Ocean fronts trigger high latitude phytoplankton blooms

Transcription

1 GEOPHYSICAL RESEARCH LETTERS, VOL.???, XXXX, DOI: /, 1 2 Ocean fronts trigger high latitude phytoplankton blooms J. R. Taylor, 1, R. Ferrari, 2 1 Department of Applied Mathematics and Theoretical Physics, University of Cambridge, Cambridge, UK., (J.R.Taylor@damtp.cam.ac.uk) 2 Department of Earth, Atmospheric and Planetary Sciences, Massachusetts Institute of Technology, Cambridge, Massachusetts, USA., (rferrari@mit.edu)

2 X - 2 TAYLOR AND FERRARI: BLOOMS AT HIGH LATITUDE FRONTS Density fronts are ubiquitous features of the upper ocean. Here, numerical simulations show that restratification at fronts inhibits vertical mixing, triggering phytoplankton blooms in low-light conditions. The stability of the water column at fronts is set by a competition between frontal instabilities, which restratify the upper ocean, and turbulent mixing, which acts to destroy this stratification. Recent studies have found that frontal instabilities can restratify the upper ocean, even in the presence of strong surface cooling and destabilizing winds. During winter at high latitudes, primary production by phytoplankton is generally limited by low ambient light levels and deep turbulent mixing. When the turbulent mixing, inhibited by frontal restratification, becomes smaller than a critical turbulence threshold, a phytoplankton bloom can develop. The finding that fronts can trigger phytoplankton blooms by reducing mixing, provides an explanation for satellite observations of high chlorophyll concentrations at high latitude fronts.

3 1. Introduction TAYLOR AND FERRARI: BLOOMS AT HIGH LATITUDE FRONTS X Free-floating photosynthetic micro-organisms, collectively known as phytoplankton, form the foundation of the marine food web. Phytoplankton are responsible for about half the global primary production (PP) [Longhurst et al., 1995] and contribute to the ocean uptake of carbon dioxide [Longhurst and Harrison, 1989], which is especially large at high latitudes [Laws et al., 2000; Takahashi et al., 2009]. Oceanic PP does not occur uniformly, but is punctuated by patchy blooms lasting a few weeks [Dutkiewicz et al., 2001]. This heterogeneity in space and time is poorly understood, and it contributes to the uncertainty in global estimates of PP and the ocean carbon uptake. Previous studies [Lévy et al., 2001; Mahadevan and Archer, 2000; Mahadevan and Tandon, 2006] have found that ocean fronts, separating waters with different densities, can generate intermittent blooms in oligotrophic subtropical waters by upwelling additional nutrients into the euphotic zone, defined here as the uppermost layer of the ocean with sufficient sunlight to support photosynthesis [Thurman and Trujillo, 1999]. Here, we show that fronts can also trigger phytoplankton blooms at high latitudes when growth is more limited by light exposure than nutrient availability. Frontal instabilities lead to a restratification of the upper ocean and reduce the turbulent flux of phytoplankton cells out of the euphotic 33 zone, thereby increasing the mean light exposure. As a result, fronts act as hotspots for biological activity not only in the subtropics, but also in the more highly productive subpolar oceans. This implies that fronts might be crucial players in the ocean ecosystem in subpolar regions where the ocean absorbs large quantities of carbon dioxide from the atmosphere, thereby impacting the global carbon cycle.

4 X - 4 TAYLOR AND FERRARI: BLOOMS AT HIGH LATITUDE FRONTS An example of a highly productive frontal region captured by the MODIS Aqua satellite is shown in Fig. 1. Here, a front is formed when warm waters carried north by the Gulf Stream encounter much colder water brought south in the Labrador Current. The dashed white contour line shows the 1000m isobath, delineating the southern tip of the Grand 42 Banks, historically one of the most productive fisheries in the world. The sea surface temperature (SST) indicates that the front is dynamically active, with evidence of eddies on scales ranging from km. At this time of the year, in early spring, a massive greening event known as the spring bloom develops in the subpolar waters across the North Atlantic. The satellite-derived chlorophyll concentration shown in Fig. 1 is more than an order of magnitude larger at the front than in the surrounding water, particularly east of 49 longitude, where the front exhibits many fine-scale bends and wrinkles. Since the highest chlorophyll concentrations are localized at the front, they are likely generated by a local process, rather than being advected from another location (the Gulf Stream flows to the east and then bends northward in the region shown in Fig. 1.) At high latitudes during winter, the ocean is strongly forced by the atmosphere, leading to a deep surface boundary layer (SBL), defined as the region of the upper ocean where turbulence is enhanced by surface forcing. Deep SBLs typically have high nutrient concentrations because they entrain deep nutrient-rich waters as they form. Despite having abundant nutrients, deep SBLs can inhibit phytoplankton growth by mixing phytoplankton cells below the euphotic zone, away from abundant light. Light limitation is particularly strong during high latitude winters when SBLs are deep and incident light levels are low. The timing of the spring bloom has traditionally been associated with the

