Warm water pathways, transports, and transformations in the northwestern North Atlantic and their modification by cold air outbreaks

Transcription

1 JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 108, NO. C4, 3129, doi: /2002jc001442, 2003 Warm water pathways, transports, and transformations in the northwestern North Atlantic and their modification by cold air outbreaks Yiyong Luo Graduate School of Oceanography, University of Rhode Island, Narragansett, Rhode Island, USA Huai-Min Zhang National Climatic Data Center, NOAA NESDIS, Asheville, North Carolina, USA Mark D. Prater and Lewis M. Rothstein Graduate School of Oceanography, University of Rhode Island, Narragansett, Rhode Island, USA Received 16 April 2002; revised 30 October 2002; accepted 24 February 2003; published 26 April [1] This paper presents a numerical study of the warm water pathways, transports, and water mass transformation in the Newfoundland Basin region and an investigation of the rectification effects of a series of cold air outbreaks (CAOs) on the above processes. An initial mean state simulation was compared with observations and showed good agreement. Then, repeated CAO events were explicitly included in the surface fluxes and illustrated the following rectification effects. The thermal regime of the entire baroclinic layer was impacted, as the thermocline deepened and the temperature anomaly was noticeable down to 500 m. Different mechanisms were responsible for the propagation of the temperature anomaly at different stages. During the onset of the CAO, vertical diffusion propagated the temperature anomaly downward near the surface, then vertical advection further propagated the anomaly downward between CAO events to about 500 m depth. With CAOs, the North Atlantic Current carried more warm water, and its pathway in the Northwest Corner shifted towards the southeast. The volume-averaged mean and eddy kinetic energy increased by 30% and 20%, respectively, and the water mass transformation rate in the Newfoundland basin was doubled. The Gulf Stream carried more heat, but heat transport to the eastern basin decreased owing to the increase in heat release to the atmosphere. INDEX TERMS: 4255 Oceanography: General: Numerical modeling; 4283 Oceanography: General: Water masses; 4532 Oceanography: Physical: General circulation; 4576 Oceanography: Physical: Western boundary currents; KEYWORDS: Newfoundland Basin, cold air outbreaks, North Atlantic Current, warm water pathways, water mass transformation, heat transport Citation: Luo, Y., H.-M. Zhang, M. D. Prater, and L. M. Rothstein, Warm water pathways, transports, and transformations in the northwestern North Atlantic and their modification by cold air outbreaks, J. Geophys. Res., 108(C4), 3129, doi: /2002jc001442, Introduction [2] The region of Newfoundland Bank and Basin in the northwestern North Atlantic Ocean is referred to as the crossroads of the North Atlantic, for it is here that major branches of the wind-driven and thermohaline circulations flow in close proximity to and interact with each other. The eastward-flowing Gulf Stream (GS) at 40 N turns sharply northward and reattaches to the western boundary as the North Atlantic Current (NAC), while the Labrador Current (LC), flowing southward toward the Tail of the Grand Banks (TGB), partially recirculates to the north along the inshore edge of the NAC. The equatorward flowing Deep Western Boundary Current (DWBC) largely wraps around the Grand Banks and its seaward extension, the Southeast Copyright 2003 by the American Geophysical Union /03/2002JC Newfoundland Ridge (SENR), and subsequently encounters the deep-reaching GS/NAC system (Figure 1). [3] The pathways and interactions of the subpolar and subtropical gyres in this region have major implications for the basin-scale and perhaps global-scale ocean circulation and climate. The GS/NAC is a major source of warm, saline water for the northern North Atlantic Ocean and Nordic Seas. As the GS passes over the SENR it bifurcates into the Azores Current, which continues east as part of the circuit of the subtropical gyre, and a northward branch as the NAC. The NAC generally follows the steep topography of the western boundary in NNE direction from the SENR (at around 40 N, 45 W) to Flemish Cap (47 N, 45 W). The NAC then turns northwest near 49 N, 44 W into the Northwest Corner [Rossby, 1996; Zhang et al., 2001]. At about N the NAC turns sharply toward the east, and the once narrow, well-defined western boundary current 26-1

