_{\mathrm{t}} $ (with respect to 0.006; Fig. 7), indicating that a larger (smaller) form-drag-based$ {\alpha }_{\mathrm{t}} $ increases (decreases) the upward oceanic heat flux and increases (decreases) the basal melt. The impact of the form-drag-based$ {\alpha }_{\mathrm{t}} $ is consistent with the results of Tsamados et al. (2015).Figure 6. The difference of (a) upward oceanic heat flux (
$-{f}_{\mathrm{w}}$ , W m–2), (b) basal melt rate (cm d–1), and (c) sea ice thickness (m) in August between the CST-DRAG-MU71 and FORM-DRAG-MU71. (d)–(f) as (a)–(c) but for CST-DRAG-P07 minus FORM-DRAG-P07.Name Variable Jun Aug Aug-Jun CST-DRAG-MU71 $ -{f}_{\mathrm{w}} $ (W m−2) 5.55 −5.18 −10.73 Meltb(cm d−1) 0.60 0.87 0.27 $ {S}_{\mathrm{i}\mathrm{o}} $ (PSU) 23.73 20.08 −3.65 $ {T}_{\mathrm{i}\mathrm{o}} $ (°C) −1.28 −1.08 0.20 FORM-DRAG-MU71 $ -{f}_{\mathrm{w}} $ (W m−2) 3.36 1.42 −1.94 Meltb (cm d−1) 0.38 0.52 0.14 $ {S}_{\mathrm{i}\mathrm{o}} $ (PSU) 24.44 23.47 −0.96 $ {T}_{\mathrm{i}\mathrm{o}} $ (°C) −1.32 −1.27 0.05 Table 4. The simulated mean values of the upward oceanic heat flux (
$-{f}_{\mathrm{w}})$ , the basal melt rate (Meltb), the interfacial salinity ($ {S}_{\mathrm{i}\mathrm{o}} $ ) and the interfacial temperature ($ {T}_{\mathrm{i}\mathrm{o}} $ ) in June and August, and the changes (August–June) in the CES region in the CST-DRAG-MU71 and FORM-DRAG-MU71.Name Variable Jun Aug Aug-Jun CST-DRAG-P07 $ -{f}_{\mathrm{w}} $ (W m−2) 4.89 −1.11 −6.00 Meltb (cm d−1) 0.43 0.81 0.38 $ {S}_{\mathrm{i}\mathrm{o}} $ (PSU) 25.65 22.30 −3.35 $ {T}_{\mathrm{i}\mathrm{o}} $ (°C) −1.38 −1.20 0.18 FORM-DRAG-P07 −$ {f}_{\mathrm{w}} $ (W m−2) 2.95 0.88 −2.08 Meltb (cm d−1) 0.24 0.48 0.34 $ {S}_{\mathrm{i}\mathrm{o}} $ (PSU) 26.56 24.36 −2.20 $ {T}_{\mathrm{i}\mathrm{o}} $ (°C) −1.43 −1.32 0.12 Table 5. Same as Table 4 but for CST-DRAG-P07 and FORM-DRAG-P07.