5 TAYLOR AND FERRARI: BLOOMS AT HIGH LATITUDE FRONTS X shoaling of the SBL at the end of winter [Gran and Braarud, 1935; Sverdrup, 1953]. According to the critical depth hypothesis, the spring bloom begins when the SBL becomes shallower than a critical depth. Sverdrup [1953] derived an expression for the critical depth, under the assumption that phytoplankton cells are well-mixed in the SBL. Recently, Taylor and Ferrari [2011], extending the work of Huisman et al. [1999], suggested that the onset of the spring bloom is more closely associated with a reduction in turbulent mixing caused by changes in the atmospheric forcing. While this hypothesis has skill in predicting the overall timing of the spring bloom, it does not explain the intensification of PP and high chlorophyll levels at fronts. We hypothesize that fronts are hotspots for the development of blooms through a suppression of vertical mixing by frontal dynamics. In a non-rotating fluid, the hydrostatic pressure gradient associated with horizontal changes in density would cause a front to quickly restratify, with water on the light side of the front flowing over the top of the dense water. In the ocean, the hydrostatic pressure gradient is often in geostrophic balance with the Coriolis acceleration associated with the Earth s rotation, the so-called thermal wind balance [e.g. Tandon and Garrett, 1995; Rudnick and Luyten, 1996]. However, this equilibrium is unstable, and in the absence of forcing, frontal instabilities cause the front to restratify, albeit at a slower rate than in the non-rotating case. This restratification, in turn, reduces the vertical mixing within the frontal zone. According to the critical turbulence hypothesis [Huisman et al., 1999], weak vertical mixing reduces the flux of phytoplankton cells out of the euphotic zone. A bloom can then develop when the

6 X - 6 TAYLOR AND FERRARI: BLOOMS AT HIGH LATITUDE FRONTS increase in biomass due to photosynthesis outpaces the loss of cells in the euphotic zone from downward mixing, sinking, grazing, and other losses. Here, we focus on two very common types of frontal instability, namely symmetric and baroclinic instability, and examine their impact on mixing in the SBL and phytoplankton growth. The characteristics of both instabilities have been described in previous studies, [e.g. Stone, 1966; Boccaletti et al., 2007; Taylor and Ferrari, 2009]. Previous studies have also shown that these frontal instabilities are capable of restratifying the upper ocean, even in the presence of strong surface heat loss and winds [Fox-Kemper et al., 2008; Ma- 89 hadevan et al., 2010; Taylor and Ferrari, 2010; Thomas and Taylor, 2010]. Here, we consider turbulence generated by convection forced by an ocean heat loss to the atmosphere. We will show that a front of moderate strength ( 0.25 C/km), forced by surface cooling of 100W/m 2, reduces the vertical mixing rate enough to trigger a phytoplankton bloom. In comparison, the largest SST gradients at the Gulf Stream front shown in Fig. 1, are 2 C/km. In the supplementary material, we summarize previously-derived scalings for symmetric and baroclinic instability, and compare the frontal restratification to destabilizing atmospheric forcing for other physical parameters. The reduction in vertical turbulent mixing at fronts does not contradict the wellestablished observation that fronts are regions of enhanced upwelling of nutrients towards the surface. We find that the rate of vertical mixing in the SBL is reduced compared to an unstratified boundary layer, following the development of stratification. Symmetric and baroclinic instability drive motions that are very effective at transporting fluid properties 102 along density surfaces, or isopycnals. However, the relatively shallow isopycnal slope,

