Ocean margin exchange—water flux estimates

Journal of Marine Systems 32 (2002) 107 – 137 www.elsevier.com/locate/jmarsys

Ocean margin exchange — water f lux estimates John M. Huthnance a,*, Hendrik M. Van Aken b, Martin White c, E. Desmond Barton d, Bernard Le Cann e, Emanuel Ferreira Coelho f, Enrique Alvarez Fanjul g, Peter Miller h, Joao Vitorino f a

Proudman Oceanographic Laboratory, Bidston Observatory, Birkenhead, Merseyside CH43 7RA, UK b Nederlands Instituut voor Onderzoek der Zee, P.O. Box 59, 1790 AB Den Burg, Texel, Netherlands c Department of Oceanography, NUI Galway, University Road, Galway, Ireland d School of Ocean Sciences, University of Wales — Bangor, Menai Bridge, Gwynedd LL59 5EY, UK e Laboratoire de Physique des Oceans, UFR Sciences 6, avenue Le Gorgeu, BP 809 29285 Brest Cedex, France f Marinha-Instituto Hidrografico, Rua das Trinas, 49, P-1249-093 Lisbon, Portugal g Puertos del Estado, Avenida del Parteno´n 10, 28042 Madrid, Spain h Plymouth Marine Laboratory, Prospect Place, West Hoe, Plymouth PL1 3DH, UK Received 27 October 2000; accepted 23 June 2001

Abstract Hydrographic data define source regions for deep, intermediate and upper-ocean water types found near the northwest Iberian margin; these water masses’ structure and geostrophic flow are described. Labrador Sea Water (LSW) moves south through the region; above this at 1000 m, particularly near the coast, Mediterranean Sea Outflow Water (MSOW) moves north. Near-surface waters are more seasonal, and we distinguish summer and winter regimes. Summer: above 600 m, poleward flow over the slope is at a minimum in spring and summer, and a maximum in the autumn/early winter. Overlying this is a general southward flow linked to wind stress. Summer upwelling along the Iberian shelf has associated offshore filaments stretching westward, and appears most intense around Finisterre (starting earlier, more persistent). The vertical upwelling flux is s/qf but lateral exchange in filaments may be more. Winter: it is proposed that upper layers of the water column offshore are fed from the north. Potential vorticity calculations suggest enhanced diapycnic mixing. However, other data indicate an east – west front down to about 300 m at 39 – 40jN, i.e., north – south convergence. This front bends to the north near the shelf edge, in accord with a warm flow along the upper slope, 0.15 m/s or more. Surface temperatures often show a much broader (hundreds of kilometers) northward extension of warm water, probably affected by winds. The thermocline from the previous summer can persist to January at least. Coastal waters show cooler river runoff. Budgeting of water fluxes is discussed in relation to direct estimates (from water mass properties and current measurements) and in relation to possible process contributions to fluxes. On/ off-shelf exchange has a relatively short time scale, estimated at about 12 days. Upwelling (mostly) and mixing may supply the nutrient content of about 200 m water depth for new production over the shelf. D 2002 Elsevier Science B.V. All rights reserved. Keywords: Water masses; Currents; Dispersion; Water mixing; Shelf edge; Exchange; ANE (NE Atlantic); Iberia; Galicia

1. Introduction *

Corresponding author. Tel.: +44-151-653-8633; fax: +44-151653-6269. E-mail address: [email protected] (J.M. Huthnance).

Ocean margin fluxes between productive shelf seas and nutrient-rich oceanic waters are important to

0924-7963/02/$ - see front matter D 2002 Elsevier Science B.V. All rights reserved. PII: S 0 9 2 4 - 7 9 6 3 ( 0 2 ) 0 0 0 3 4 - 9