2 26-2 LUO ET AL.: MODIFICATION BY COLD AIR OUTBREAKS Figure 1. Model domain and major currents of the study region. Thick dashed lines indicate the boundaries across which transports are calculated and discussed in the text. Legends are GS, Gulf Stream; NAC, North Atlantic Current; LC, Labrador Current; DWBC, Deep Western Boundary Current; GB, Grand Banks; TGB, Tail of the Grand Banks; SENR, Southeast Newfoundland Ridge; FC, Flemish Cap; NC, Northwest Corner; CGFZ, Charlie-Gibbs Fracture Zone; and MAR, Mid-Atlantic Ridge. The bathymetry shown are 200 m, 500 m, and 1000 to 5000 m by 1000 m increments. becomes a weaker, more diffusive eastward flow known as the Subpolar Front (SPF), which separates the cold, fresh subpolar waters from the warm and saline subtropical waters [Carr and Rossby, 2001]. It is along and across the SPF that the warm subtropical waters make their ways into the subpolar region. The flow of the warm, subtropical waters into the subpolar region and Nordic Seas is an important part of the meridional overturning circulation (MOC) of the Atlantic Ocean. These waters lose large amounts of heat to the atmosphere (which moderates the climate of northern Europe) and are transformed into the intermediate and deep-water masses that spread southward by the LC and DWBC as part of the thermohaline circulation. [4] One of the most striking weather features over the northwestern North Atlantic Ocean is the frequent occurrence of winter-time, synoptic-scale cold air outbreaks (CAOs), which move eastward off the North American continent and can persist for several days. When a CAO moves over the ocean surface, the ocean not only receives a large momentum flux from the atmosphere but also releases large amounts of heat and moisture to the atmosphere as a result of the increased surface stresses and large sea-air temperature differences. Simpson [1969] suggested that significant heat, moisture, and momentum fluxes are concentrated almost entirely within the synoptic-scale forcing events. Results by Elsberry and Camp [1978] showed that over two-thirds of the total wind energy input at three Ocean Weather Stations in the North Pacific Ocean was by synoptic forcing events, which occurred roughly one-third of the year. The presence (or absence) of just a few of these strong events might have significant effects on the seasonal variability. Budyko [1974] showed that ocean-to-atmosphere latent plus sensible heat fluxes for an average December reached a global maximum along the GS off the east coast of the United States. Observations indicate that the combined (latent plus sensible) ocean-to-atmosphere heat flux can reach 1350 W m 2 during an intense CAO [Grossman and Betts, 1990; Wayland and Raman, 1989; Xue et al., 1995]. [5] During a CAO the strong exchanges of heat, moisture and momentum affect both the ocean and atmosphere. In the atmosphere these effects include the generation or intensification of extratropical cyclones [Newton and Holopainen, 1990], coastal frontogenesis [Riordan, 1990; Huang and Raman, 1992, Xue et al., 2000], and enhanced precipitation concentrated in rainbands [Hobbs, 1987; Huang and Raman, 1992]. Oceanic responses can be divided into local and nonlocal components. The local response includes surface temperature decreases and deepening of the surface mixed layer, both as a result of locally intense one-dimension processes [Elsberry and Camp, 1978; Bane and Osgood, 1989; Xue et al., 1995; Xue and Bane, 1997]. The nonlocal response is composed of thermal advection, heat flux at the base of the surface mixed layer, and modification of the local currents by buoyancy gradient effects [Worthington, 1976, 1977; Adamec and Elsberry, 1985a, 1985b; Nof, 1983; Xue et al., 1995; Xue and Bane, 1997]. [6] The above studies focused on the comparatively short-term (of the order of weeks) oceanic response to a single CAO event. On the other hand, previous studies inferred that a series of such events might influence the thermal regime of the entire baroclinic layer (upper ocean) and lead to long-term climatic consequences [Frankignoul and Hasselmann, 1977; Frankignoul and Reynolds, 1983; Polonsky et al., 1992]. The model simulations by Polonsky et al. [1992] showed that storm-induced vertical advection in the main thermocline exceeded vertical diffusion by one and a half orders of magnitude, which considerably influenced the thermal regime of the main thermocline if there were several consecutive storms of the same sign and similar trajectory. Using stochastic models, the results of Frankignoul and Hasselmann [1977] inferred that largescale, long-time sea surface temperature anomalies might be explained naturally as the response of the oceanic surface layers to short timescale atmospheric forcing; the white noise spectrum of synoptic-scale atmospheric fluxes produced a red response spectrum, with most of the variance concentrated in very long periods. Frankignoul and Reynolds [1983] also demonstrated that the atmospheric forcing acted as a white noise; in regions of strong currents the advection effects were important at low frequencies. [7] For this study we used a regional ocean general circulation model to quantity the modification of the warm water transports and pathways in the northwestern North Atlantic Ocean when a series of realistic synoptic-scale CAO events were reintroduced into the wintertime climatological fluxes. Descriptions of the model and configuration are given in section 2. Results from the mean state run are presented in section 3. Section 4 investigates the oceanic response to the CAO events. Section 5 discusses the water