7 TAYLOR AND FERRARI: BLOOMS AT HIGH LATITUDE FRONTS X following restratification, suppresses vertical mixing. At the same time, circulations associated with the formation and instability of the front extend into the thermocline, and draw nutrient-rich fluid into the SBL [Thomas et al., 2008]. The reduction of vertical mixing in the SBL at fronts increases light exposure and leads to blooms in light-limited regions, while the upwelling of nutrients into the SBL leads to blooms in nutrient-limited conditions. 2. Numerical Simulations Numerical simulations provide a means to test the hypothesis that reduced mixing at fronts can trigger phytoplankton blooms. The simulation domains shown in Fig. 2 were chosen to examine the response of a phytoplankton population to distinct frontal dynamics. Domain 1 is assumed to be far enough away from the front so that its effects are not felt, while Domains 2 and 3 are embedded within the frontal zone. The simulations each start with a mixed layer of uniform density and depth H = 150m, and a linearly 115 stratified fluid below. The strength of the front, measured by the horizontal density gradient, is constant in each simulation. Turbulence is generated by cooling the surface with a constant heat flux of Q 0 = 100W/m 2. The phytoplankton concentration is initially uniform in the mixed layer (P = P 0 ) with no phytoplankton below the mixed layer. Based on the phytoplankton growth and loss rates used in the simulations (see Supplementary material, Eq. A9), the critical depth, H C = 100m, is shallower than the mixed layer depth, and according to the critical depth hypothesis [Sverdrup, 1953], net losses should outpace phytoplankton growth.

8 X - 8 TAYLOR AND FERRARI: BLOOMS AT HIGH LATITUDE FRONTS Initially, the front is in equilibrium, with the lateral pressure gradient due to changes in density balanced by the Coriolis acceleration associated with an along-front flow. This equilibrium is unstable, and the front proceeds to slump in a two-step process. In the first stage, the front becomes unstable to symmetric instability, which has recently been observed at the Kuroshio [D Asaro et al., 2011] and Gulf Stream fronts [Thomas et al., 2011]. By definition, symmetric instability is independent of the along-front direction, and it is captured in the two-dimensional slice of Domain 2. Later, meanders and eddies develop along the front, and slumping accelerates as a result of baroclinic instability. Although the eddies that eventually develop are three-dimensional, the most unstable modes of the instability are independent of the cross-front direction [Stone, 1970]. Since Domain 3 is a two-dimensional slice in the along-front and vertical directions, our analysis strictly applies only to the early stages of baroclinic instability, before strong three-dimensionality develops. Using two-dimensional slices has the advantage that only one frontal instability can develop in each simulation and allows for a cleaner discussion of the dynamics. Coupling between the various instabilities and three-dimensional effects will be important to examine in a future study, but they are not likely to alter the basic conclusions discussed below. In Domain 1, far from the front, the phytoplankton concentration remains well-mixed in the SBL, and decreases in time, in agreement with the prediction from Sverdrup s critical depth theory. After 6 days, the phytoplankton concentration is about 17% lower than its initial value in Domain 1, as shown in Fig. 3. In contrast, the phytoplankton response is fundamentally altered by the front. In Domains 2 and 3, the SBL restratifies,

9 TAYLOR AND FERRARI: BLOOMS AT HIGH LATITUDE FRONTS X and the phytoplankton grow near the surface where light is more abundant. Previous studies [Huisman et al., 1999; Taylor and Ferrari, 2011] have shown that phytoplankton growth is possible when the turbulent diffusivity 1, a measure of the intensity of mixing, drops below a critical value. Using the parameters from the simulations shown in Fig. 3, the critical turbulent diffusivity is κ c m 2 /s (see Supplementary Material). Far from the front, in Domain 1, the turbulent diffusivity resulting from surface cooling is 151 κ T m 2 /s, much larger than the critical value. In comparison, the turbulent diffusivity is much smaller in the frontal zone, with κ T m 2 /s in Domain 2 and κ T m 2 /s in Domain 3. The formation of surface-intensified blooms in Domains 2 and 3 is consistent with the fact that κ T is less than or close to the critical turbulence level. The reduction in the rate of vertical mixing at fronts can be viewed as a consequence of the vertical stratification that develops in the SBL as the inclined isopycnals at the front slump in response to frontal instabilities. In Domain 1, convective plumes transport phytoplankton from the surface to the base of the SBL, mixing rapidly in the vertical direction. In Domains 2 and 3, the turbulent velocities are comparable or larger than those in Domain 1, but the dominant motion is now oriented along the isopycnals. Mixing of phytoplankton slows down because it takes much longer to transport biota away from the surface along very slanted trajectories - the aspect ratio of the isopycnals is very small. The restratification process is noticeably different in Domains 2 and 3. Symmetric instability quickly restratifies the boundary layer in Domain 2, but the level of stratification remains relatively weak. In this case, the stratification develops over the first 12 hours until the