108

J.M. Huthnance et al. / Journal of Marine Systems 32 (2002) 107–137

carbon and nutrient budgets in the ocean. A further aim in determining exchange processes at the ocean margin is a basis for developing global models to predict the impact of environmental changes on the ocean and especially coastal regimes. These issues have been considered in several interdisciplinary studies for varied contexts at the edge of the continental shelf: the Shelf Edge Exchange Processes (SEEP) study in the Middle Atlantic Bight, eastern USA (Walsh et al., 1988, et seq.; Biscaye et al., 1994, et seq.); the Coastal Transition Zone Experiment off western USA (Brink and Cowles, 1991, et seq.); the UK Land– Ocean Interaction Study (LOIS) Shelf Edge Study west of Scotland (Souza et al., 2001); ECOMARGE in the Gulf of Lions (Monaco et al., 1990). In its first phase (1993 –1996), the EU OMEX (Ocean Margin Exchange) study focused on Goban Spur southwest of Ireland (Wollast and Chou, 2001). OMEX-II focused on the northwest Iberian seasonal upwelling region in 1997– 2000. The region (Fig. 1) potentially represents extensive and productive eastern Atlantic and Pacific margins with fairly narrow shelves and upwelling for at least part of the year. OMEX was highly inter-disciplinary; it used Eulerian and quasi-Lagrangian approaches to help understand the physically driven biogeochemical system with important spatial and temporal variability. Previous studies of this area have concerned alongslope flow, associated eddies (Haynes and Barton, 1990), changing transport according to adjacent Atlantic baroclinic flow (Frouin et al., 1990; Maze´ et al., 1997), the progress of Mediterranean Sea Outflow Water (MSOW) along the slope (Shapiro and Maschanov, 1996) and eddies formed from it (e.g., Pingree and Le Cann, 1993; Pingree, 1995). Van Aken (2000a,b, 2001) has described the regional hydrography. The ARCANE study of intergyre north-east Atlantic circulation on broad and meso-scales took place contemporarily with OMEX-II, focusing on slope currents, instabilities, coherent structure generation and baroclinic tides (Le Cann et al., 2002a,b, in preparation). The EU project MORENA (1993 – 1996) included studies of hydrography, currents and production west of Iberia, with models of the physics (Stevens et al., 2000). The EU project SEFOS also included hydrographic sections and some current measurements in this region. Relevant data from these projects are used here.

Fig. 1. The north – west Iberian region studied in OMEX-II showing locations referred to in the text and of current meter data used.

This paper aims to estimate shelf-slope – ocean exchange transports for the northern part of the west Iberian margin, the OMEX-II focus. The outline basis is that as MSOW flows northwards along the slope, its salinity decreases (Sy) via lateral exchange. Hence, an effective lateral diffusivity K f 500 m2 s
1 can be estimated (Daniault et al., 1994), or a lateral exchange rate, Sy (slope current transport)/(salinity excess in current) f (slope current transport)/700 km (Huthnance, 1995). Our approach differs from the inverse box model of Maze´ et al. (1997), which spanned all of west Iberia: our area is smaller; we make greater use of measured constituent differences and currents, so that a budget balance can be checked, not only inferred. We also consider independent estimates of specific

J.M. Huthnance et al. / Journal of Marine Systems 32 (2002) 107–137

109

process contributions to the transports, exchange velocities or effective diffusivity, to support and add interest to the budgeting. We do not aim to provide values of fluxes in all detail. For further resolution in space and time, it is proposed that models such as Stevens et al. (2000) and Coelho et al. (this volume; an OMEX companion paper) should be used. Nevertheless, measurements are needed to test models and the present estimates can provide such a test. Its validity is checked by whether flux contributions to ‘‘boxes’’ form balanced budgets. In Section 2 we describe the hydrography of the NW Iberian context: water masses, currents, processes and seasonal phenomena. Section 3 discusses the budgeting basis and choice of budget boxes; the data and calculations are in Section 4. Independent estimates of process contributions to fluxes are given in Section 5 with discussion and conclusions in Sections 6 and 7.

2. Hydrography 2.1. Water masses, related circulation and inter-annual variability The hydrography of the European ocean margin has been described in a series of papers on the deep, intermediate, and thermocline water masses in the northeast Atlantic by Van Aken (2000a,b, 2001). These papers are based on an extensive historic data set, including hydrographic surveys from WOCE, JGOFS, ‘‘Galicia,’’ MORENA and SEFOS, as well as the OMEX data. Local hydrography in the OMEXII area should be considered in this context, because of limitations in quality, quantity and spatial coverage of the OMEX-II data. For example: the salinity accuracy during some OMEX cruises is relatively poor ( f 0.05), of the order of inter-annual changes in the salinity of the Eastern North Atlantic Central Water (Van Aken, 2001); excluding the upper ocean, oxygen and nutrient concentrations are relatively scarce, with calibration problems for some OMEX cruises. The water mass structure is well described by a H– S diagram (Fig. 2) in combination with a salinity section along the west European continental slope (Fig. 3). Between approximately 2500 and 3000 m is found a

Fig. 2. Potential temperature – salinity diagram of the OMEX-II stations from the summer of 1997 (open circles) against the background of a series of mean hydrographic lines for the continental slope between 40jN and Porcupine Bank (lines). Approximate positions of the water types discussed in this paper are indicated; the dashed oval shows the narrow ENACW envelope.