Figure 7. Spatial distribution of the form-drag-based heat coefficient
$ {\alpha }_{\mathrm{t}} $ in (a) FORM-DRAG-MU71 and (b) FORM-DRAG-P07. In each plot, the$ 6\times {10}^{-3} $ form drag contour is shown by the dashed black line, and the region (CES) of downward heat flux in the CST-DRAG-MU71 and CST-DRAG-P07 is shown by the red contour line in (a) and (b), respectively.Moreover, compared with the form-drag-based
$ {\alpha }_{\mathrm{t}} $ , a constant$ {\alpha }_{\mathrm{t}} $ (0.006) seems to play an important role in creating the condition for a reversed (downward)$ {f}_{\mathrm{w}} $ (Fig. 8). We found a downward oceanic heat flux in the CST-DRAG-MU71 (Fig. 3f). Figure 8b also shows a similar downward oceanic heat flux in the CES region in the CST-DRAG-P07. In contrast, there are no downward oceanic heat fluxes in the two form-drag-based$ {\alpha }_{\mathrm{t}} $ runs (Figs. 8a, c). To study the influences of the summer IO freshening process beginning in June (Fig. 1) on the reversed oceanic heat flux in the CES region, we calculated the mean changes in this region from June to August. The results are listed in Table 4 (Table 5) for the CST-DRAG-MU71 and FORM-DRAG-MU71 (for CST-DRAG-P07 and FORM-DRAG-P07). First, for the CST-DRAG-MU71, the mean value of the upward oceanic heat flux ($-{f}_{\mathrm{w}}$ ) in June in the CES region is ~5.6 W m–2, this is associated with the ~0.60 cm d–1 of basal melt rate, ~23.7 PSU of interfacial salinity, and ~ –1.28°C of interfacial temperature in this region. As the oceanic heating continues on the sea ice during the summer season (Fig. 1e), the basal melt rate increases by ~0.27 cm d–1 (from 0.60 in June to 0.87 in August, cm d–1). Accordingly, there is a decrease of ~3.6 PSU in the interfacial salinity (from 23.7 PSU in June to 20.1 PSU in August), indicating a significant IO freshening process. Given the reduced interfacial salinity, the interfacial temperature is enhanced by ~0.2°C (from –1.28°C in June to ~ –1.08°C in August). As a result, there is a negative temperature difference (~ –0.2oC; Fig. 3f) between the SST (–1.25°C ) and$ {T}_{\mathrm{f}\mathrm{i}\mathrm{o}} $ (–1.08°C ), which leads to a reversed heat flux (from the ice to the ocean ML) of ~5.2 W m–2 in the CES region in August.Figure 8. The simulated upward oceanic heat flux (
$-{f}_{\mathrm{w}}$ , W m–2) in (a) FORM-DRAG-MU71, (b) CST-DRAG-P07, and (c) FORM-DRAG-P07. Note that the negative values in (b) denote the downward oceanic heat flux.After analyzing the results of CST-DRAG-MU71, we now look at how different changes in the CES region can affect FORM-DRAG-MU71. Our results show that by the smaller form-drag-based
$ {\alpha }_{\mathrm{t}} $ in the CES region (< 0.006; Fig. 7) in FORM-DRAG-MU71, the upward oceanic heat flux in June (3.4 W m–2) in the FORM-DRAG-MU7 is less than that in the CST-DRAG-MU71 run (5.6 W m–2). Accordingly, the basal melt rate is relatively slower (0.38 cm d–1) compared with that of 0.60 cm d–1 in the CST-DRAG-MU71. Associated with the slower basal melt rate, the decrease in the interfacial salinity from June to August (~ –0.1 PSU) is less than that (–3.6 PSU) in the CST-DRAG-MU71 run, indicating a relatively weaker IO freshening process in the FORM-DRAG-MU71 in comparison to the CST-DRAG-MU71. Therefore, although the IO freshening also causes an increase of 0.05°C for the interfacial temperature (from –1.32°C in June to –1.27°C in August), the interfacial temperature –1.27°C is still lower than the SST (~ –1.22°C ), and hence the heat exchange at the IO interface is still from the ocean ML to the ice (upward) in the FORM-DRAG-MU71 run.Similar differences in the CES region can be seen between the runs of a larger constant and smaller form-drag-based
