Convection in the Western Labrador

Although the occurrence of deep convection in the ocean is fairly easy to discern after the fact, it is a difficult process to observe in real time. This is partly because of the harsh wintertime conditions surrounding this phenomenon (high winds, cold temperatures, rough sea state, and often times ice), and also due to the fact that the lateral scales of the convective plumes are very small (Marshall and Schott 1999). It was not until March 1976 that deep convection was directly observed in the Labrador Sea using a shipboard conductivity/temperature/depth (CTD) profiler (Clarke and Gascard 1983). A typical storm that drives overturning in the western Labrador Sea is shown in Fig. 26.2. The storms generally follow the North Atlantic storm track past Newfoundland toward Iceland (Hoskins and Hodges 2002), and the cyclonic circulation draws bitterly cold air off of the Labrador landmass. Total ocean-to-atmosphere heat fluxes from the storms often exceed 500 W m-2, with the largest fluxes occurring near the marginal ice zone (Renfrew and Moore 1999; Pagowski and Moore 2001; Renfrew et al. 2002).

During positive phases of the NAO, the storm track tends to shift to the northeast and the frequency of cyclones increases (Rogers 1990). While this makes it more conducive for overturning to occur in the western Labrador Sea (Dickson et al. 1996), numerous other factors come into play. These include: (1) advection of freshwater from the Arctic, for instance the great salinity anomaly (Dickson et al. 1988) that shut down convection in the early 1970s (Talley and McCartney 1982), as well as other smaller events in the 1990s (Belkin et al. 1998); (2) interannually varying input of warm and salty subtropical water (Curry et al. 1998); and (3) the "memory" of the system (e.g. Straneo and Pickart 2001). Since it takes roughly 5-6

Heat flux due to Labrador Sea storm

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Heat flux due to Labrador Sea storm

W/rrf

Fig. 26.2 Typical Labrador Sea winter storm (16 February 1997) from NCEP. The sea level pressure is contoured, and the vectors are the 10 m winds. The total heat flux (sensible + latent) is in color, where positive flux corresponds to heat loss from the ocean. The flux corrected product of Moore and Renfrew (2002) has been used. The center of the storm is denoted by the L, and the marginal ice zone along the Labrador shelf is colored white. The isobaths are 1,000, 2,000, and 3,000 m (gray lines)

25 mi"1

Fig. 26.2 Typical Labrador Sea winter storm (16 February 1997) from NCEP. The sea level pressure is contoured, and the vectors are the 10 m winds. The total heat flux (sensible + latent) is in color, where positive flux corresponds to heat loss from the ocean. The flux corrected product of Moore and Renfrew (2002) has been used. The center of the storm is denoted by the L, and the marginal ice zone along the Labrador shelf is colored white. The isobaths are 1,000, 2,000, and 3,000 m (gray lines)

years for all of the Labrador Sea Water to be flushed from the basin after formation (Yashayaev 2007), this means that repeated winters of deep convection will prime the system for continued overturning even if a subsequent winter is not very cold or windy. This was the case for the winter of 1996-1997 which produced deep convection with only moderate atmospheric forcing over much of the season (Pickart et al. 2002).

Although the region of strong heat flux from storms such as that in Fig. 26.2 is fairly broad, convection does not readily occur over the entire Labrador basin. This is partly due to the circulation of the sea. As explained in Marshall and Schott (1999), one of the factors, in addition to the atmospheric forcing, that promotes convection is the presence of cyclonic circulation. This both weakens the upper-layer stratification due to the doming of the isopycnals, and traps the water thereby allowing numerous storms to influence the same water parcels. The circulation of the western sub-polar gyre consists of a strong boundary current over the continental slope, and a series of closed cyclonic recirculations adjacent to this (Lavender et al. 2000; Fig. 26.3). It has been argued that these recirculations are driven by the enhanced wintertime windstress curl to the east of Greenland (see below), governed by the dynamics of topographic beta plumes (Spall and Pickart 2003). As seen in Fig. 26.3, it is clear that the deepest mixed-layers in the Labrador Sea occur within

Fig. 26.3 Absolute geostrophic pressure anomaly at 700 m (contours) from Lavender et al. (2000), overlaid on the distribution of surface eddy speed (color) from Lilly et al. (2003). The locations of convection measured by profiling floats in winter 1997 are denoted by the symbols (see legend). The contour interval for the pressure anomaly is 1 cm, and the isobaths are 1,000, 2,000, and 3,000 m. Regions of low geostrophic pressure are indicated by an L

Fig. 26.3 Absolute geostrophic pressure anomaly at 700 m (contours) from Lavender et al. (2000), overlaid on the distribution of surface eddy speed (color) from Lilly et al. (2003). The locations of convection measured by profiling floats in winter 1997 are denoted by the symbols (see legend). The contour interval for the pressure anomaly is 1 cm, and the isobaths are 1,000, 2,000, and 3,000 m. Regions of low geostrophic pressure are indicated by an L

the recirculations. This notion is supported as well by the observations of Clarke and Gascard (1983), Pickart et al. (2002), and Lavender et al. (2002).