core of Northeast Atlantic Deep Water (NEADW), which originates from overflow from the Nordic Seas into the Iceland Basin. West of Porcupine Bank ( f 52jN), NEADW still contains over 40% Iceland Scotland Overflow Water (ISOW); near the continental slope off Galicia the amount of ISOW in the NEADW core is reduced to about 20%, due to additional mixing with the underlying Lower Deep Water (LDW) and the overlying cores of Labrador Sea Water (LSW) and Mediterranean Sea Outflow Water (MSOW; Van Aken, 2000a). At a depth of about 1800 m, a core of Labrador Sea Water moves from the Charlie– Gibbs Fracture Zone southwards through the region. In its undiluted form, LSW is characterised by a deep salinity minimum. Off Galicia, these salinity minima can occasionally be observed, but over the continental slope and further to the south, these deep minima disappear due to diapycnal and isopycnal mixing (Van Aken, 2000b).

110

J.M. Huthnance et al. / Journal of Marine Systems 32 (2002) 107–137

Fig. 3. Salinity distribution over the European continental slope between the 200 and 4000 m isobaths. The horizontal coordinate gives the along-slope distance to the latitude of 40jN.

Above the LSW layer, at 1000 m, a core of northwardmoving saline MSOW is observed (Figs. 2 and 3), hugging the continental slope. The salinity of this core decreases polewards along the continental slope due to isopycnal and diapycnal mixing with less saline water types (Van Aken, 2000b). Near northwestern Spain, MSOW is lost to Meddies (see Section 2.2) so that strong along-slope salinity gradients form off northern Spain between 8j and 6jW (at f 400 to 600 km in Fig. 3). Changes of temperature and salinity in the lower MSOW at approximately 1200 m in May 1999, not correlated with currents, were symptomatic of other variability, possibly a vertical movement of the lower boundary. In the upper layers, the ocean margin is influenced by southward transport in the subtropical anticyclonic gyre. Over the continental slope, however, the flow is seasonal: typically poleward over the upper slope in winter (Frouin et al., 1990); southward surface flow during the upwelling season. Above the MSOW core, the H – S curves have an inverted ‘‘S’’ shape (Fig. 2) with a salinity minimum at f 450 m. At shallower levels, both temperature and salinity increase upwards to the bottom of the seasonal thermocline, where a

slight salinity maximum is generally observed at about 100 m. In summer, strong temperature gradients are observed in the upper 100 m, with only limited salinity gradients. In the permanent thermocline, between the seasonal thermocline and the f 450 m salinity minimum, the distribution of Apparent Oxygen Utilization (AOU) shows a trend of southward ageing. Between 20jW and the continental slope, the typical southward AOU increase on isopycnals is
AOUy = 0.0275 Amol kg
1 km
1, with concurrent increases of nutrients due to oxidation of organic matter. The inference is subducted water moving south, perhaps meeting Antarctic Intermediate Water moving north. The oxygen utilization rate AOUt is estimated by Van Aken (2001) to be 9.6 Amol kg
1 year
1. AOUt/AOUy then gives a southward velocity
v c 11 mm/s. Diapycnal mixing is estimated to be a small contribution to AOUt (from KV f 10
5
10
4 m2 s
1, KVD2AOU/Dz2f
0.3 AM year
1). Water in the sub-surface salinity minimum off Galicia has probably been subducted in the northern Bay of Biscay. The AOU ageing gradient in the thermocline is also southward over the continental slope from Porcupine Bank to 40jN. This suggests either that annual mean flow in the thermocline is to

J.M. Huthnance et al. / Journal of Marine Systems 32 (2002) 107–137

the south or that the ageing trend over the slope is mainly determined by interaction with the anticyclonic gyre in the open ocean, despite the poleward slope currents in winter and occasional inferences of a permanent poleward flow in the lower thermocline. Potential vorticity on isopycnals has north – south gradients further north where it subducts into the thermocline, but is homogeneous west of Iberia where potential vorticity calculations suggest enhanced diapycnal mixing. However, other data, e.g., hydrographic sections during cruise CECIR XVII (March– April 1990), indicate an east – west front down to about 300 m at latitude 39 –40jN, implying convergence between waters to the south and to the north (Vitorino and Marreiros, in preparation). This front bends to the north near the shelf edge, east of about 10jW between 300 and 600 m, suggesting a tongue of northward moving water near the shelf edge as also simulated by Dubert (1998). Against the Portuguese slope, salinity (S) on the 27.1 isopycnal has an isolated maximum at 350 – 400 m depth (Van Aken, 2001). The source cannot be MSOW below the intervening f 450 m salinity minimum. Vertical mixing (from above) is implied with KV f 10
3 m2/s if Szz f 5.10
6 m
2, DS f 0.125, time scale f 1 year (from length scale 330 km/v f 0.011 m/s); larger KV is implied if lateral diffusion is effective in tending to remove the maximum. Further north, off Galicia and in the Bay of Biscay,

