Constraints on the extent of the Ross Sea phytoplankton bloom

Tasha E. Reddy1 and Kevin R. Arrigo1

Received 3 October 2005; revised 20 January 2006; accepted 21 March 2006; published 4 July 2006.

[1] An intense phytoplankton bloom forms over the continental shelf in the Ross Sea,Antarctica, every spring as the sea ice retreats during the expansion of the Ross Seapolynya. The extent of this bloom rarely covers the full area of the shelf for reasonsthat are not well understood. One clue may be found in satellite imagery spanning the last25 years that shows frequently recurring, previously undescribed, large-scale plumes ofhigh-chlorophyll water and jets of low-chlorophyll water demarking the northernboundary of the spring phytoplankton bloom. The plumes, with horizontal length scales of100–200 km, appear at the same location each year. The spatial consistency of thispattern of jets and plumes suggests regulation by forcing that does not vary from year toyear, such as the interaction between bathymetry and prevailing ocean currents. Here wepresent a modeling study indicating that this pattern in the spring phytoplankton bloom iscaused by intrusions of offshore surface waters onto the continental shelf, guided byrelatively deep polar shelf bathymetry. Our study suggests that currents from the Ross Seagyre follow along the shelf break and prevent the phytoplankton bloom from extendingfarther toward the edge of the shelf, even when the waters there are ice-free, perhapsbecause the offshore waters have low iron concentrations. These results illustrate howphytoplankton distributions can be constrained by the interaction of deep bathymetryand surface geostrophic currents along the highly productive Antarctic continentalshelf.

Citation: Reddy, T. E., and K. R. Arrigo (2006), Constraints on the extent of the Ross Sea phytoplankton bloom, J. Geophys. Res.,

111, C07005, doi:10.1029/2005JC003339.

1. Introduction

[2] Phytoplankton chlorophyll a (Chl) concentrations inthe Southern Ocean are generally higher on continentalshelves, likely owing to the greater abundance of iron inshallow waters [Comiso et al., 1993]. The Ross Sea conti-nental shelf (Figure 1) is no exception, containing some ofthe most biologically productive waters in Antarctica[Arrigo and van Dijken, 2003a]. Phytoplankton biomassoften exceeds 10 mg Chl m�3 [Smith and Gordon, 1997]during intense spring phytoplankton blooms that form whenlight availability in surface waters increases due to theexpansion of the Ross Sea polynya (an area of open watersurrounded by sea ice).[3] While the extent of the sea ice is one factor constrain-

ing the ultimate size of the phytoplankton bloom, evenwhen the continental shelf is virtually ice-free, the bloomrarely extends to the outer reaches of the shelf. Furthermore,the extent of the bloom can vary dramatically from year toyear [Arrigo et al., 1998; Arrigo and van Dijken, 2004].Clues to what is limiting the spatial extent of the phyto-plankton bloom may be found in ocean color satelliteimagery of the region collected between 1978 and 2003

[Arrigo and Van Dijken, 2004; Arrigo et al., 1998] whichshow large plumes of high-Chl water (Figure 2, areas Athrough C) separated by regions of low-Chl water (Figure 2,lines 1 through 3) that develop during the peak of the springbloom. The area of low-Chl along the northwestern bound-ary of the bloom (line 1) is not spatially fixed from year toyear and may be associated with the position of the retreatingice edge, explaining its absence in 2003 (Figure 2d). The othertwo jets of low-Chl (lines 2 and 3), however, are quiteconsistent from year to year. They flank both sides of thePennell Bank, suggesting a correlation between low surfaceChl and deep bathymetry. During years when bloom extent islarger than average, the easternmost jet of low Chl (line 3)does not appear. However, virtually every year that thesouthwestern Ross Sea was ice-free during spring (largeblooms did not develop during some years because of heavysea ice due to icebergs or unusual climate conditions[Arrigo and van Dijken, 2003b, 2004]), the center jet oflow Chl (line 2) bordered by two plumes of high Chl isspatially fixed and strikingly apparent in satellite imagery.[4] The formation of these mesoscale features is likely

driven by surface circulation in this region of the SouthernOcean. Circulation in the southwestern Ross Sea isdominated by the swift surface-intensified slope current[Assmann et al., 2003], the southern part of the cyclonic(clockwise) Ross Sea Gyre that runs northwest along theshelf break (Figure 1). Some of this water flows onto theRoss Sea continental shelf, which averages approximately

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1Department of Geophysics, Stanford University, Stanford, California,USA.