A second factor controlling the spatial extent of convection in the Labrador Sea is the eddy field. It is now known that the eastern boundary of the sea, near 61-62° N, is a site of eddy formation, apparently due to both barotropic and baroclinic instability of the boundary current (Eden and Boning 2002; Katsman et al. 2004; Bracco and Pedlosky 2003). One of the main factors is the local variation in topographic slope, which is conducive for instability (Wolfe and Cenedese 2006). The anti-cyclones formed from this region are long-lived and generally translate to the southwest (Prater 2002; Lilly et al. 2003). These features contain warm and salty boundary current water in their cores (hence the name Irminger Rings, Lilly et al. 2003), and they are an effective means of transporting buoyant water into the interior. This is believed to play an important role in the restratification after convection (Katsman et al. 2004), and also seems to influence the location where the deepest convection occurs in the basin. Note in Fig. 26.2 that the heat loss due to the storms is strong in the northern Labrador Sea, within one of the regions of cyclonic circulation (Fig. 26.3). This implies that deep convection should occur there, but observations show that the spreading of buoyant water from the boundary by the Irminger rings inhibits deep overturning (Pickart et al. 2002). This is consistent with the distribution of surface eddy speed (Fig. 26.3) which shows that the northeast part of the basin is strongly influenced by the eddies. The low occurrence of deep mixed-layers in this region (Fig. 26.3) implies that the eddy field helps to confine the deepest overturning to the western part of the basin.

The newly convected Labrador Sea Water leaves the basin by one of three general pathways (Talley and McCartney 1982). The first pathway is in the Deep Western Boundary Current, which is an effective means of transporting the water to the subtropics (e.g. Molinari et al. 1998; Pickart et al. 1997). The second pathway is with the North Atlantic Current, which advects the water to the eastern Atlantic (e.g. Read and Gould 1992). The third pathway is into the Irminger basin. It is this pathway that lies at the heart of the issue of whether or not Labrador Sea Water is formed entirely within the Labrador basin. Based on hydrographic data and models, it has recently been established that the travel time for Labrador Sea Water to reach the Irminger Sea via this pathway is approximately 2 years (Pickart et al. 2003a; Yashayaev et al. 2007; Falina et al. 2007; Kvaleberg et al. 2008). This raises the possibility that past observations of relatively newly convected Labrador Sea Water in the Irminger basin might have been incorrectly interpreted as local formation south and east of Greenland. Judging by the station map of Defant (1936), it is possible that some of the data from the 1935 Meteor cruise that validated Nansen's hypothesis were in fact taken within this pathway from the Labrador Sea. However, this is unlikely the case for all of the stations. Furthermore, based on recently collected hydrographic data, the distribution of Labrador Sea Water within the Irminger basin is inconsistent with a Labrador Sea-only source.

Sy et al. (1997) argued that, in the mid-1990s, Labrador Sea Water took only 6 months to reach the Irminger basin, which is inconsistent with the above estimates and the model results. Furthermore, using early springtime measurements, Pickart et al. (2003a) showed that if this pathway were the sole means by which Labrador Sea Water entered the Irminger basin, then the advective timescale would at times have to be less than 3 months. This is clearly unrealistic, especially since the observations of Pickart et al. (2003a) were taken inside the Irminger gyre. Advective-diffusive models (Straneo et al. 2003; Kvaleberg et al. 2008) show that it takes on the order of 3 years for Labrador Sea Water to penetrate the gyre from the outside. Additionally, observations collected during the most recent positive phase of the NAO during the early 1990s show two separate extrema of weak mid-depth stratification: one in the western Labrador Sea, and one in the southwest Irminger Sea (Fig. 26.4). It is impossible to reproduce such a lateral tracer pattern from a single source of convection in the Labrador Sea (Straneo et al. 2003), and the double-source of Labrador Sea Water considered by Kvaleberg et al. (2008) fits the observations better. Next, we discuss new insights regarding the atmospheric patterns associated with the high orography of Greenland - patterns that are conducive for convective overturning in the vicinity of Cape Farewell.

Potential vorlicity at 1000m (1989-1997)

Potential vorlicity at 1000m (1989-1997)

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