111

the highest salinities in isopycnals from the thermocline are also found over the upper slope. The high estimated KV and generic salt enrichment over the continental slope indicate that diapycnal mixing is quite important near the shelf. Salination of thermocline waters near the continental slope is by local downward mixing from higher surface values in the poleward slope current (Van Aken, 2001). Salinity in the upper seasonal thermocline varies seasonally because of the variable slope current; highest salinities are at the end of winter. On the more local scale of OMEX-II west of Galicia, average surface salinity in coastal water over the shelf is low due to river runoff, especially in winter. This causes a sub-surface salinity maximum over the shelf at about 100 m depth. In winter, stability due to the salt stratification suffices to allow a sub-surface temperature maximum over the shelf. Because the poleward slope current prevails in winter, advecting saline water, a lateral surface salinity maximum develops in winter near the upper slope between the less saline open-ocean water and fresher coastal water. Inter-annual variations observed in Eastern North Atlantic Central Water (ENACW; Fig. 2) are attributable to changes in the salinity and/or temperature of the subducting Central Water. The characteristic salinity of the ENACW envelope at 12 jC shows an interannual variation (Fig. 4), with a salinity maximum in 1992. In the preceding period, the characteristic sal-

Fig. 4. Characteristic salinity of the Eastern North Atlantic Central Water over the Galician slope (the OMEX-II area, black dots) and the Armorican slope (Bay of Biscay, SW off Brittany, open figures).

112

J.M. Huthnance et al. / Journal of Marine Systems 32 (2002) 107–137

inities were about 0.1 lower, while in 1997 –1999, characteristic salinities at the 12 jC isotherm were 0.05 lower, with variations of about 0.01. The latter changes are probably seasonal and were observed by regular hydrographic surveys in OMEX-II. These data off Galicia are well correlated with characteristic salinities of ENACW at the 11.5 jC isotherm from the Armorican slope in the Bay of Biscay (Fig. 4). Thus, ENACW variability appears to have a relatively large scale. Near Porcupine Bank, more influenced by the cyclonic sub-arctic gyre, the Central Water shows less variability; salinity variability is about one-third of that observed in the OMEX-II area, and not very well correlated. The cause of this inter-annual variability pattern is yet not clear; it may be due to either atmospheric forcing in winter or advection of Central Water from further west. 2.2. Currents and processes Maze´ et al. (1997) analysed Bord-Est hydrographic data to infer transports in the region. They used three boxes (latitude boundaries 36j, 37.2j, 40j, 43jN; coast to 12j or 13jW) and their combination, times six layer depths and six water properties (volume, mass, heat, salt, pre-formed nitrate and phosphate). Neglecting diapycnal exchange, these gave 144 constraints on 82 unknowns: 50 reference velocities underlying interior flow (assumed geostrophic), 24 net Ekman transports and eight surface fluxes. A least-square-error fit yielded 2 Sv depth-integrated transport eastwards across the western boundaries north of 37.2jN and northward meridional transport at all levels: 10 Sv at 37.2jN, 12 Sv at 43jN. In the narrower strip between 10.5jW and 1000 m water depth, the transports were 9 Sv at 37.2jN, 4 Sv at 43jN. Frouin et al. (1990) estimated a smaller baroclinic oceanic flow eastwards above 300 m between 38 and 42jN, increasing the northward transport in the winter slope current from f 0.3 to f 0.6 Sv at these latitudes. Their findings support the model of Stevens et al. (2000; under 1986 – 1988 mean monthly wind forcing) showing upper-slope poleward flow, separated in summer from poleward flow of Mediterranean Water by a flow minimum at f 500 m. Jorge da Silva (1996) estimated the poleward transport using November – December 1994 hydrographic data at six across-shelf sections (Table 1).

Table 1 Slope current transport at standard SEFOS sections (Jorge da Silva, 1996) using cruise SEFOS-IH-9401 hydrography Section

Position

Transport (Sv)

I II III IV V VI

37j40VN (zonal) 39j40VN (zonal) 41j30VN (zonal) 43jN (zonal) 43j30N 43j30VN/8jW (meridional)

5.7 5.1 3.0 4.3 2.1 2.9

Geostrophic currents relative to 1350 dbar were extrapolated to the upper slope and shelf after Reid and Mantyla (1976). Transport across each section was estimated for the upper 550 dbar, from the shelf edge to the outer limit of the slope current.