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600 m in depth but which has three shoals 300 m deepand a steep shelf break (Figure 1). Previous modelingstudies have found that the circulation on the shelf isstrongly affected by this irregular bottom topography[e.g., Dinniman et al., 2003].[5] There are a variety of ways in which relatively deep

bathymetry can alter surface circulation [Heath, 1981;Holland, 2001; Semtner and Chervin, 1992] and affectphytoplankton distributions. The most common mechanismis the topographically enhanced entrainment of nutrientsinto surface waters and subsequent stimulation of phyto-plankton growth. This process has been observed in manylocations, including thewesternAntarctic Peninsula [Prezelinet al., 2000], the Subtropical Convergence at Chatham Rise[McClatchie et al., 1997], and upwelling filaments off of thecoast of California [Flament et al., 1985] and southwestAfrica [Shillington et al., 1990]. However, the upwelling ofnutrients does not likely affect phytoplankton distributions inthe Ross Sea, as springtime nutrients (including micro-nutrients such as iron) are generally high on the conti-nental shelf owing to convective overturning duringwinter. Another way that phytoplankton distributions areaffected by deep bathymetry is through lateral advectionof surface waters that are bathymetrically constrained tocreate heterogeneous patterns in surface phytoplanktondistributions. For example, Moore and Murdoch [1993]observed a nonupwelling, bathymetrically constrained,offshore-directed plume of high Chl near the coast ofNew Zealand. However, there are few other examples ofsuch features occurring along nonupwelling coasts. Finally,the depth of the well-mixed surface layer can be altered as itflows over irregular bathymetry [Stewart, 2004]. Thesechanges in mixed layer depth (MLD) can alter light avail-ability, and consequently, phytoplankton growth rates andabundance.[6] The purpose of the present study was to investigate

the physical processes that control the size and shape of thephytoplankton bloom in the southwestern Ross Sea. Inparticular, we were interested in whether the surface ex-

pression of the bloom is bathymetrically constrained, and ifso, by what mechanism. This was investigated using athree-dimensional ocean ecosystem model of the RossSea, specifically, the Coupled Ice, Atmosphere, Ocean(CIAO) model [Arrigo et al., 2003; Worthen and Arrigo,2003]. This model simulates phytoplankton dynamics as afunction of seasonal changes in temperature, light, andnutrient distributions, which are in turn controlled bypatterns of advection and vertical mixing. By determiningthe conditions under which CIAO could accurately simulatethe major distributional patterns of the phytoplanktonbloom, we were able to assess whether the observedmesoscale features form regularly due to bathymetric con-trol of either advective flow or mixed layer depth. While thefeatures themselves are not likely to have a significant effecton carbon uptake or determine upper trophic level foragingpatterns, identifying the mechanisms of their formation willhelp us understand how mesoscale physical oceanographicprocesses constrain phytoplankton distributions on thishighly productive continental shelf.

2. Methods

2.1. Remote Sensing

[7] Spatial patterns in Chl were derived from Sea-viewingWide Field-of-view Sensor (SeaWiFS 1997–2003) andCoastal Zone Color Scanner (CZCS 1978–1981) satelliteimagery obtained from the Goddard Earth Sciences Dataand Information Services Center (DAAC). We processedSeaWiFS data using the NASA SeaDAS image processingsoftware and the OC4v4 algorithm. CZCS pigment concen-trations were calculated using the standard algorithm ofGordon et al. [1983]. Despite heavy cloud cover onaverage, clear images were obtained during ice-free yearswith which to characterize patterns of Chl and their rela-tionship to bathymetry so that averaging over multiple dayswas not necessary. Sea ice distributions were determinedfrom daily Special Sensor Microwave Imager (SSM/I) dataavailable at the EOS Distributed Active Archive Center

Figure 1. (a) Map of the Antarctic coastline and 1000-m bathymetric contour with study region outlinedin blue. A schematic of the Ross Sea Gyre is shown in green. (b) Bathymetry map for the study area alsoincluding Ross Sea Gyre circulation schematic (green). Purple line marks the boundaries of imagesdisplayed in Figure 2.

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(DAAC) at the National Snow and Ice Data Center, Uni-versity of Colorado, Boulder, Colorado.

