Response of the Arctic freshwater budget to extreme nao forcing




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Response of the Arctic freshwater budget to extreme NAO forcing


Alan Condron*,1,2 , Peter Winsor2, Chris Hill1 and Dimitris Menemenlis3


* Corresponding author, email: acondron@whoi.edu


1 Department of Earth, Atmospheric and Planetary Sciences, Massachusetts Institute of Technology, Cambridge MA 02139 USA.


2 Woods Hole Oceanographic Institute, Physical Oceanography Department, Woods Hole, MA 02543


3 NASA Jet Propulsion Laboratory, California Institute of Technology., M/S 300-323, 4800 Oak Grove Drive, Pasadena, CA 91109


Submitted to Journal of Climate


Abstract


To investigate the response of the Arctic Ocean freshwater budget to changes in the NAO, we use a high resolution (1/6 degree) regional-ocean versionconfiguration of the ocean-only MITgcm and carry out several different 10- and 30-year integrations. At this resolution the model resolves the major Arctic transport pathways, including the Bering Strait and Canadian Archipelago. We perform two main calculations by repeating the wind fields of two contrasting NAO years in each run for the extreme negative and positive NAO phases of 1969 and 1989, respectively. We compare these calculations with a control run and the compiled observationally-based freshwater budget estimate of Serreze [2007].

Our results show a clear response in the Arctic freshwater budget to NAO forcing: Repeat NAO negative wind forcing results in virtually all freshwater being retained in the Arctic, with the bulk of the freshwater content being pooled in the Beaufort Gyre. In contrast, repeat NAO positive forcing accelerates the export of freshwater out of the Arctic to the North Atlantic, primarily via Fram Strait (~900 km3 yr-1) and the Canadian Archipelago (~500 km3 yr-1), with a total loss in freshwater storage of ~13,000 km3 (15%) after 10 years. The large increase in freshwater export through the Canadian Archipelago highlights the important role that this gateway plays in redistributing the freshwater of the Arctic to subpolar seas by providing a direct pathway from the Arctic basin to the Gulf Stream system and the Atlantic Ocean.

We discuss the sensitivity of the Arctic Ocean to long-term fixed extreme NAO states, and show that the freshwater content of the Arctic is able to restore to initial values after only 15-20 years.


1. Introduction

The stability and response of the meridional overturning circulation (MOC) to identical freshwater perturbations and emissions scenarios varies greatly between different models [e.g. Manabe and Stouffer, 1988; Rahmstorf, 1995; Stouffer et al., 2006]. One of the main factors controlling the strength and sensitivity of the MOC is the representation of the freshwater budget of the high latitudes. In the open ocean deep convection regions of the North Atlantic a delicate density balance exists, so that only slight variations in Arctic-North Atlantic freshwater exchanges have an affect effect on the strength of the overturning cell by interrupting the formation of North Atlantic Deep Water (NADW) [Aagaard and Carmack, 1989; Mysak et al., 2005; Rahmstorf et al., 2005; Zhang and Vallis, 2006]. In recognition of the importance of understanding the impact of freshwater exchanges between the Arctic and the Atlantic, the Arctic/Subarctic Ocean Fluxes (ASOF) program was set up in 1999, with the aim of measuring and improving the modeling of these fluxes.

A change in the Arctic freshwater budget between 1965 and 1995 is believed to have increased the export of freshwater to the North Atlantic, as observed by a freshening of the North Atlantic by an amount equivalent to ~19,000 ± 5000 km3 (31536 km3 yr-1 = 1 Sv) of freshwater [Curry and Mauritzen, 2005; Peterson et al., 2006]. About half of this freshwater appears to have been released from the Arctic during the Great Salinity Anomaly (GSA) that began in 1968, when increased ice export in Fram Strait discharged ~2000 km3 yr-1 of freshwater into the northern North Atlantic over a 5-years period [Dickson et al., 1988]. Similar GSA’s occurred in the 1980's and 1990's [Belkin et al., 1998; Belkin, 2004], and all appear to have their origin in the Arctic [Haak et al., 2003; Sundby and Drinkwater, 2007]. Observations of the 1960's GSA indicate that the additional freshwater added to the Labrador Sea inhibited the production of Labrador Sea Water (LSW) [Lazier, 1980], while modeling studies of this event have shown that the reduced LSW formation may have in fact weakened the strength of MOC by 1-3 Sv [Mysak et al., 2005; Zhang and Vallis, 2006]. Considering this rather large change in the MOC, we note that an additional annual release of only 2% of the freshwater stored in the Arctic over a five-year period would generate a signal comparable to that of the GSA.