2.2.1. Mean and seasonal currents Current meter data have been analysed; historical data complement the measurements made during OMEX-II (Table 2). From these data, an atlas of mean monthly currents has been established (Fig. 5a, b). Moreover, f 100 mono- and multi-cycle RAFOS floats were deployed at depths 450 dbar (near the 450 m salinity minimum), 1000 and 1500 dbar in 40– 50jN during ARCANE. Surfacing each 3 months, multi-cycle floats gave data for typically 3 years each; mono-cycle floats had a 1 year underwater mission. These data show generally poleward flow over the slope, but also variability, some of which is seasonal, especially above the MSOW. In the upper water column ( < 600 m depth), winter currents (Fig. 5a) are poleward with a mean 0.03– 0.10 m/s. Satellite imagery also shows a warm tongue of water advecting north along the margin (Pingree and Le Cann, 1990, 1992a). North of 41jN, the 450 dbar ARCANE floats generally went to the west or northwest, but anticyclonic flow north of Cape Ortegal (44jN, 8jW) merged to flow generally to the southeast in Biscay. Drifting buoys drogued at 150 m in ARCANE were found to take an initial direction related to season: northwards in autumn –winter, westwards and southwards in spring –summer. Distinct seasonality was found by ARCANE moorings over the slope at 42j07VN off Vigo (Le Cann et al., 2002a, in preparation). Maximum poleward flow at all depths was in autumn/early winter: a first maximum between August and October in both ENACW and MSOW (monthly means f 0.05 – 0.1 m/s); a second surface-intensified maximum (monthly mean f 0.1 m/s) tending to occur in December – January. In

J.M. Huthnance et al. / Journal of Marine Systems 32 (2002) 107–137

113

Table 2 Current-meter measurements (in OMEX II and historical including SEFOS and OMEX I) contributing to flux estimates Latitude

Longitude

Water depth

CM depth

Start date

Duration (months)

Source

40.0873 40.102 40.198 40.283 37.832 37.802 37.771 37.753 42.267 42.218 42.218 40.999 41.003 41.016 41.036 38.217 38.408 41.07 41.005 41.098 42.645 42.64 42.64 41.318 41.37 41.315 42.1 39.752 39.752 39.783 39.788


9.843
10.125
10.46
11.00
9.509
9.726
10.235
11.00
10.15
9.803
9.509
9.475
9.748
10.254
10.908
9.788
9.763
8.807
9.058
9.337
9.697
10.028
10.028
8.982
9.305
8.987
9.387
9.61
9.617
9.808
9.81

1171 3838 4405 4927 1112 1990 3238 5050 2700 2337 1338 1293 2543 4085 3853 2000 2000 40 104 302 1453 2238 2238 84 2010 84 323 908 978 2300 2298

308, 713, 1001, 1021 277, 686, 996, 1472, 2999, 3639 243, 646, 1460, 2962 3030 320, 722 293, 990, 1484, 1795 298, 648, 1011, 1503, 1804, 2773, 3023 292, 680, 967, 1491, 3002 56, 256, 756, 1156, 2000 337, 837, 1237, 2037 404, 849, 1238 294, 794, 1194 843, 1243, 2243 885, 1285, 3085 653 1200 200, 700, 1200 36 42, 70, 97 38, 120, 170, 280 620, 1070 620, 1071 1075 27, 51, 71, 81 80, 130, 430, 1230 29, 53, 76, 82 5 97, 145, 639 95, 149, 359, 674 72, 120, 312, 615 46, 102, 290, 582

26/03/88 27/03/88 27/03/88 28/03/88 31/03/88 31/03/88 01/04/88 01/04/89 28/05/93 06/05/94 06/05/94 31/05/93 30/05/93 22/11/93 22/11/94 04/01/94 04/01/95 07/05/87 08/05/88 08/05/89 21/07/97 23/07/97 14/03/98 29/01/98 05/05/99 17/11/96 06/07/98 23/11/94 10/10/95 14/12/94 23/10/95

7 – 13 9 – 12 2 – 13 13 3 – 12 8 – 12 4 – 13 8 – 12 11 6.5 6.5 12 12 10 10 23 23 4.5 4.5 4.5 7.5 7.5 9.5 2.5 – 4 4 1 1 6 – 12 5 – 11 10 6 – 11