2.2. Modeling

[8] To investigate possible interactions between surfacephytoplankton blooms and bathymetry, we increased thehorizontal resolution of CIAO, a complete description ofwhich is given by Arrigo et al. [2003] and Worthen andArrigo [2003], from 25 km to 8 km in the southwesternRoss Sea. The physics of CIAO are based on the PrincetonOcean Model (POM), a primitive equation, sigma-coordi-nate ocean circulation model [Blumberg and Mellor, 1987]with vertical mixing calculated using the turbulence closurescheme of Mellor and Yamada [1982]. CIAO specifies dailysea ice concentration from a 20-year climatology derivedfrom passive microwave satellite data to calculate surfaceheat and salt fluxes.[9] The biological component of the model includes the

two dominant phytoplankton groups in the region (diatomsand Phaeocystis antarctica), macronutrients (silicate, phos-phate, and nitrate), the micronutrient iron (Fe), zooplankton,and detritus. The Fe cycle used in this version of the modelis simple and includes neither photochemistry of Fe norscavenging of Fe by particles, but still captures first orderdynamics of this essential micronutrient through remineral-ization, mixing, and advection [Worthen and Arrigo, 2003].Remineralization of detritus is a source of dissolved Fe, andthere is weak restoring of Fe in the bottom layer of themodel which simulates a sedimentary Fe flux. Phytoplank-ton growth rate is computed as a function of temperature,light, and nutrient concentration.[10] Using climatological data sets as described by Arrigo

et al. [2003], the model was spun up for 3 years and thenrun for 1 year to simulate the bloom over an average annualcycle. The model is able to accurately simulate the temporaland spatial evolution of two well-understood phytoplanktonblooms that occur annually in the Ross Sea polynya. Theprogression of the blooms and the magnitude of primaryproduction are in excellent agreement with both in situobservations and satellite ocean color data [Arrigo et al.,2003; Tagliabue and Arrigo, 2005; Worthen and Arrigo,2003].

3. Results

[11] As described previously, the Chl distributions in theRoss Sea during the spring phytoplankton bloom appear tobe related to the bathymetry on the continental shelf, 300–600 m below the surface (Figure 2). The regions of low Chl(Figure 2e, lines 1 through 3) run parallel to the bathymetrycontours of the troughs between the banks, but are actuallyneither in the troughs nor on the shoals. Line 1 (Figure 1)runs down the center of Mawson Bank and lines 2 and 3border the largest shoal on the shelf, the Pennell Bank.[12] Modeled surface velocities exhibit flow almost par-

allel to bathymetric contours, both along the continentalshelf break and on the shelf (Figure 3). Currents along theshelf break move at approximately 0.1 m s�1, similar toflow estimates from previous modeling studies [Assmann etal., 2003; Van Woert et al., 2003; Dinniman et al., 2003].Our results suggest that jets of oceanic water move onshoreon both sides of the Pennell Bank and along the center of

Figure 2. Satellite imagery of chlorophyll from December(a) 1978, (b) 1998, (c) 1999, and (d) 2003, and(e) bathymetry. Areas of low chlorophyll associated withthe bloom are indicated by white lines (1–3). The labels areincluded on all images regardless of whether the feature ispresent or not during that year. Gray areas indicate sea ice,and black regions represent areas of missing satellite oceancolor data, most often due to cloud cover. Note that in 2003(Figure 2d), the western boundary of the bloom is locatedwestward of the usual location (see text).

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the Mawson Bank while offshore flow occurs along thecanyon between Mawson and Pennell banks. Althoughmost previous modeling studies [Assmann et al., 2003;Van Woert et al., 2003] were too coarse to resolve featuresin the circulation on the shelf at a resolution suitable for thepurpose of this investigation, our results agree with circu-lation patterns derived from a higher resolution study[Dinniman et al., 2003] showing onshore flow at 300 mdepth flanking both sides of the Pennell Bank. Also con-sistent with our findings is the observation that an influx ofwater from offshore may be concentrated along the westernflanks of the submarine banks on the shelf [Hofmann andKlinck, 1998]. Current patterns calculated by CIAO areconsistent with the few measurements that are available,including long-term current meter observations of northwestflow along the shelf break of 6.4 cm s�1 at 230 m over aduration of 399 days and one observation on the shelf at230 m with velocities of <1 cm s�1 (Picco et al. [1999]cited by Dinniman et al. [2003]). Locarnini [1994] usedcurrent meter measurements on the Ross Sea continentalshelf between 250 m and 450 m to identify a northwardflow along the western side of the shelf (near 170�E) as wellas in the central shelf region (near 175�W), in agreementwith our model results. Unfortunately, their data are verysparse and observations were not available near the shoals.A study of sea ice trajectories (determined by tracking icefloes over time) show ice in the northwest area of the shelfmoving in a northward direction [Arrigo et al., 1998],consistent with our modeled circulation patterns.[13] The model produces plumes of high and low-Chl