10 X - 10 TAYLOR AND FERRARI: BLOOMS AT HIGH LATITUDE FRONTS isopycnal slope is about 1/30, at which point restratification ceases. In Domain 3 the SBL remains unstratified for about 2 days, until baroclinic instability develops. After this stage, the SBL quickly restratifies, and by day 6 the isopycnal slope is much shallower than in Domain 2, about 1/250. Notice that the turbulent diffusivity is slightly smaller in Domain 2 than Domain 3, and therefore is not purely a function of the level of restratification, but depends on details of the frontal dynamics. The scalings discussed in section A2 of the supplementary material predict when frontal instabilities will be able to restratify the SBL, but do not directly predict the level of vertical mixing. Parameterizing the vertical mixing at fronts will be an important step towards improving biogeochemical models, but this is beyond the scope of the present study. 3. Discussion Permanent fronts like the Gulf Stream, shown in Fig. 1, are confined to a few regions like the western boundaries of midlatitude oceans and the Southern Ocean. While we expect these exceptionally large fronts to support phytoplankton blooms, the dynamical processes described here should apply much more generally. A network of fine submesoscale fronts on scales of 1 10 km are found in many regions of the ocean [Lévy et al., 2001; Capet et al., 2008; Klein and Lapeyre, 2009]. While the density difference across submesoscale fronts is much smaller than large fronts like the Gulf Stream, the more dynamically relevant quantity, the horizontal density gradient, can still be large since the density change occurs over a short distance. The strength of submesoscale fronts is commonly comparable to, or even larger than, the value considered here [Capet et al., 2008].

11 TAYLOR AND FERRARI: BLOOMS AT HIGH LATITUDE FRONTS X Reduced vertical mixing and enhanced mean light exposure at fronts provides a means for phytoplankton populations to survive high-latitude winters in sufficient numbers to re-emerge in the spring. Several previous studies have proposed mechanisms for the persistence of a small phytoplankton population in winter [Huisman et al., 2002; Backhaus et al., 2003; Ward and Waniek, 2007], but none considered the role of spatial heterogeneity. Fronts could maintain the wintertime phytoplankton population by locally enhancing mean light exposure and supporting intermittent mid-ocean blooms. It has been speculated that fronts play a similar role in oligotrophic subtropical waters by drawing more nutrients into the euphotic zone, triggering patchy blooms [Lévy et al., 2001; McGillicuddy et al., 2007; Mahadevan et al., 2008]. In both cases fronts act like oases, allowing the phytoplankton population a temporary respite in an otherwise forbidding environment. The mechanism described here is particularly relevant to high latitudes where light exposure is often a more limiting factor than nutrient availability. Since subpolar regions are generally more productive than the subtropics [Lévy, 2005], this mechanism could have a strong impact on the global PP. Reduced mixing at fronts also allows blooms to develop earlier in the season when the SBL is deeper than the critical depth, as seen in Fig. 3. Previous work has found that early blooms in deep SBLs are capable of generating 204 significantly more biomass than shallow blooms [Stramska et al., 1995]. Through this mechanism, blooms at high latitude fronts could increase the ocean uptake of carbon dioxide and play an important role the global carbon cycle.