3 LUO ET AL.: MODIFICATION BY COLD AIR OUTBREAKS 26-3 mass transformation and heat budget, and a summary is given in section Model Description [8] The Princeton Ocean Model (POM) [Blumberg and Mellor, 1987; Mellor, 1996] was configured specifically for our study region. Rowley [1996] used the POM in the Newfoundland Basin to study the structures and topographic control of the NAC. He used relatively coarse resolution (11 levels vertically and 1/8 by 1/8 horizontally), but the model was capable of simulating the basic structure, pathways, and meanders of the NAC as compared with those observed by the RAFOS floats [Zhang et al., 2001]. The POM has a free surface, bottom-following vertical sigma coordinate, and a turbulence closure submodel [Mellor and Yamada, 1982]. Our model domain extended from 35 N to55 N and 27 W to55 W (Figure 1) with an eddy-resolving grid. The horizontal grid was chosen to give better resolution over the region where strong jets and the largest velocity shears were expected and varied from 7 km to 9 km over the GS/NAC and expanded to 17 km along the boundaries. The sigma coordinate had 16 vertical layers; the vertical resolution was higher near the surface and lower near the bottom. (For instance, where the water depth is 3000 m, the layer thickness was 5 m at the surface and about 300 m near the bottom.) The horizontal time differencing was explicit, whereas the vertical differencing was implicit. The latter eliminated instability time constraints for the vertical grid and permitted the use of finer vertical resolution near the surface and in shallow water regions. The topography (Figure 1) was derived from Smith and Sandwell [1997] 1/30 data, and the maximum bottom slope allowed between two adjacent grid points was H/H < 0.5 to reduce the so-called pressure gradient errors [Ezer and Mellor, 1997]. Further details on the model configuration are given in the following section. 3. Mean Oceanic State [9] Recently, Smith et al. [2000] presented a basin-scale eddy resolving model simulation for the North Atlantic Basin (20 S 72.6 N). Their comparison with observations showed that the model performed well for the large-scale circulation, especially in the region of our interest [Smith et al., 2000]. Our strategy was to use these basin-scale model results as initial and boundary conditions for the regional model simulations Model Setting [10] The Smith et al. [2000] model results used in this study consisted of 10-day snapshots of forcing (wind stress and surface heat flux) and prognostic variables (surface elevation, velocity, temperature, and salinity) from 1 January 1992 to 31 December This time period included the years ( ) during which 100 RAFOS floats in this region had collected Lagrangian trajectory data on the 27.2 and 27.5 s t surface [Rossby, 1996; Zhang et al., 2001]. A 4-year mean of each field was then obtained by averaging these snapshots. The mean wind stress and heat flux were used as surface forcing, and the mean model results as initial and boundary conditions for our regional model. Smith et al. [2000] used the daily wind stress and the climatological seasonal surface heat flux, which were averaged from the 6-hourly ECMWF Analysis. [11] For the regional POM simulations, Sommerfeld-type radiation conditions were used for velocities normal to the open boundaries, and an upwind advective scheme was applied to the temperature and salinity as well as the velocity components tangential to the boundaries. For the mean state run, temperature and salinity were restored to the means with a relaxation time scale varying linearly from 5 to 40 days over eight-gridpoint-wide buffer zones at the open boundaries. A Laplacian smoothing-desmoothing technique was applied to the surface elevation, velocity fields, temperature, and salinity to reduce small-scale noise, particularly near the open boundaries [Rowley, 1996]. [12] The model was initialized with the ocean at rest and the temperature and salinity set to the 4-year means. To reduce the sigma-coordinate pressure gradient errors, the mean density field was subtracted before integrating the density field to obtain pressure from the hydrostatic equation [Beckmann and Haidvogel, 1993; McCalpin, 1994; Mellor et al., 1998]. The time step was 10 s for the external mode and 800 s for the internal mode. The model was then run for 10 years under the mean wind stress and surface heat flux. The domain-averaged kinetic energy reached an equilibrium value after three years of simulation, which is a typical timescale for baroclinic adjustment of the velocity to the initial density field. The results discussed below were the means taken from the final 4 years of the simulation Mean Features [13] We begin with an overview of the model s regional mean circulation in terms of mean flow, features of the 27.5 s t isopycnal surface, kinetic energy, and the warm water transports and pathways. Whenever possible, the simulations are compared with observations Circulation and Transport [14] The sea surface elevation and surface velocity field are shown in Figures 2a and 2b, respectively. The salient features of the mean current were in remarkably good agreement with those from observations by surface drifters and subsurface RAFOS floats (Figure 2d), also described by Rossby [1996], Kearns and Rossby [1998], Carr and Rossby [2001], and Zhang et al. [2001]. Southeast of the Grand Banks near (40 N, 45 W), the GS split into the NAC (flowing northeastward) and a broader flow to the south. As the NAC flowed northeastward along the continental slope, it exhibited three major meanders in its mean path with troughs at 41 N, 44 N, and 47 N, which agreed with the mean currents derived from RAFOS float observations (Figure 2d). Beyond Flemish Cap, the NAC pathway turned into the Northwest Corner where a series of four anticyclonic quasi-stationary rings appeared on its east side at roughly (47 N, 39 W), (49 N, 41 W), (51 N, 44 W), and (52 N, 46 W). All these anticyclonic rings also appeared in the float observations (Figure 2d). The mean current retroflected sharply toward the southeast and then proceeded eastward between 48 N and 51 N, and when it reached 37 W, it fanned out into a broader region as a more diffusive flow.

4 26-4 LUO ET AL.: MODIFICATION BY COLD AIR OUTBREAKS Figure 2. Results from the mean experiment: (a) sea surface elevation with contour interval of 10 cm; (b) surface velocity; (c) barotropic stream function with contour interval of 5 Sv; and (d) Eulerian mean velocity on 27.5 s t surface from RAFOS float observations. [15] In the following discussions the four boundaries for volume transport are the thick dashed lines in Figure 1. Specifically, the transport across the eastern boundary is actually the transport across the curved dash line along the MAR, since this is more physically meaningful, representing cross-basin transport at this section. (Note that the model domain eastern boundary actually extends to 27 W.) [16] The volume transport stream function of the whole water column, which reflects the mean barotropic transport, is shown in Figure 2c. Of the 90 Sv entering from the western boundary near 40 N mainly as the GS, about 55 Sv left from the western and southern boundaries in the recirculation of the Stream. North of the GS, about 15 Sv flowed west over the continental slope at the Tail of Grand Banks. Roughly 10 Sv of this exited from the western boundary between 42 N and 44 N, and the other 5 Sv returned northward joining the remaining 35 Sv of the GS transport to form the NAC. The NAC lost about 10 of its 40 Sv to the Mann Eddy around (44 N, 41 W) and another 25 Sv was lost to east of Flemish Cap which looped around