The changes outlined have been paralleled with changes in the water mass structure and properties of the Arctic Ocean [Quadfasel et al., 1991; Steele and Boyd, 1998; Morison et al., 2000]. Atlantic Water (AW) entering the Arctic via the Barents shelf and the West Spitzbergen Current have warmed since the early 1990s, producing temperature anomalies in the Nansen basin of roughly +1°C in 2004 [Polyakov et al., 2005]. The increased penetration of AW caused the cold halocline (CHL) to retreat from the Amundsen Basin to the Makarov Basin in the mid-1990s [Steele and Boyd, 1998], although measurements showed that it had returned to the Amundsen Basin in 2001 [Bjork et al., 2002].

Large-scale interannual changes in the atmospheric circulation over the Arctic appear to have played a dominatedominant role in the observed changes in the freshwater budget and water mass structure of the Arctic Ocean [Morison et al., 2000; Hakkinen and Proshutinsky, 2004; Karcher et al., 2005; Peterson et al., 2006; Koberle and Gerdes, 2007]. The atmospheric circulation of the Arctic is dominated by the North Atlantic Oscillation(NAO)/Arctic Oscillation (AO) [Hurrell, 1995; Thompson and Wallace, 1998], which switched from its most extreme negative state in the 1960s, to its most extreme prolonged positive state in the early 1990s (Figure 1a)1. Hakkinen and Proshutinsky [2004] and Koberle and Gerdes [2007] both simulated the Arctic freshwater budget from the 1950s to the early 2000s using two different coupled ocean-ice regional Arctic models, forced with atmospheric National Centers for Environmental Prediction (NCEP) reanalysis data [Kalnay et al., 1996]. Both models show a reduction in the freshwater content of the Arctic by ~5000 km3 between the 1980s and the mid-1990s, followed by an increase of similar magnitude by the early 2000s. The change in freshwater storage occurs in parallel with the weakening of the NAO from its persistent high state in the early 1990s, towards a more neutral state after the mid-1990s.

The direct response of the Arctic freshwater budget to the positive and negative phases of the NAO pattern were recently examined by Zhang et al. [2003], and then by Houssais et al. [2007], using regional coupled ocean-ice models of the Arctic, forced with NCEP data. To create atmospheric forcing fields for the NAO, the NCEP data were regressed onto the positive and negative phases of the NAO index over the last 50 years. Interestingly, despite both authors using different models their findings are quite similar. We note, however, that Houssais et al. [2007] primarily focused on the response of the Arctic to positive NAO forcing. During the positive NAO forcing, the freshwater storage in the Arctic reduced due to an increase in the export of sea ice and liquid freshwater in Fram Strait and the Canadian Archipelago.

It is hypothesized that the response of the freshwater budget to the NAO is caused by a switch from anticyclonic to cyclonic circulation over the central Arctic [Hakkinen and Proshutinsky, 2004]. During the negative phase of the NAO, the strong Beaufort high pressure (Figure 1b) results in an anticyclonic atmospheric circulation over the central Arctic, causing freshwater to be stored/retained in the Beaufort Gyre, due to Ekman convergence. In contrast, during the positive phase of the NAO, the Beaufort High weakens, and is replaced by a cyclonic circulation (Figure 1c). This pattern is associated with an intensification of the North Atlantic storm track, and an increased northward atmospheric heat transport. The reduced Ekman convergence can no longer retain the freshwater in the Beaufort Gyre, causing it to be exported into the North Atlantic [Proshutinsky et al., 2002; Zhang et al., 2003; Hakkinen and Proshutinsky, 2004; Peterson et al., 2006]. At the same time, an intensified sea level pressure gradient in the Fram Strait drives sea ice into the North Atlantic [Kwok et al., 2004].