waters (Figure 4a) that correspond remarkably well to thehigh and low-Chl waters seen in satellite imagery (Figure 4b).Although the horizontal extent and Chl concentrations of themodeled bloom are greater than the bloom observed mostyears in the satellite imagery, the model is able to capture thetwo areas of low Chl on the west side of the shelf (Figure 4a,lines 1 and 2). Surface salinity and temperature calculated bythe model (Figures 5a and 5b) also exhibit alternating spatialpatterns of high and low values similar to that of Chl.Specifically, areas of reduced salinity and temperature corre-

spond to regions of low Chl. This pattern was most pro-nounced in the salinity distributions (Figure 5a), which is notsurprising considering that density differences in polar watersare more strongly influenced by salinity than by temperature.Vertical profiles of temperature and salinity (Figures 5c and5d) show that the horizontal patterns of these two properties inthe surface plots are restricted to water masses isolated abovethe pycnocline.

4. Patterns in Chl and Mechanisms for TheirFormation

[14] Two mechanisms by which deep bathymetry canaffect phytoplankton distributions in surface waters, asidefrom facilitating the upwelling of nutrients, a process notlikely to be important in the high-nutrient Ross Sea, includealteration of MLD and modification of surface advectionpatterns. MLD can be altered when stratified waters flowover irregular bottom topography. If the flow velocity overbottom topography is subcritical (slower than the speed atwhich waves travel), the pycnocline will be displacedupward and MLD will decrease. If the flow is supercritical(faster than the speed at which waves travel), the pycnoclinewill be displaced downward and MLD will thicken. Thelocation of the deepening or shoaling relative to the topog-raphy in each of these cases is impacted by local Coriolisstrength. In areas where MLD has deepened, phytoplanktonare mixed out of the euphotic zone and bloom formation isinhibited. A shallower MLD is often characterized by highermean light levels, enhanced phytoplankton growth, andhigher Chl concentrations. Changes to MLD through thismechanism can be substantial in a stratified oceanic envi-ronment where the interface between the two internal layersof different densities can be depressed or raised to a muchgreater extent than the sea surface [Stewart, 2004].[15] If this process were important in the Ross Sea, we

would expect to see a shoaling or deepening of the mixedlayer associated with the stratified flow over the three largebanks on the continental shelf (Figure 1). Depth profiles oftemperature and salinity from CIAO (Figures 5c and 5d)reveal that the pycnocline remains relatively uniform overthe shoals, ranging from 10 to 30 m. Furthermore, waters onthe Ross Sea continental shelf are generally weakly strati-fied [Arrigo et al., 1998, 2000], rendering it less likely thatstratified flow could alter the mixed layer enough to affectphytoplankton growth. This indicates that stratified flowover the shoals is not markedly altering the depth of thepycnocline to create the patterns in Chl, salinity, andtemperature, either in the model or in nature.[16] Model results suggest instead that the surface pat-

terns in Chl are caused by horizontal advection of differentsurface water masses, as revealed by vertical profiles oftemperature and salinity above the pycnocline (Figures 5cand 5d). Waters with low Chl are associated with distinctwater masses having low salinity and temperature (Figure 5)that advect in fromnortheast of the continental shelf.Modeledinstantaneous horizontal surface velocities (Figure 4a) givefurther indication that the patterns of high and low Chl arecaused by horizontal advection. The surface velocities showthat onshore transport of oceanic waters are bathymetricallyconstrained to the troughs in between the shoals (Figure 3).This pattern of modeled surface velocity persists throughout

Figure 3. Instantaneous surface velocity streamlines over-lain on bathymetry.