12 X - 12 TAYLOR AND FERRARI: BLOOMS AT HIGH LATITUDE FRONTS Appendix A: Supplementary Material A1. Numerical simulation setup The computational method used here to study the evolution of phytoplankton at fronts solves the non-hydrostatic rotating Navier-Stokes equations in two-dimensional slices. Although variations are neglected in the direction normal to the computational domain, the velocity field retains all three components, and evolves according to u + u t T u fv = 1 p + ρ 0 x ν 2 u, v + u t T v + w dv G + fu = 1 p + dz ρ 0 y ν 2 v, w + u t T w = 1 p + b + ρ 0 z ν 2 w, b + u t T b + um 2 = κ 2 b, u = 0, (A1) (A2) (A3) (A4) (A5) where f is the Coriolis frequency. The two-dimensional computational domains are arranged as in Fig. 2 so that the gradient operator is ( / x, 0, / z) in Domains 1 and 2, and (0, / y, / z) in Domain 3. Here, x and y are the cross-front and along-front directions, respectively. For the simulations in Domains 2 and 3, the front is included by imposing a constant background buoyancy gradient, M 2, while M 2 = 0 in Domain 1. The buoyancy, b, is then defined as the departure from the background state b = b T M 2 x, (A6) where b T is the total buoyancy. Similarly, the total velocity u T is decomposed into departures u from a flow V G in thermal wind balance with M 2 : u = u T V G ĵ, V G M 2 /f. (A7)

13 TAYLOR AND FERRARI: BLOOMS AT HIGH LATITUDE FRONTS X - 13 The viscosity and diffusivity, ν and κ, are both set to m 2 /s, which maintains numerical stability and is much smaller than the turbulent diffusivity in the SBL in each case. The buoyancy is initialized with a mixed layer depth of H = 150m with a constant stratification below: b T (t = 0) = { HN M 2 x if H < z < 0, N0 2 z + M 2 x if z H, (A8) where N 0 = s 2. The frontal strength is set to M 2 = s 2 in Domains 2 and 3, and M 2 = 0 in Domain 1. Assuming that the buoyancy gradient is controlled by the temperature (with constant salinity), and using a thermal expansion coefficient of α = C 1, this frontal strength corresponds to a temperature gradient of about 0.25 C/km. The physical model is coupled with an idealized phytoplankton model, P t + u P = µ 0e z/h l P mp + κ 2 P, (A9) where µ 0 = 1day 1 is the maximum growth rate at z = 0, m = 0.1day 1 is a constant loss rate, h l = 10m, and κ = m 2 /s is a weak diffusion added to maintain numerical stability. Based on these parameters, Sverdrup s critical depth is H c = (µ 0 /m)h l = 100m, and less than the mixed layer depth, H = 150m. The initial phytoplankton profile is P (t = 0) = { 1 if H < z < 0, 0 if z H. (A10) The turbulent diffusivity κ T was calculated from the resolved vertical turbulent flux and the mean phytoplankton gradient in each simulation: κ T = w P P / z. (A11)

14 X - 14 TAYLOR AND FERRARI: BLOOMS AT HIGH LATITUDE FRONTS Here, the angled brackets denote an average over the full horizontal extent, the upper 100m in the vertical, and for one inertial period in time ending at t = 6days. Primes denote a departure from this average, e.g. P = P P. An approximate expression for the critical turbulent diffusivity is given in Eq. (16) of Taylor and Ferrari [2010], κ c h2 l m (µ 0 m) 2. (A12) Phytoplankton growth is expected when κ T κ c, and in our simulations κ c m 2 /s. A2. Scaling of restratification by frontal instability The stratification in the upper ocean is determined by a competition between restratification and mixing. As discussed in Thomas and Ferrari [2008], restratification can be driven by several processes including frictional effects, frontogenesis, and the potential energy release associated with frontal instabilities. Here, we focus on restratification associated with mixed layer baroclinic and symmetric instabilities as discussed in the main 223 text. When restratification driven by the frontal instability is able to overcome a de stabilizing atmospheric forcing, the mixed layer will become stably stratified, suppressing vertical mixing, and possibly triggering a phytoplankton bloom. In the simulations presented in this paper, the frontal strength and surface forcing have been selected so that symmetric and baroclinic instabilities are both able to restratify the mixed layer. In order to see how the restratification by frontal instabilities compare to the surface forcing in other conditions, it is helpful to review scalings for mixed layer restratification developed in previous studies. It should be noted that a general scaling for the vertical turbulent diffusivity as a function of the frontal strength and surface forcing remains unknown. Al-