5 LUO ET AL.: MODIFICATION BY COLD AIR OUTBREAKS 26-5 Figure 3. Temperature at 45 N in the mean experiment. (49 N, 41 W) and then flowed eastward forming the southern branch of the SPF. The last 5 Sv, augmented with another 5 Sv of LC transport which retroflected to north (at south of Flemish Cap), continued north of Flemish Cap and then retroflected in the Northwest Corner and eventually flowed eastward forming the northern branch of the SPF. About 25 Sv flowed over the Mid-Atlantic Ridge (MAR) entering the eastern basin, of which 15 Sv crossed the MAR near the Charlie-Gibbs Fracture Zone (CGFZ) (Figure 2c). These transport features were in reasonable agreement with observations as follows. Krauss [1986] reviewed the results from drifters launched in the vicinity of the NAC, the Azores Current, and the SPF and concluded that the NAC lost 10 of its 35 Sv around the Mann Eddy and that the rest (25 Sv) continued northward (Figure 19 of Krauss [1986]). Results of two hydrographical surveys of the area southeast of Newfoundland (May June 1990), where the GS branches into the NAC and the Azores Current, showed that the geostrophic volume transports of the NAC were 35 Sv and 46 Sv, respectively [Koshlyakov and Sazhina, 1994]. Bubnov [1995] applied the geostrophic method to hydrographic sections taken along 36 W between 47 N and 53 N in April June 1990 and calculated m transports of 12.4 Sv and 14 Sv for the northern and southern branches of the SPF using a 2000 m reference level. According to a hydrographic survey conducted in the Newfoundland Basin during January and February 1997, Ganiaux et al. [2001] identified three fronts (the Northern Subarctic Front, the Southern Subarctic Front, and the Middle-Atlantic Front) along 35 W between 45 N and 52 N. The currents associated with these fronts transported 26 Sv toward the east before crossing the Mid-Atlantic Ridge and supplying the eastern part of the North Atlantic Basin Warm Water Transports and Pathways [17] In this paper, warm water is defined as the thermocline waters warmer than 7 C. The model temperature at 45 N is shown in Figure 3. The 7 C isotherm approximately corresponded to the bottom of the thermocline, which was deeper than 1000 m in the southeastern domain and shallower than 500 m in the northwestern domain. [18] The warm water originated from the western boundary current as the upper layer water of the GS (Figure 4). A large fraction (22 out of 36 Sv) passed over the MAR and entered the eastern basin, of which 12 Sv flowed over the MAR near the CGFZ. The remaining 14 Sv exited from the southern (10 Sv) and northern boundaries (4 Sv) west of the MAR. Of the 36 Sv of warm water carried by the GS at 50 W, 10 Sv were lost around the southeast of the Grand Banks to the southern boundary. The rest (26 Sv) flowed toward the north as the NAC. Of the 26 Sv NAC warm water, about 10 Sv left the NAC/western boundary to the Mann Eddy and flowed into the Newfoundland Basin; the rest (16 Sv) continued flowing to the northeast along the steep topography. About 12 Sv of this turned to the east at 47 N and was transported by the southern branch of the SPF, while the other 4 Sv turned northwest and retroflected to the east at (51 N, 44 W) and was transported by the northern branch of the SPF. Along 37 W between 42 N and 52 N, all of the 26 Sv warm water flowed northeastward. Before crossing the MAR, about 4 Sv turned northwestward and exited from the northern boundary, and the rest (22 Sv) crossed the MAR and entered the eastern basin Features of the 27.5 s t Isopycnal Surface [19] The model depth of the 27.5 s t isopycnal surface is shown in Figure 5. The dynamic structure of the surface was similar to that obtained from RAFOS floats deployed in the Newfoundland Basin (Figure 4b of Carr and Rossby [2001] and Figure 6b of Zhang et al. [2001]). The depth of the isopycnal underwent large changes across the NAC and SPF, and varied from 200 m on the onshore/subpolar side to m near the 4000 m isobath and to 700 Figure 4. A warm water transport schematic (in Sverdrups) in the mean experiment superimposed on the vertically averaged velocity of the warm water (>7 C).

6 26-6 LUO ET AL.: MODIFICATION BY COLD AIR OUTBREAKS 600 cm 2 /s 2 due to the low variability of that feature [Carr and Rossby, 2001]. [21] Note that in the regions of highest EKE, such as the NAC, the level of EKE was comparable to or even greater than the MKE due to the growth of meanders and formation of eddies. Thus although peak level of MKE exceeded those of EKE, the volume-averaged MKE and EKE over the entire domain were 7.9 and 9.4 cm 2 /s 2, respectively, so the EKE contributed about 54% to the total energy. 4. Effects of Cold Air Outbreak [22] The Barnier et al. [1995] heat flux, a seasonal climatology derived from 6-hour ECMWF data, and daily averages of ECMWF wind stress were used as surface Figure 5. Depth of the 27.5 s t isopycnal surface (solid lines) superimposed on bathymetry lines of 200 m, 2000 m, and 4000 m (dashed lines). 800 m on the offshore/subtropical side. The depth contours largely reflected the upper layer current patterns, such as the trough locations, the NAC, and the broadening of the SPF Kinetic Energy [20] The near-surface ( m) mean kinetic energy (MKE) and eddy kinetic energy (EKE) are shown in Figure 6. The statistics were calculated from the 10-day snapshots of the last 4 years of the simulations. Large values of the MKE were constrained to the main axes of the NAC and LC (Figure 6a). A band of high MKE values (above 600 cm 2 /s 2 ) occurred along the NAC axis between the SENR and Flemish Cap except around (45 N, 45 W) offshore of the 4000 m isobath, and values larger than 1600 cm 2 /s 2 were found along the 4000 m isobath from 46 N to47 N. These features were in good agreement with observations from surface drifters [Carr and Rossby, 2001] that showed MKE value in the NAC of more than 400 cm 2 /s 2 with the maximum of 1800 cm 2 /s 2 around (46 N, 43 W). Comparing the EKE with the MKE, it is clear that maximum EKE values were concentrated in the vicinity of the major currents, except near stable, topographically controlled currents such as the LC (Figure 6b). The pattern of EKE was essentially the same as those found in previous studies [Richardson, 1983; Krauss and Käse, 1984; Horne and Petrie, 1988; Brügge, 1995; Carr and Rossby, 2001; Stammer, 1997; Heywood et al., 1994; Zhang et al., 2001]. A band of high EKE values was wider at 41 N, 44 N, and 47 N, consistent with the widening NAC path and pinching off of eddies at these locations [Rossby, 1996; Kearns and Rossby, 1998]. Maximum EKE values around SENR were above 900 cm 2 /s 2 due to a combination of strong flow and high variability, and EKE values in Mann Eddy were below Figure 6. Results from the mean experiment: (a) mean kinetic energy and (b) eddy kinetic energy averaged between 50 and 100 m depths.