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December and January. Velocities along the shelf break areapproximately 10 cm s�1; onshore velocities between theshoals are approximately 4 cm s�1, and offshore velocities areapproximately 2 cm s�1.[17] Chl concentrations in offshore waters that advect

onto the shelf in association with the jets are most likelylow because these waters have not been ice-free longenough for a bloom to form. As the temperature and salinitycharacteristics of waters within the jets shows (relativelycold and fresh relative to surrounding waters, Figure 5),these jets advected onto the shelf from regions that hadrecently been ice covered. Because the lack of light in ice-covered waters reduces phytoplankton growth rates dramat-ically, these newly ice-free source waters for the jets (suchas along line 2 in Figure 4) have Chl concentrations of only

approximately 0.2 mg m�3 (Figure 2b). Given phytoplank-ton growth rates at 0�C of 0.2 d�1 in open waters of theRoss Sea, it would take approximately 30 days to reachbiomass levels characteristic of the surrounding high Chlwaters within the bloom. This is longer than the approxi-mately 17 days that it takes for the jets to move fromoffshore into the bloom. Thus there is insufficient time forphytoplankton abundance within the jet to reach concen-trations characteristic of surrounding waters.[18] Low iron concentrations in the waters being

entrained into the plume from offshore may also play arole in maintaining the persistently low Chl concentrationswithin the dominant jet (Figure 6 line 2). Results fromCIAO show that iron concentrations in December are low(below growth-limiting levels for phytoplankton) both

Figure 4. Comparison of December chlorophyll distributions from (a) model results (velocitystreamlines are also shown) and (b) SeaWiFS imagery from 1998.

Figure 5. Model results showing surface (a) salinity and (b) temperature and vertical profiles along atransect of 74.8�S of (c) salinity and (d) temperature.

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within this jet and in the offshore waters entrained into it(Figure 6b, line 2). In situ iron concentrations measuredthroughout the southwestern Ross Sea support the idea ofiron limitation within offshore source waters for the jets.Although iron abundance in surface waters on the RossSea continental shelf before the spring bloom is relativelyhigh (0.1–0.4 nM), concentrations decrease markedlynorthward, falling to 0.025 nM (in midsummer) at theedge of the continental shelf and to <0.05 nM (yearround) in the Antarctic Circumpolar Current [Coale et al.,2005]. However, it should be noted that iron concentrations inseawater and particularly in ice and snow are still not wellknown and have been parameterized in CIAO based on only alimited number of samples from the Ross Sea [Coale et al.,2005; Edwards and Sedwick, 2001]. More in situ data areneeded to firmly establish surface Fe concentrations in theRoss Sea and to ensure that iron concentrations are beingproperly represented in nonshelf waters within CIAO.[19] If the areas of reduced Chl concentration within the

phytoplankton bloom result from the advection of offshorewaters low in Chl into the bloom, this may explain why theeasternmost jet of low-Chl water (Figure 2, line 3) is lesspronounced during years when the phytoplankton bloom ismore extensive (Figures 2a–2c) and why it is not wellrepresented by CIAO (Figure 4a, line 3). Unlike thedominant low-Chl jet (line 2), which is formed from thesouthwestward flow of mostly nonshelf waters ontothe continental shelf (Figure 3), much of water formingthe easternmost jet of low Chl water originates on the shelf,flows northward toward the shelf break, then turns backtoward the southwest before forming the jet and flowinginto the bloom (Figure 3, line 3). During years when thebloom is more extensive (e.g., 1978 and to a lesser extent,1998, Figures 2a and 2b), waters on the shelf where the jetoriginates will have a high Chl content. Consequently,although this northward flow entrains some low-Chl off-shore water before moving back onto the shelf, the jet ismuch less pronounced than in years when the bloom on theshelf is less extensive (Figure 2c).

5. Advection via Ekman Layer, Tidal, andGeostrophic Flow

[20] Having determined that horizontal advection controlsphytoplankton distributions in the Ross Sea, we nextconsider which of a number of different advective processes

(surface Ekman dynamics, tides, and geostrophy) couldrespond to bathymetric features hundreds of meters belowsurface and may be responsible for the patterns of advectionseen in surface waters. On low-latitude continental shelves,phytoplankton blooms are often affected by the relativelyshallow bathymetry through turbulent mixing to the seafloor[Lucas and Cloern, 2002] or topographically induced eddiesbringing nutrients into the euphotic zone [Roman et al.,2005]. However, Antarctic continental shelves are deepcompared to low-latitude shelves (typically hundred ofmeters versus tens of meters, respectively), and theseprocesses are likely to be of little importance.[21] Surface circulation is most directly affected by wind

stress creating flow in the surface Ekman layer, but this typeof flow is unlikely to form consistent patterns in surfacecirculation related to bathymetry on the Ross Sea continen-tal shelf. Winds blowing on the sea surface produce a thinboundary layer which is at most a few hundred meters thick,but usually only 5–30 m thick [Stewart, 2004]. SurfaceEkman layer flow does not extend to the ocean floor in theRoss Sea, and cannot connect surface circulation to deepbathymetric features.[22] Tidal models [Padman et al., 2003] predict that tidal

currents are generally strong in the southwestern Ross Sea,particularly over the shoals on the continental shelf. Fur-thermore, tidally driven flow is spatially consistent overthese features from year to year. Nevertheless, model resultsrule out tides as a first-order factor in the formation of these

Figure 6. Modeled results for surface waters in December of (a) chlorophyll and (b) iron.