BORD-EST BORD-EST BORD-EST BORD-EST BORD-EST BORD-EST BORD-EST BORD-EST MORENA MORENA MORENA MORENA MORENA MORENA MORENA BODC-Pingree BODC-Pingree IH IH IH Ifm Kiel Ifm Kiel Ifm Kiel IH IH IH RAYO IH IH IH IH

February – March, flows at ENACW and MSOW levels tended to reverse (equatorward monthly means f 0.05 –0.10 m/s; a possible relaxation?). The poleward component of flow then increased until September. In phase this appears similar to the ‘‘SOMA’’ effect found on the northern Biscay slope by Pingree et al. (1999). However, a secondary minimum in June –July is linked to upwelling. This minimum is marked by equatorward reversals, observed from May to August (typically near the surface, 0.15 m/s equatorward, 10 days in duration). Increases in poleward flow are loosely correlated with weak ( f 0.02 – 0.04 m/s) increases in onshore flow (likewise decreases in each). Stronger episodes of off- and onshore flows ( f 0.1 m/s) occur on 5 –10-day scales. Surface offshore episodes tend to occur from May to September, apparently linked to upwelling events. Variance for

surface flows tends to be high in winter (poleward flow), with a secondary maximum in summer (upwelling events). Vitorino (1989) gives more evidence of summer upwelling conditions forcing near-surface southward mean flow, over the shelf and slope off northern Portugal, to as deep as 170 m when northerly winds were strongest during May – October 1987. Nevertheless, a poleward undercurrent can persist through summer, in this case, at 280 m depth during the whole period, only reversing briefly during the strongest upwelling wind events. Currents over the slope during May to September 1999 (OMEX; Fig. 6) were also predominantly poleward, especially in the record at 430 m depth. Further north, near Cape Ortegal, mean annual currents are surface-intensified and equatorward (westward) in the ENACW and MSOW ranges, and tend to

114

J.M. Huthnance et al. / Journal of Marine Systems 32 (2002) 107–137

be directed slightly offshore. Mean monthly currents are westward most of the year. Maxima are during summer for ENACW, with high variability. At the level of Mediterranean influence (600 –1400 m depth), currents are stronger, more persistent and less seasonal in character (Fig. 5(ii)). In particular, OMEX meters at f 620 and 1070 m (‘‘Ifm’’ in Table 2) showed mostly poleward mean flows O (0.1 m/s) in July 1997 to March 1998 (Fig. 7). March –December 1998 (an extended upwelling season) showed a different character; at 1070 m, there was more southward flow in late summer and early autumn. Over the steepest portion of the slope, currents are constrained along the mean slope direction; notably mooring 2 over the steeper slope in 1453 m water depth has little cross-

slope flow. The 1000 dbar ARCANE floats clearly showed a slope current going north at about 0.02 m/s. Higher values of stability in current direction are found in 600 –1400 m. A northward decrease of the poleward slope current is suggested by current meter analyses and Table 1. External pressure gradients associated with the oceanic density field are believed to favour the poleward flow, especially along the upper slope but also nearshore at the Algarve. A northward decrease in current is consistent with topographic guidance, including deflection around the Tagus Basin and Galicia Bank, and with hydrography (Coelho et al., 2002). The meridional gradient of density and inferred sea surface height gradient is concentrated further to the south (but Mo-

Fig. 5. Monthly mean current vectors for alternate months. (i) 0 – 600 m, (ii) 600 – 1400 m.

J.M. Huthnance et al. / Journal of Marine Systems 32 (2002) 107–137

115

Fig. 5 (continued).

roccan influence past Gibraltar is an open question), has a maximum in 39 – 41jN, almost disappears (as forcing for slope current) north of 41jN and increases again in Biscay (Arhan et al., 1994). Coelho et al. (2002) find that model agreement with observed currents is much better if actual (detailed) density observations are used rather than Levitus climatology. This density forcing is seasonal; weaker in summer. 2.2.2. Variability ‘‘Eddy’’ kinetic energy was typically 0.004 m2 s
2 for floats at 450 dbar; more north of Galicia. Current meters in 600 –1400 m often show relatively large low-frequency variance. Variance ellipse axes are generally aligned along the topography above and

below 600 m. Some across-slope variability below 600 m along the southern Iberian Margin is possibly due to off-slope flow of Mediterranean Water. The 1000 dbar floats show slope –ocean exchange, as do current meters in 600– 1400 m, 38 –39jN (Fig. 5(ii)). Further west, the floats show westward flow at 39 – 40jN, 42 – 43jN and eastward flow at 46 – 47jN, indicating possible bifurcation around Galicia Bank. At 1500 dbar (floats), flow was mostly to the west in 39– 43jN, with a southward component further north. Spreading of Mediterranean Water was successfully modelled by Stevens et al. (2000). Evidence of Mediterranean Water eddies (Meddies) shows in the enhanced low-frequency variance around 1000 m depth, and in ARCANE CTD and