$ {\alpha }_{\mathrm{t}} $ combined with P07 conductivity (CST-DRAG-P07 and FORM-DRAG-P07, Table 5). We can see that there is also a relatively greater (smaller) oceanic heat flux, a relatively greater (smaller) decrease in the interfacial salinity, a relatively greater (lower) interfacial temperature than the SST, and thus a downward (upward) oceanic heat flux associated with a larger (smaller) constant (form-drag-based)$ {\alpha }_{\mathrm{t}} $ in the CES region in the CST-DRAG-P07 (FORM-DRAG-P07).In summary, the above results suggest a coupled mechanism in the three-equation IO boundary approach for a downward heat flux from the ice to the ocean ML in summer. It involves the increased oceanic heat flux since the early period of summer, the continuously strong IO freshening process during JJA, the reduced interfacial salinity, the increased interfacial temperature above the SST, and the final downward heat flux from the ice to the ocean in the later summer, where the parametrization method applied to the
$ {\alpha }_{\mathrm{t}} $ plays an important role in amplifying the magnitude of the oceanic heat flux to the IO interface in early summer. The latter can cause a greater freshening effect at the IO interface.Finally, we describe the impacts of the conductivity on the model simulations. Hunke (2010) demonstrated that the conductivity P07 decreases the e-folding scale of the ice thickness redistribution function, reduces the ocean heating on the sea ice, and adjusts the upward albedo, each of which acts to thicken the ice. In this study, we also observe an increase in the sea ice thickness in our P07 runs compared to our MU71 runs (Fig. 9a). Additionally, we find that the impacts of the P07 conductivity reported by Hunke (2010) can be strengthened by the three-equation boundary condition, which can be demonstrated as follows. In Fig. 1g, we see that the positive heat conduction
$ {f}_{\mathrm{c}\mathrm{o}\mathrm{n}\mathrm{d}\mathrm{b}\mathrm{o}\mathrm{t}} $ is increased in the P07 runs compared with the MU71 runs. According to the three-equation scheme, the increased$ {f}_{\mathrm{c}\mathrm{o}\mathrm{n}\mathrm{d}\mathrm{b}\mathrm{o}\mathrm{t}} $ tends to be associated with a reduced oceanic heat flux to avoid overestimating the heat convergence under a given melt condition (see Eq. 10). Ultimately, this process, controlling the IO energy balance, results in a decreased upward oceanic heat flux (Fig. 9b) and a reduced basal melt (Fig. 9c) in the P07 runs, contributing to the increase in the sea ice thickness (Fig. 9a).
Name | IO boundary | Heat/salt transfer coef. | Ice conductivity |
REF | two-equation | Constant drag, $ {\alpha }_{\mathrm{t}}=0.006 $, $ {\alpha }_{\mathrm{s}}={\alpha }_{\mathrm{t}}/50 $ | Maykut and Untersteiner |
CST-DRAG-MU71 | three-equation | Constant drag, $ {\alpha }_{\mathrm{t}}=0.006 $, $ {\alpha }_{\mathrm{s}}={\alpha }_{\mathrm{t}}/50 $ | Maykut and Untersteiner |
FORM-DRAG-MU71 | three-equation | Form drag (${ {\rm{Cdn} } }\_{\mathrm{o}\mathrm{c}\mathrm{n} }$), ${\alpha }_{\mathrm{t} }={{\rm{Cdn}}}\_{\mathrm{o}\mathrm{c}\mathrm{n} }/2$, $ {\alpha }_{\mathrm{s}}={\alpha }_{\mathrm{t}}/50 $ | Maykut and Untersteiner |
CST-DRAG-P07 | three-equation | Constant drag, $ {\alpha }_{\mathrm{t}}=0.006 $, $ {\alpha }_{\mathrm{s}}={\alpha }_{\mathrm{t}}/50 $ | Pringle et al.(2007) |
FORM-DRAG-P07 | three-equation | Form drag (${ {\rm{Cdn} } }\_{\mathrm{o}\mathrm{c}\mathrm{n} }$), ${\alpha }_{\mathrm{t} }={{\rm{Cdn}}}\_{\mathrm{o}\mathrm{c}\mathrm{n} }/2$, $ {\alpha }_{\mathrm{s}}={\alpha }_{\mathrm{t}}/50 $ | Pringle et al.(2007) |