Figure 7. Instantaneous surface velocity streamlines over-lain on modeled surface density anomaly.

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jets and plumes, as these features appear in the model resultswhen tidal forcing is excluded. Additionally, tidal currentellipses [MacAyeal, 1984] are on the order of 10 to 100times smaller than the mesoscale patterns observed in thephytoplankton bloom. There is no obvious mechanism bywhich tidal flow, circling around in these small-scale tidalellipses, could create the observed patterns in the bloomthrough advection alone.[23] We conclude then that the plumes of water observed

in the model output and in the satellite imagery are causedby geostrophic flow, whereby fluid motions are maintainedas a result of near balance between the Coriolis effect andgravity-induced horizontal pressure gradient forces. Thisconclusion is based on the previously described relationshipbetween surface advection and the shoals on the continentalshelf, as well as velocity streamlines that are nearly parallelto isopycnal contours (Figure 7), a characteristic of flow ingeostrophic balance. Furthermore, polar regions experiencestrong rotation due to their high latitudes and are stronglyinfluenced by geostrophic flow. Consistent with this, cur-rents in the Ross Sea study region are dominated bygeostrophic flow, with well-documented evidence of thisflow apparent in geologic features from the Ross Sea upperslope [Anderson et al., 1984]. This type of flow extendsdeep into the water column and tends to move aroundobstacles such as shoals, instead of over them, connectingdeep bathymetric features to surface waters above.

6. Conclusions

[24] Previous studies provide evidence that annual pro-duction in the southwestern Ross Sea is limited by ironavailability [Cochlan et al., 2002; Olson et al., 2000; Arrigoet al., 2003]. In the present investigation we extend thesestudies to show that the horizontal distribution of high-productivity water is most likely constrained by geostrophicsurface currents that continually move waters low in Chl,and most likely, iron, both along the shelf break and ontothe shelf. Interannual variation in the maximum size of thephytoplankton bloom in the Ross Sea during years ofnormal sea ice cover may be the result of annual differencesin the balance between the strength of geostrophic circula-tion along the shelf break and upward mixing of iron fromdeeper waters on the shelf. Reduced flow within thesouthwestern leg of the Ross Sea gyre or more intensevertical mixing in winter in some years may allow thephytoplankton bloom to extend farther toward the conti-nental shelf break.[25] This study also gives some insight into questions

regarding the importance of Circumpolar Deep Water(CDW) moving onto the Ross Sea continental shelf as asource of nutrients for the spring phytoplankton bloom.Studies on the west Antarctic Peninsula continental shelfhave shown that topographically induced upwelling atshallow locations near the shelf break brings nutrient-richCDW into the euphotic zone and increases biologicalactivity [Prezelin et al., 2000]. Dinniman et al. [2003] havesuggested that similar processes may be important in theRoss Sea. Warm CDW upwelling onto the Ross Seacontinental shelf has been proposed to contribute to theformation of the seasonal polynya [Jacobs and Comiso,1989], but there remains little evidence that this water mass

is entrained into surface waters during the summer or isbiologically significant. In fact, CIAO results presented heresuggest that on the shelf and across the shelf break, there isa reorganization of surface flow as opposed to theentrainment of intermediate waters into the surface. Whileprior studies suggested that deep offshore water enhancesbiological activity in the Ross Sea polynya, we suggestalternately that in some regions, the offshore surfacewaters moving onto the shelf are inhibiting phytoplanktongrowth.

[26] Acknowledgments. We would like to thank D. Fong, O. Fringer,D. Holland, R. Labiosa, and S. Monismith for helpful discussions andcomments regarding this research. We also acknowledge the valuablesuggestions from two anonymous reviewers of this manuscript. This workwas supported by NASA headquarters under the Earth System ScienceFellowship grant NGT5-30489.

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�����������������������K. R. Arrigo and T. E. Reddy, Department of Geophysics, Stanford

University, Stanford, CA 94305-2215, USA. (tereddy@stanfordalumni.org)

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