116

J.M. Huthnance et al. / Journal of Marine Systems 32 (2002) 107–137

Paillet et al. (2002, in press) and Le Cann et al. (2002b, in preparation). OMEX spectra show peaks at inertial frequencies (notably at the 2338 m mooring near 42j40VN) and decrease with depth down the water column. There is some variability at 2 – 3-day periods and some smaller magnitude variability at 10 – 25 days. 2.2.3. Tidal currents and mixing The assembled data were analysed for tidal currents; model results in Rodriguez et al. (1991) were also used. Tidal currents in the area are predominantly semi-diurnal with typical amplitudes 0.1 m/s on the

Fig. 6. Stick plots for daily mean vector currents near 41j34VN, 9j18VW at f 80, 130, 430, 1230 m in water depth 2010 m.

ADCP sections in Biscay and off Iberia. In one case, the signal extended up to the surface and down to more than 1500 m; the Meddy had perhaps come from the slope in winter (Paillet et al., 1999). Eddies tracked by floats moved in an overall south – westward direction northwest of Finisterre. About four to five Meddies may exist at the same time in the Iberian Margin region. Some appear to come from the straight slope near 41jN. The southern Portuguese capes also appear to be sources of Meddies: e.g., Pingree (1995) tracked one, finding anticyclonic circulation up to 0.3 m/s; Bower et al. (1997) tracked floats deployed in MSOW, finding Cape Saint Vincent and Estremadura Promontory to be regular sources of Meddies; Cherubin et al. (2000) detailed mesoscale variability from baroclinic instability – filaments, small eddies and Meddies near Cape Saint Vincent and Cabo da Roca. Eddies and filaments observed in ARCANE, especially by floats and drogues, are further discussed in

Fig. 7. Stick plots for daily mean vector currents near 42j40VN at f 620 and 1070 m in water depths 1453 m (mooring 2) and 2338 m (mooring 3).

J.M. Huthnance et al. / Journal of Marine Systems 32 (2002) 107–137

shelf, 0.03 – 0.05 m/s in 2000– 4000 m depth where they assume greater relative importance as other forms of current are weaker. Amplitudes increase around and north of Finisterre. Deep-water ellipses are alongslope but gain an increased cross-slope component on the upper slope and outer shelf, especially close to the bottom. One mooring in 1293 m at 41jN has relatively strong cross-slope diurnal currents f 0.05 m/s. Spectra show peaks at the principal M2 tidal frequency and its harmonics M4, M6. Depth-variability, strong non-linearity, modulated forms (perhaps from non-linear rectification or varying stratification) or local enhancement are often apparent. Examples of bottom-intensification are near the 2338 m OMEX mooring, with tidal currents as large as 0.2 m/s, in 1200 m at 41j22VN and at mid-shelf (86 m) inshore of Porto Canyon (Vitorino et al., 2002a). Here, tidal currents have an M2 major axis f 0.03 m/s and are strongly polarised, roughly aligned with the canyon axis. An internal tide appears in Nazare´ canyon as a propagating wave with (potential) energy density decreasing towards the sides and head of the canyon (Quaresma et al., 2000), while in Setubal canyon, it appears as a standing wave. Model runs for a section at 42j40.5VN (Davies et al., 2002) show internal tides enhancing currents on the uppermost slope, and further enhancement by upwelling-induced stratification. Tidal cycles with overturning/enhanced mixing and currents near the bottom, up to 300 m above the bed, were shown by ARCANE 27 h yoyo CTDs and current moorings in 1700 m near 42j6VN off Vigo. Semi-diurnal currents were as large as 0.25 m/s with phase propagating down (energy spreading up and down from near 1000 m). Possibly tidal vertical velocities f 0.01 –0.02 m/s and inversions were also measured. Vertical diffusivity, KV, was roughly estimated assuming that the Ozmidov (1965) scale, LO f e1/2N – 3/2, was approximately equal to the Thorpe (1977) scale LT of density inversions, which could be estimated from re-ordered profiles (see e.g., Ferron et al., 1998). Then from the steady turbulence kinetic energy equation