7 LUO ET AL.: MODIFICATION BY COLD AIR OUTBREAKS 26-7 reasonable T/S values at the open boundaries, a model simulation was performed by replacing the Smith et al. [2000] T/S values by the Levitus [1982] climatology at the open boundary conditions in the CAO experiment. No significant difference in the results was found between the two boundary cases Thermal Modification and Mechanisms [24] The most obvious and expected response was the decrease of the near-surface temperature, which was largely due to the huge oceanic heat loss associated with the CAO events. On average, SST decreased by 1.8 C. An unexpected effect was that the thermal regime of the entire baroclinic layer was influenced, i.e., the thermocline lowered and the temperature anomaly reached down to 500 m, which is approximately the 11 C surface in the southeastern domain (Figures 8a and 8b). This was noticeably different from the oceanic response to a single CAO event, where only the upper mixed layer (generally 10 to 100 m) Figure 7. Positions of the largest heat flux of the chosen CAO every 6 hours from February 1993 to February 1993 and associated heat flux (W m 2 ) and wind stress (Pa) at February 1993 from ECMWFAnalysis. forcing by Smith et al. [2000]. In the mean oceanic state simulation of section 3, the forcing data were obtained by averaging the 10-day snapshots of Smith et al. [2000] which filtered out high-frequency forcing events, such as the CAO. In this section, an experiment was designed to investigate the magnitude and characteristics of the modification to warm water transports and pathways in the model region when a series of synoptic-scale CAO events were explicitly expressed in the surface forcing of fluxes during the winter season. Based upon the ECMWF analysis, a representative example of the CAO was the event between 9 12 February 1993, during which the total heat flux (latent plus sensible) exceeded 1200 W m 2 over the NAC as it moved off the continent and toward northeast. The position of the largest heat flux of that CAO every 6 hours, and the associated heat flux and wind stress midway through the event from ECMWF analysis is showed in Figure 7. Beginning with the model circulation from the 10-year spin-up (the mean oceanic state described in section 3), the model was then run for another 3 years using the same mean forcing except for periods when the representative CAO s surface wind stress and heat flux were imposed. The CAO fields were imposed every 10 days for 4 months in order to represent a series of 12 CAO events per winter season. The domain-averaged kinetic energy of the simulation indicated that an equilibrium state had been reached after 1 year of the 3-year experiment, the following analysis was from time-averaged results over the last 2 years of the simulation. [23] During the CAO experiment the buffer zone technique was removed from the open boundaries and only the upwind advection scheme was applied to temperature and salinity (T/S). To demonstrate that the results of the CAO experiment were not critically sensitive to different but Figure 8. Section contours at 45 N: (a) temperature in the CAO experiment and (b) temperature anomalies: CAOmean.

8 26-8 LUO ET AL.: MODIFICATION BY COLD AIR OUTBREAKS Figure 9. Terms (row-wise) in the heat conservation equation at depths of 20 m, 200 m, and 400 m (column-wise): (a) during an active CAO; (b) between intermission of two CAO events; and (c) in the summer time free of CAO events. All terms except temporal variation have been moved to the righthand side of the equations. deepened and the temperature of the upper mixed layer decreased but the thermal structure of the main thermocline was not influenced [Xue et al., 1995; Xue and Bane, 1997]. [25] To understand the mechanism of the CAO-induced, deep-reaching cooling, all the terms in the governing equation of temperature conservation were calculated and examined at three depths (near the surface (20 m), and at 200 and 400 m depths) and during three separate times of forcing regimes (at an active CAO, during the intermission between two CAO events in the winter season, and during the CAO-free time period of the year). The results are

9 LUO ET AL.: MODIFICATION BY COLD AIR OUTBREAKS 26-9 Figure 9. (continued) shown in Figures 9a, 9b, and 9c for the three time stages, respectively. During an active CAO (Figure 9a) and near the surface (top left panel), most of the area was being cooled. At 200 m (top middle panel) about half of the area was being cooled and the rest was being warmed. At 400 m (top right panel), much of the area was still being warmed. The sequence indicated that the cooling was progressively downward but still had not been reached to 400 m depth at this time. In fact, during the intermission between two CAO events and at 400 m depth (Figure 9b), the cooling area was larger than that at the active CAO. [26] To illustrate the mechanism of cooling and warming, the dominant terms in the thermal (temperature) equation were examined, i.e., column wise in Figure 9. During an active CAO and in the surface layer at 20 m (Figure 9a, left column), the large-area cooling (top) was mainly balanced by the vertical diffusion term, which included the air-sea flux term at the sea surface. However, the warming was

10 26-10 LUO ET AL.: MODIFICATION BY COLD AIR OUTBREAKS Figure 9. (continued) mainly balanced by the horizontal advection. At 200 m and 400 m the temporal changes in temperature were mainly caused by vertical advection, with contribution from horizontal advection. At the intermission between two CAO events in the wintertime (Figure 9b), the role of vertical diffusion (including air-sea flux) was reduced to secondary even in the surface layer (left panels); the temperature changes were mainly brought by horizontal advection. However, at 200 and 400 m, although horizontal advection still played an important role, the vertical advection dominated the changes in temperature. During the CAO-free time period of the year (Figure 9c), the dominant term was horizontal advection (with contribution from horizontal diffusion) in the surface layer and at 200 m; both horizontal and vertical advection dominate at 400 m. [27] In summary, during the onset of the CAO, vertical diffusion propagated the temperature anomaly downward in the upper layer, then vertical advection propagated the