In light of the growing body of literature attributing changes in the Arctic to the NAO, and the sensitivity of the MOC to freshwater release, the main focus of this paper is to investigate changes in the Arctic freshwater budget, and the storage and release of freshwater to the North Atlantic, in response to the positive and negative phases of the NAO. Despite previous work, the different response of the Arctic freshwater budget, in terms of storage, release, and exchanges with the North Atlantic to the negative and positive phases of the NAO, continues to remain unclear. We note that many modeling studies of the Arctic freshwater budget discussed previously suffered from two shortcomings that likely lead to inaccuracies. Firstly, previous model studies restore salinity and/or temperature to climatological values to avoid model drift [Zhang et al., 1998; Maslowski et al., 2000; Zhang et al., 2003; Karcher et al., 2005]. The strength of the restoring typically varies from days to months in these simulations, but any restoring makes deriving an accurate freshwater budget of the Arctic difficult. Secondly, we note that the resolution of the regional models (typically 1º) used to examine the freshwater budget requires a simplification ofneed to simplify the exchange with the northwest Atlantic through the complex network of channels in the Canadian Archipelago. This is typically done by either prescribing the flow using climatological values, or by simplifying the region to one or two large straits [Zhang et al., 1998; Maslowski et al., 2000; Karcher et al., 2005].

We begin this paper with a brief description of the model and its set-up (Section 2), and then evaluate the performance of the model in the Arctic by comparing it to an observed climatological Arctic freshwater budget (Section 3). The response of the Arctic freshwater budget to changes in the NAO is then determined by forcing the model with two extreme positive and negative phases of this oscillation (Section 4). Conclusions are then drawn (Section 5).


2) Model Description and Calculations

a) Model Description

The numerical calculations presented here employ a coupled ocean sea-ice configuration of the MIT General Circulation Model (MITgcm) [Marshall et al., 1997]. The MITgcm configuration used covers a limited area Arctic domain that has open boundaries at ~55° in the Atlantic and Pacific sectors. The exact boundary lines coincide with grid cells in a global cube sphere MITgcm configuration [Menemenlis et al., 2005] (Figure 2). This global configuration is used to provide monthly boundary conditions of potential temperature, salinity, flow, and sea-surface elevation to the simulations presented in this paper.

The grid covering the Arctic domain is locally orthogonal and has a variable horizontal resolution with an average spacing of ~18 km. Although this grid spacing is not eddy resolving (deformation radius in the Arctic is ~5-10 km), the mesh resolves major Arctic Straits, including many of the channels of the Canadian Archipelago. The sea-ice and ocean equations are solved on the same horizontal mesh. The height-based vertical gridding option of MITgcm is utilized. There are 50 vertical levels and vertical spacing is set to vary from ~10 m near the surface to ~450 m at a depth of ~6 km. The vertical resolution is greatest in the upper ocean with 20 vertical levels in the top 300m, which permit a good representation of freshwater in the Arctic halocline. Bathymetry is derived from the U.S National Geophysical Data Center (NGDC) two-minute global relief dataset (ETOPO2), which uses the International Bathymetric Chart of the Arctic Ocean (IBCAO) product for Arctic bathymetry. The ETOPO2 data is smoothed to the coupled model horizontal mesh and mapped to the ocean vertical levels using a ``lopped cell'' strategy [Adcroft et al., 1997], which permits accurate representation of the ocean bottom boundary.