15 TAYLOR AND FERRARI: BLOOMS AT HIGH LATITUDE FRONTS X though it is beyond the scope of this paper, parameterizing the vertical mixing at fronts will be an important step towards improving ocean models and their representation of primary production. Since the process of restratification is different for symmetric and baroclinic instability, we consider each separately. A2.1. Baroclinic Instability The baroclinic instability of mixed layer fronts (commonly referred to as mixed layer instability, or MLI) was studied in detail by Boccaletti et al. [2007]. The instability extracts potential energy from the front, leading to a slumping of the isopycnal surfaces and restratification. Fox-Kemper and Ferrari [2008] and Fox-Kemper et al. [2008] derived a scaling for the rate of restratification by MLI. The restratification by MLI can be compared to the destabilizing effect of surface forcing by comparing the associated vertical buoyancy fluxes. According to the scaling, the maximum vertical buoyancy flux associated with MLI is: H 2 ( b) 2 w b MLI = C MLI, (A13) f where C MLI 0.06 is an empirically determined scaling constant. The stabilizing effect of MLI competes against the destruction of buoyancy through a surface buoyancy flux, B 0, and an Ekman buoyancy flux (EBF ) associated with an unstable cross-front Ekman flow, EBF = ρ 1 0 τ (f 1ˆk b). The latter is important when the wind stress, τ, has a component in the direction of the thermal wind [Thomas and Lee, 2005]. A stability parameter can be defined as the ratio of the stabilizing and destabilizing buoyancy fluxes: R MLI = C MLI H 2 ( b) 2 f (B 0 + EBF ). (A14)

16 X - 16 TAYLOR AND FERRARI: BLOOMS AT HIGH LATITUDE FRONTS When R MLI << 1, the surface forcing is strong relative to the frontal restratification, and the mixed layer is expected to remain unstratified, while stratification can develop when R MLI >> 1. For the parameters used in this study, R MLI 57, consistent with the fact that the mixed layer restratifies in Domain 3. A2.2. Symmetric Instability The restratification by symmetric instability (SI) is significantly different than MLI. The stability criteria for SI can be expressed in terms of the potential vorticity: P V = (f +ω) b where ω = u is the relative vorticity. SI develops when P V < 0 [Hoskins, 1974] and quickly brings the P V to zero by restratifying the mixed layer (by increasing f z b) as seen in Domain 2 and in Taylor and Ferrari [2010]. SI grows faster than baroclinic instability in most cases when P V < 0 [Stone, 1970] and dominates the initial phases of frontal instability. However when P V 0 the symmetric instability stops. There is recent observational evidence that SI can also maintain a stable stratification under very intense atmospheric forcing [D Asaro et al., 2011]. Unlike MLI, which restratifies the entire mixed layer, SI is only capable of generating significant stratification below a convective layer, z < h. Taylor and Ferrari [2011] presented a scaling for the convective layer depth, h, in terms of the frontal strength, M 2 = b, the depth of the unstable SI layer, H, and the surface wind and buoyancy forcing. The ratio of the convective layer depth to the boundary layer depth, h/h, determines the relative importance of SI. When h/h << 1, a large fraction of the boundary layer will become stratified as a result of SI, while in the limit when h/h 1, the boundary layer will be unstratified. For a given set of parameters, the ratio h/h can be readily found by numerically solving the following

17 TAYLOR AND FERRARI: BLOOMS AT HIGH LATITUDE FRONTS X - 17 fourth order polynomial equation (see Taylor and Ferrari [2011]): ( M 4 f 2 ) 3 ( ) 4 h B 0 + EBF C SI H H 4 ( 1 h H ) 3 = 0. (A15) where C SI 14 is an empirical scaling constant, and the small entrainment coefficients have been neglected. Using the parameters from the simulations shown in Fig. 3, h/h 0.035, indicating that SI is expected to develop and restratify most of the mixed layer, as is indeed seen in Domain Acknowledgments. Notes 1. The turbulent diffusivity is the rate at which the area occupied by a tracer patch grows as a result of turbulent mixing. The turbulent diffusivity is typically many orders of magnitude larger than the molecular diffusivity, i.e. the rate of 247 spreading of a tracer patch by molecular agitation References Backhaus, J., E. Hegseth, H. Wehde, X. Irigoien, K. Hatten, and K. Logemann, Convection and primary production in winter, Marine Ecology Progress Series, 251, 1 14, Boccaletti, G., R. Ferrari, and B. Fox-Kemper, Mixed layer instabilities and restratification, Journal of Physical Oceanography, 37 (9), , Capet, X., J. McWilliams, M. Molemaker, and A. Shchepetkin, Mesoscale to submesoscale transition in the California Current System. Part I: Flow structure, eddy flux, and observational tests, Journal of Physical Oceanography, 38 (1), 29 43, 2008.