11 LUO ET AL.: MODIFICATION BY COLD AIR OUTBREAKS evident in the three meanders in the NAC s path and the three anticyclonic rings in the Northwest Corner. However, significant modifications were seen in other aspects. For example, the anticyclonic ring at (52 N, 46 W) disappeared, which was consistent with most observations, although some of the float trajectories showed the presence of such a feature (Figure 2d). Furthermore, the width of the GS and NAC in the south narrowed and their strength increased, which indicated that stronger flows occurred along narrower pathways. Figure 10. Results from the CAO experiment: (a) sea surface elevation with contour interval of 10 cm and (b) surface velocity. anomaly further downward between CAO events to about 500 m depth. During the summer, the anomaly was reduced by the horizontal advection of subtropical warm waters into the domain Surface Circulation [28] The sea surface elevation and surface velocity field in the CAO experiment are shown in Figures 10a and 10b. Compared with the results shown in Figures 2a and 2b, it is apparent that the fundamental features of the circulation were not affected; only slight changes were 4.3. Kinetic Energy [29] The strong atmospheric forcing in the CAO experiment produced stronger flows and increased eddy activity and thus increased both the MKE and EKE. The volumeaveraged MKE and EKE increased by 30% (from 7.9 to 10.2 cm 2 /s 2 ) and 20% (from 9.4 to 11.3 cm 2 /s 2 ), respectively. In the surface layer a band of large MKE values were about 200 cm 2 /s 2 higher (a total of 800 cm 2 /s 2 )inthe CAO experiment along the NAC axis between SENR and Flemish Cap except around (45 N, 45 W), and the maximum values exceeded 1900 cm 2 /s 2 (300 cm 2 /s 2 higher than in the mean experiment) along the 4000 m isobath from 46 N to47 N (Figure 11a). The MKE values in the LC were increased in the CAO experiment as well. Compared to the mean state (Figure 6b), the CAO experiment resulted in a band of large EKE values along the NAC in the south that were even higher (Figure 11b). Meanwhile, away from the strong flows, the EKE values were usually larger in the CAO experiment, meaning that eddy activity was more active there than in the mean experiment Warm Water Transports and Pathways [30] The warm water transports and pathways in the CAO experiment are shown in Figure 12 and, when compared with the results from the mean state (Figure 4), indicated significant modifications. Four Sverdrups more warm water (a total of 40 Sv) entered from the western boundary transported by the GS, and eventually, half of this (2 Sv) left from the southern boundary (a total of 12 Sv) and another half (2 Sv) left from the northern boundary (a total of 8 Sv). Instead of crossing the MAR near the CGFZ and entering the eastern basin, the additional two Sverdrups of water exited from the northern boundary. The NAC carried an additional 2 Sv of warm water (a total of 28 Sv) along its entire pathway. The warm water to the Mann-Eddy was not affected and was still 10 Sv in the CAO experiment. However, the warm water carried by the southern branch of the SPF was 2 Sv less (a total of 10 Sv). Along 37 W between 42 N and 52 N, two Sverdrups more warm water (a total of 28 Sv) flowed toward the northeast. On the other hand, in addition to transporting 4 Sv more warm water (a total of 8 Sv), in the CAO experiment the pathway of the northern branch of the SPF moved slightly towards the southeast, which was due to some warm water being transformed to cold water (<7 C) there. [31] In short, the GS carried more warm water that entered from the western boundary, and part of this flowed southeast and left from the southern boundary, with the rest flowing along the NAC toward the northeast and eventually leaving the northern boundary. Instead of crossing the MAR

12 26-12 LUO ET AL.: MODIFICATION BY COLD AIR OUTBREAKS The depth ranges of the three water masses and associated volume transports and water mass conversions (in terms of volume transport) are shown in Table 1 and Figure 13a. Note that the depth ranges were values from the mean experiment, and those from the CAO experiment were comparable. The SLW had a layer thickness of 250 m in the west, 400 m in the south, 200 m in the east, and zero (outcropping) in the north. For the mean experiment, 19 Sv of SLW entered the domain at the west boundary (part of the GS), and 9 and 6 Sv exited from the south and east boundaries, respectively. This resulted in 4 Sv of the warm water being transformed to cooler UTW through the 14 C isotherm. The CAO events drew 2 Sv more SLW at the west boundary and exported 1 Sv more at the south but 2 Sv less at the east, which resulted in a net 3 Sv increase in the warm-to-cooler water mass conversion. [33] The UTW also outcropped to the north, but the layer thickness at the west, south, and east were about 100, 200, and 300 m, respectively. In the mean experiment, 9 Sv of this UTW entered from the west and 1 and 8 Sv exited from the south and east, respectively. There were 4 Sv input through the upper surface (14 C isotherm) and nearly the same amount (4 Sv) output to the lower LTW through the 11 C isotherm. The nearly same water mass transformation rates at the upper (14 C) and lower (11 C) surfaces indicated that this layer was almost transparent to the water mass conversion, thus acting as a free passage of water mass conversion from the SLW to the LTW. For this water mass class, the CAO events brought 1 Sv more water through the west boundary, exported 1 Sv more through the south but 2 Sv less through the east, which resulted in a net water mass conversion from UTW to LTW of 9 Sv (5 Sv more than through the upper surface of 11 C isotherm). [34] The LTW layer was still quite shallow along the northern boundary (bottom depths of about 450 m) but quite deep along the southern boundary (bottom depths of about 1200 m). Therefore the flow in the south in this layer was Figure 11. experiment. Same as Figure 6 but from the CAO and entering the eastern basin, the additional two-sverdrup warm water exited from the northern boundary. 5. Discussions 5.1. Water Mass Transformation [32] Combining the transport and cooling propagation calculations, the following summary cartoons were produced to show the three-dimensional water mass transformation in this region [e.g., Speer and Tziperman, 1992; Zhang and Talley, 1998]. In the following we describe the warm-to-cold water mass conversion rates by temperature class and show how they were affected by the CAO events. Three water masses were defined for the upper ocean of this region: the surface layer water (SLW) with T > 14 C, the upper thermocline water (UTW) with 11 C < T < 14 C, and the lower thermocline water (LTW) with 7 C < T < 11 C. Figure 12. experiment. Same as Figure 4 but from the CAO