Initial ocean hydrography is taken from the Polar science center Hydrographic

Climatology (PHC) 3.0 database [Steele et al., 2001]. Initial sea-ice distributions is from the Pan-Arctic Ice-Ocean Modeling and Assimilation System datasets [Zhang and Rothrock, 2003]. Atmospheric state (10-m surface wind, 2-m air temperature and humidity, and downward long and short-wave radiation) is taken from the six-hourly National Centers for Environmental Prediction reanalysis [Kalnay et al., 1996]. Monthly mean estuarine fluxes of fresh water are based on the Arctic Runoff database [Lammers et al., 2001; Shiklomanov et al. 2000].

The ocean component is configured to use an equation of state formulated according to Jackett and McDougall [1995]. Ocean surface fluxes (in the absence of sea-ice) are calculated using the bulk formula of Large and Pond [1981]. Boundary layer and convective mixing in the ocean is parameterized according to Large et al. [1994]. Background vertical diffusivity of temperature and salinity is set to 4×10-6 m2s–1. An enhanced vertical diffusivity of 1×10-4 m2s–1 is active at depth motivated by Bryan and Lewis [1979]. Tracer transport equations are solved using a high order monotonicity-preserving scheme [Daru and Tenaud, 2004]. Non-linear momentum terms are solved using a vector invariant formulation [Adcroft et al., 2004] with viscous dissipation following Leith [1968], but modified to dissipate divergence as well as vorticity [Fox-Kemper and Menemenlis, 2008].

The sea-ice component of the coupled system follows the viscous plastic rheology formulation of Hibler [1979] with momentum equations solved implicitly on a C-grid [Arakawa, 1977] using a procedure based on [Zhang and Hibler, 1997]. Fluxes of momentum into ice due to the overlying atmospheric winds and momentum fluxes between sea-ice and the ocean are calculated by solving for the momentum balance at each surface grid column [Hibler and Bryan, 1987]. The ocean sea-ice coupled system is stepped forward synchronously with a time-step of 1800s.

The freezing and melting of sea-ice and associated fluxes of heat and fresh water between the ocean, sea-ice, and atmosphere are calculated by solving a heat balance equation for each surface grid column at each time-step [Zhang et al., 1998; Hibler, 1980; Parkinson and Washington, 1979; Semtner 1976]. When air temperatures are below freezing, precipitation falls as snow. Snow falling on the sea ice is advected with the ice movement, and will gradually be transformed into sea ice as it accumulates. The numerical model calculations were performed using a coupled ocean-sea ice model that was forced by prescribed atmospheric fields. The model configuration is a regional version of the global cube-sphere Estimating the Circulation and Climate of the Ocean, Phase II (ECCO2) model and is based on the hydrostatic, Boussinesq approximated, primitive equation form of the MIT general circulation model [Marshall et al., 1997; Menemenlis et al., 2005] (Figure 2a). The limited area model extends from ~55°N to the pole at the prime meridian, with open boundary conditions at the north, east and west edges of the domain, where the zonal and meridional velocity, potential temperature, and salinity are prescribed (Figure 2b). The boundary conditions were taken from a global integration of the ECCO2 model (as depicted in figure 2a) from 1992 to 2001, and are at a monthly temporal resolution.

The ocean model has an average horizontal resolution of ~18 km (1/6°). Although it is not eddy resolving (deformation radius in the Arctic is ~5-10 km), it does allow for the representation of the major Arctic straits, and in particular the complex network of channels in the Canadian Archipelago. In the vertical, 50 levels increase in thickness from 10 m at the surface to a maximum thickness of 450 m, while a non-linear free surface is used to represent sea surface height. The first 20 model levels are in the top 300 m of the ocean, giving a good representation of freshwater in the Arctic halocline. The model topography is taken from ETOPO2. The model uses the K-profile parameterization (so called kpp) based on Large et al. (1994) for vertical mixing vis. [add values for viscosity, advection, diffusion; Chris and Dimitris]

The ocean model is coupled to a sea ice model, in which ice dynamics follow the viscous plastic model of Hibler (1979), and has two-category zero-layer thermodynamics Hibler (1980) [more; Chris and Dimitris]. The sea ice model is forced by wind stress, internal ice stress, and ocean drag, while heat fluxes over the ice are calculated from standard bulk formula (needs a ref). When air temperatures are below freezing, precipitation falls as snow. Snow falling on the sea ice is advected with the ice movement, and will gradually be transformed into sea ice as it accumulates. The sea ice and ocean model use the same time step, set to 1800 s in the current calculations.