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20 X - 20 TAYLOR AND FERRARI: BLOOMS AT HIGH LATITUDE FRONTS McGillicuddy, D., et al., Eddy/wind interactions stimulate extraordinary mid-ocean plankton blooms, Science, 316 (5827), 1021, Rudnick, D., and J. Luyten, Intensive surveys of the azores front 1. tracers and dynamics, Journal of Geophysical Research, 101 (C1), , Stone, P., On non-geostrophic baroclinic stability, Journal of the Atmospheric Sciences, 23, , Stone, P., On non-geostrophic baroclinic stability: Part II, Journal of the Atmospheric Sciences, 27 (5), Stramska, M., T. Dickey, A. Plueddemann, and R. Weller, Bio-optical variability associated with phytoplankton dynamics in the North Atlantic Ocean during spring and summer of 1991, Journal of Geophysical Research, 100 (C4), , Sverdrup, H., On conditions for the vernal blooming of phytoplankton, Journal du Conseil International pour l Exploration de la Mer, 18, , Takahashi, T., et al., Climatological mean and decadal change in surface ocean pco2, and net sea-air co2 flux over the global oceans, Deep Sea Research Part II: Topical Studies in Oceanography, 56 (8-10), , Tandon, A., and C. Garrett, Geostrophic adjustment and restratification of a mixed layer with horizontal gradients above a stratified layer, Journal of Physical Oceanography, 25, , Taylor, J., and R. Ferrari, On the equilibration of a symmetrically unstable front via a secondary shear instability, Journal of Fluid Mechanics, 622, , 2009.

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22 X - 22 TAYLOR AND FERRARI: BLOOMS AT HIGH LATITUDE FRONTS SST (ºC) 30 Sea Surface Temperature Composite, March 14-21, Latitude Longitude SST (ºC) Sea Surface Temperature, March 20, Grand Banks Latitude km 10 km Longitude Chla (mg/m3) Chlorophyll Concentration, March 20, Latitude Longitude Figure 1. An example of a highly productive front captured by NASA s MODIS Aqua satellite. A weekly composite of the sea-surface temperature (SST) is shown in the top panel, and daily images of SST and inferred Chlorophyll concentration in the highlighted region are shown in the lower two panels. The dashed white line shows the 1000m isobath. The frontal zone is dynamically active on a range of scales from O(10 100km). At the same time, the chlorophyll concentration is much higher at the front than in the surrounding water, indicating that this front is richly productive. D R A F T October 23, 2011, 8:49pm D R A F T

23 TAYLOR AND FERRARI: BLOOMS AT HIGH LATITUDE FRONTS X - 23 Baroclinic Instability Thermal Wind Domain 3 Symmetric Instability Domain 1 Domain 2 Figure 2. Schematic of a frontal zone separating water masses with two different densities. Initially, the along-front thermal wind balances the tilted density surface. In time, however, the front becomes unstable first to symmetric instability, and later to baroclinic instability, both resulting in an increase in the stratification. Three shaded regions indicate the domains used in the simulations shown in Fig. 3. Domain 1 is assumed to be far enough away from the front that its effects are not felt, while Domains 2 and 3 are embedded within the frontal zone.

24 X - 24 TAYLOR AND FERRARI: BLOOMS AT HIGH LATITUDE FRONTS Domain 1 0 P/P Cross-front distance (m) Cross-front distance (m) Domain 3 0 Depth (m) P/P Depth (m) Depth (m) -50 Domain 2 0 P/P Figure Along-front distance (km) Numerical simulations in three two-dimensional slices though a frontal zone as illustrated in Fig. 2. The phytoplankton concentration (color) and density contours (black) are shown after 6 days. In each simulation, the initial SBL depth (H = 150m) was deeper than the critical depth (Hc = 100m), indicated by the dashed black line. Isopycnals are plotted at an interval of 0.1kg/m3. The initial phytoplankton concentration was uniform in the SBL P = P0, with P = 0 below the SBL. Unlike Domain 1 which is outside of the frontal region, a bloom develops in the frontal zones of Domains 2 and 3. D R A F T October 23, 2011, 8:49pm D R A F T