13 LUO ET AL.: MODIFICATION BY COLD AIR OUTBREAKS Table 1. Water Volume Transports and Balances (in Sv) for the Three Layers of the Warm Water Masses Defined in the Text a a Negative values indicate transports out of the boxes and vice versa. Values in parentheses are approximate depths of the bottoms of the three layers. Values in the Total Warm Water row are the vertical sum, i.e., the values for the whole warm water column between sea surface and 7 C isotherm. very weak, and the associated volume transport was very small (0 Sv) compared with those along other boundaries. For the mean experiment, the GS carried in 8 Sv water, and the same amount exited from the east boundary. The upper layer (UTW) transformed 4 Sv water to this layer through the interface (11 C isotherm), and the same amount exited from the north boundary instead of through the lower surface (7 C isotherm) of the layer, which was very small (0 Sv) compared to other surfaces. The CAO events did not produce any significant transport across the 7 C isotherm either but did affect the transports and mass transformation within this layer. For instance, with the CAO events, this layer exported 4 and 2 Sv more water to the north and east, respectively, of which 5 Sv more was transformed from the upper layer (UTW) and 1 Sv more entered from the GS. [35] It is important to notice that in Figure 13a, the 4 Sv (9 Sv with CAO events) LTW was not further transformed to the colder than 7 C water in this region (the bottom transport through 7 C isotherm was near zero). Instead, the LTW eventually turned north and exited the north boundary (with 4 Sv and 8 Sv with CAO events). In other words, in the Newfoundland Basin, the strong atmospheric forcing transformed the SLW and UTW into the LTW, and then the LTW was transported to the north of this region but not transformed to the deep cold water (no Deep North Atlantic Water or DNAW formation there). It is also shown that the modification of the CAO events doubled the water mass transformation rates in this region. With CAO events, of the 9 Sv transformed to the LTW, 8 Sv exited from the north and 1 Sv from the east Heat Budget [36] Along with the warm water, a large amount of heat is transported to the northwestern North Atlantic Ocean and partly to the Nordic Seas eventually, while releasing heat to the atmosphere along the way. The ocean-to-atmosphere heat flux is a predominant forcing in both ocean and atmosphere, and in our simulation experiments the heat release from ocean to atmosphere dramatically increased by 90% (from 0.19 to 0.36 PW) with the CAO events. Similar to Table 1 and Figure 13a for the mass transformation, the warm water heat transports for the three water masses are shown in Table 2 and Figure 13b, with an extra component of air-sea heat flux. (The precipitation was ignored in the mass balance). Overall, the CAO events increased the heat transport by 9% (from 2.22 to 2.42 PW) at the west boundary, 10% (from 0.76 to 0.84 PW) at the south, but no much change at the north (0.18 PW). However, over the entire MAR at the east, the CAO events decreased the heat transport by 6% (from 1.08 to 1.02 PW, with a large portion of it over the narrow CGFZ, from 0.52 to 0.48 PW), due to a combination of decrease in both temperature and mass transport.

14 26-14 LUO ET AL.: MODIFICATION BY COLD AIR OUTBREAKS lower surface of 11 C), it was not the case for heat transport. The lateral heat transports at the four boundaries (0.44, 0.04, 0.46, 0.00 PW at the west, south, east and north, respectively) resulted in a net lateral divergence of 0.06 PW. This layer also lost 0.06 PW to the atmosphere at the outcropping area at the northwest. Adding 0.25 PW from the SLW, the heat passing through the bottom of UTM (11 C isotherm) was 0.13 PW (downward), 0.12 PW less than that at the upper surface of this layer. The CAO events drew the same heat (0.44 PW) through the west boundary but exported 0.02 PW more through the south and 0.14 PW less through the east and released 0.03 PW more to the air, which was balanced by 0.47 PW heat gain from the SLW and 0.44 PW heat loss to the LTW. [39] Even the lowest layer, the LTW, had a small outcropping area in the northwest, resulting in a heat loss to the atmosphere of 0.04 PW from this layer in the mean experiment. Adding to the lateral heat transports of 0.38, 0.02, 0.26, 0.18 PW at west, south, east, and north, respectively, and the heat input from the upper layer (UTW) of 0.13 PW, this layer exported a small amount of heat, 0.01 PW, to the water below. The CAO events did not produce much significant heat transport across the 7 C isotherm ( 0.02 PW) either but did affect the heat transports through the other boundaries of this layer, i.e., the LTW received 0.31 PW more heat from the UTW, released 0.08 PW more to the air and exported 0.22 PW more to the east. Figure 13. Schematics of (a) volume transport and water mass transformation (Sv) and (b) heat budget (10 2 PW) for three layers of the warm water defined in the text and Tables 1 and 2. [37] For the SLW, in the mean experiment, the heat transports in the west, south, east, and north boundaries were 1.40, 0.70, 0.36, and 0.00 PW, respectively. The heat loss through the air-sea interface was 0.09 PW. The five components resulted in a net heat transport through the bottom of the SLW (14 C isotherm) of 0.25 PW (downward). The CAO events increased/decreased the above numbers to 1.60, 0.76, 0.22, 0.00, 0.15, and 0.47 PW, respectively. [38] Although the UTW, in the mean experiment, was vertically transparent for water mass (4 Sv entered the upper surface of 14 C and the same amount got out of the 6. Summary [40] In this paper the Princeton Ocean Model has been used to study the warm water pathways, transports and water mass transformation in the crossroads of the North Atlantic Ocean-Newfoundland Basin region and to investigate the response of the above processes to the consecutive occurrence of wintertime cold air outbreaks. Previous studies have focused on the short-time and local modification of the ocean to a single CAO event, while our study has focused on the ocean s large scale and low frequency response to a series of wintertime high-frequency CAO events. The model was first spun up for 10 years using a constant forcing, and the simulations indicated that the regional mean circulation, in terms of mean flow, transports, features of the 27.5 s t isopycnal surface, and kinetic energy, was in good agreement with observations, giving us confidence for proceeding with the CAO experiment. [41] To investigate the CAO effects, the model, from the 10-year spin-up, was run for another 3 years using the same forcing except for periods when the representative CAO s surface wind and heat flux were imposed. Results show that the thermal structure in the thermocline was significantly impacted: the thermocline deepened and the signal of changing temperature reached down to 500 m. This was different from the oceanic response to a single CAO event, where only the upper mixed layer deepened and the thermal structure of the main thermocline was not much influenced. Dynamical investigations indicated that the temperature anomaly was propagated by vertical diffusion during the CAO, and by vertical advection and horizontal advection between CAO events. The volume-averaged MKE and EKE increased by 30% and 20%, respectively. The CAO events