Initial temperature and salinities were taken from the Polar Science Center Hydrographic Climatology (PHC) of the Arctic [Steele et al., 2001] and are made publicly available from http://psc.apl.washington.edu/Climatology.html. Ice concentrations and ice thickness were taken from the Pan-Arctic Ice-Ocean Modeling and Assimilation System Sea ice distribution (PIOMAS), available at http://psc.apl.washington.edu/IDAO/data_piomas.html. Atmospheric forcing data are provided by 6 hourly National Centers for Environmental Prediction (NCEP) reanalysis data [Kalnay et al., 1996].

We use 6 hourly 2-m air temperature, 10-m zonal and meridional wind speed, relative humidity, and precipitation. The latent and sensible heat fluxes are calculated from the bulk aerodynamic flux formula of Large and Pond (1981), and evaporation is calculated by dividing the latent heat flux by the latent heat of vaporization. Freshwater from continental runoff is included from monthly gauged Arctic river data taken from the Arctic runoff database, R-ArcticNet, available from http://www.r-arcticnet.sr.unh.edu. Freshwater is exchanged with the ocean through the melting and growing of sea ice. Sea ice is assumed to be entirely fresh and have a density of 920 kg m-3, while the density of snow is fixed at 330 kg m-3. It is important to note that there is no salinity or temperature restoring in the model, allowing us to more accurately model the Arctic freshwater budget.


b) Model Calculations

We initially perform a control run by forcing the model with 6-hourly NCEP reanalysis data for 10 years from January 1992 to December 2001. To understand the response of the Arctic to the different states of the NAO we then re-ran the model twice from rest, keeping all reanalysis fields exactly the same as the control simulation, except that we repeatedly cycled the 6-hourly wind fields for two contrasting NAO years for 10 years in each integration. We note that due to the tendency for the NAO to persist in one phase on a decadal scale, a 10-year integration of the model in each phase seems appropriate. Furthermore, a significant number of coarser resolution (approx. 1º) global climate models predict a positive increase in the NAO index in the next century in response to rising levels of atmospheric carbon dioxide [Stephenson et al., 2006]. We use the 6-hourly wind velocity data from the two most extreme negative and positive NAO years in the instrumental record of 1969 and 1989, respectively (Figure 1a). The mean sea-level pressure maps show the very different atmospheric circulation patterns in the Arctic for these two years. During the NAO negative year of 1969 (hereafter referred to as NAO-), the Beaufort High pressure is evident over the western Arctic, with a mean sea level pressure over 1020 hPa, resulting in an anticyclonic circulation pattern. In contrast, the NAO positive year of 1989 (hereafter referred to as NAO+) shows an intensified Icelandic low pressure and a cyclonic circulation pattern persisting over the Arctic Ocean. In summer (not shown) the Beaufort High is replaced by a low pressure, giving a cyclonic circulation pattern. We therefore note that for NAO+ years the atmospheric circulation remained largely cyclonic for the entire year.

In order to visualize the export pathways of the stored freshwater from the Arctic to the North Atlantic we add a passive dye to the Arctic in the upper 200 m everywhere the vertical thickness of freshwater exceeds 16 m. In addition, to get a first-order understanding about the system's response we performed a 30-year long NAO+ integration, and a 20-year NAO- integration that starts from the conditions in the Arctic after 10-years of NAO+ forcing. During these experiments, both the 6-hourly NCEP data (excluding wind) and the monthly open boundary conditions, for the period from 1992 to the end of 2001, are cycled 3 times.

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