15 LUO ET AL.: MODIFICATION BY COLD AIR OUTBREAKS Table 2. Same as Table 1 but for Heat Transports and Balances (in PW) a a Note there is an extra air layer for heat balances. drew more warm water from the GS to the NAC towards the northeast, and its pathway in the Northwest Corner shifted towards the southeast owing to warm-to-cold water transformation. [42] The transports shown in the water transformation schematic (Figure 13a) were largest in the warm surface layer along the western boundary but were mostly in the cooler layers along the eastern boundary. This layer change in transport was an indication of water mass transformation that occurred in this region. In the mean state, 21% (4 out of 19 Sv) of SLW carried in by the GS was transformed to the cool UTW. However, since 4 Sv of UTW was transformed to the even cooler LTW, the net volume of UTW was not affected by this vertical flux. In the lower thermocline layer no LTW was further transformed to cold water; therefore there was no formation of Deep North Atlantic Water. With the influence of the CAO events, the water mass transformation rate was doubled: in the surface layer, 33% (7 out of 21 Sv) of SLW was transformed to UTW, while 9 Sv of UTW was transformed to LTW. However, similar to the mean state case, the rectification effects from the 12 consecutive wintertime CAO events did not reach the bottom of the LTW, i.e., the water mass transformation at 7 C isotherm was near zero. [43] We conclude, based on the results of this study, that high-frequency surface forcing is necessary to accurately model both mass and heat transports, especially in regions where cooling is dominated by extreme events of short duration. This is true even beyond regional simulations. For example, the incorporation of high-frequency forcing in global models that are used to predict the effects of climate change across the North Atlantic should improve estimate of heat that is transported to Europe. New research has shown that the number of explosive extratropical cyclones has increased over the last 21 years, and although these events occur predominately in cold seasons, the increase in occurrence might be linked to greenhouse warming [Lim and Simmonds, 2002]. [44] Acknowledgments. Richard Smith and Mathew Maltrud (Los Alamos National Laboratory) graciously shared their model results with us. We thank Tom Rossby for the RAFOS data and for sharing his insights of the study region. Critical comments and suggestions from two reviewers greatly helped us in the revision. This research was supported by National Science Foundation Grants OCE References Adamec, D., and R. L. Elsberry, Numerical simulations of the response of intense ocean currents to atmospheric forcing, J. Phys. Oceanogr., 15, , 1985a. Adamec, D., and R. L. Elsberry, The response of intense oceanic currents systems entering regions of strong cooling, J. Phys. Oceanogr., 15, , 1985b. Bane, J. M., Jr., and K. E. Osgood, Wintertime air-sea interaction processes across the Gulf Stream, J. Geophys. Res., 94, 10,755 10,772, Barnier, B., L. Siefridt, and P. Marchesiello, Thermal forcing for a global ocean circulation model using a three-year climatology of ECMWF analyses, J. Mar. Syst., 6, , Beckmann, A., and D. Haidvogel, Numerical simulation of flow around a tall isolated seamount, J. Phys. Oceanogr., 23, , Blumberg, A. F., and G. L. Mellor, A description of a three-dimensional coastal ocean circulation model, in Three-Dimension Coastal Ocean Models, Coastal Estuarine Ser., vol. 4, edited by N. Heaps, 208 pp., AGU, Washington, D.C., Brügge, B., Near surface mean circulation and kinetic energy in the central North Atlantic from drifter data, J. Geophys. Res., 100, 20,543 20,554, Bubnov, V. A., The North Atlantic Current by the Atlantex 90 experiment data, Oceanology, 34, , Budyko, M. I., Climate and Life, 508 pp., Academic, San Diego, Calif., Carr, M.-E., and H. T. Rossby, Pathways of the North Atlantic Current from surface drifters and subsurface floats, J. Geophys. Res., 106, , Elsberry, R. L., and N. T. Camp, Oceanic thermal response to strong atmospheric forcing: I. Characteristics of forcing events, J. Phys. Oceanogr., 8, , Ezer, T., and G. L. Mellor, Simulations of the Atlantic Ocean with a free surface sigma coordinate ocean model, J. Geophys. Res., 102, 15,647 15,657, 1997.