production P ¼ buoyancy loss B þ dissipation e, B ¼ N 2 KV ,

Rf ¼ B=P

117

we have B = Ce, where C = R f /(1-R f ) = 0.2; then KV = Ce/N2 = CNLT2. LT is large near the surface, has another range of larger values in 100– 400 m and is especially large in 1200– 1600 m. Hence, KV = O(10
2 m2 s
1) near the bottom, O(10
3 m2 s
1) at 100 –400 m. The internal wave regime and associated turbulent dissipation on the shelf near Vigo have been studied by Sherwin et al. (2002). Many 5– 10-m internal waves (i.e., range up to 20 m) are spread throughout the tidal period, especially in the troughs of the 12.4-h M2 internal tide, which has amplitude up to 20 m. Measurements in August 1998 showed solitons at 15 –20min intervals most of the time. SAR images also show a fairly chaotic internal wave field. The overturning length scale was about 5 m. In the band between 13.5 and 15.5 jC, the vertical diffusion coefficient, KV, was measured to be about 3.7  10
4 m2 s
1 (tidal average); the associated energy dissipation f 0.003 Wm
3 was the greatest found in OMEX measurements. By contrast, the shelf edge at 41jN studied in MORENA had packets of very large non-linear internal waves phase-locked to the barotropic tide (Jeans and Sherwin, 2001). Maximal vertical diffusion was estimated in the thermocline; KV f 2.2  10
3 m2 s
1 averaged over a tidal cycle. At the base of the thermocline, values were an order of magnitude smaller. We note an analogy with calm conditions in the North Sea; Van Haren et al. (1999) found average KV f 3  10
5 m2 s
1 but a peak up to 6  10
4 m2 s
1 in the thermocline where the buoyancy frequency N was greatest ( > 0.025 s
1); near-inertial motion caused shear and combined with the tide to form the internal wave spectrum. The greatest turbulent vertical diffusivities found in OMEX, KV of order 0.01 m2 s
1, were during winter storms when mixing extended to 100 m depth, but values were less than 10
4 m2 s
1 below a fresher surface layer over the shelf. In filaments, maximum measured diffusivities were about 1.7  10
5 m2 s
1 with a mean 1.3  10
5 m2 s
1 (relatively small values in the context of only small internal waves; Sherwin et al., 2002). (Note: convergence at filament boundaries can cause vertical extension of water properties beyond that caused by diffusion).

118

J.M. Huthnance et al. / Journal of Marine Systems 32 (2002) 107–137

Table 3 1981 – 1999 upwelling index at 42jN, 9jW, m3 s
1(100 m)
1 Month

January

February

March

April

May

June

July

August

September

October

November

December

Mean smd sdd sd6h mabsd sumup


31 38 113 125 86 297


15 32 70 78 51 596

7 13 69 76 45 648

13 16 60 68 43 806

17 26 52 59 39 613

38 18 41 47 42 1121

58 27 36 43 52 1562

41 21 38 44 42 1192

14 12 46 52 34 704


3 16 68 74 46 404


22 30 76 85 54 428


30 40 99 115 74 417

Data from NOAA Pacific Fisheries Environmental Laboratory, derived from 6-hourly Fleet Numerical meteorological analyses. Divide by 100 for values in m2 s
1. ‘‘smd’’ is a standard deviation for the individual monthly values about the 19-year mean for the month. ‘‘sdd’’ is a standard deviation for daily values relative to their mean for the individual month and year; ‘‘sd6h’’ is the equivalent for 6-hourly values; the years 1993 – 1999 were used to estimate these values. ‘‘mabsd’’ is the mean of the absolute daily values in 1993 – 1999 (by month). ‘‘sumup’’ is the 1993 – 1999 mean of the sum of positive daily values for the month.

2.3. Summer regime Above 600 m, poleward flow is a minimum in spring and summer, and a maximum in the autumn (Section 2.2). If or when strong enough, southward (upwelling-favourable) wind stress forces southward flow overlying and countering the poleward flow. Satellite images in summer show strong evidence of upwelling along the Iberian shelf, enhanced at Finisterre (43jN; starts earlier, more intense and more persistent). Different phases of upwelling should be distinguished. From the onset of upwelling in May or June, the seasonal thermocline is initially displaced, later surfaces, moves offshore and becomes established at the shelf break. Secondary upwelling fronts

can form inshore of the established front. Where the Rias Baixas break the coastline, the corresponding portion of upwelling appears to occur within the rias ´ lvarez-Salgado et al., 2000). Associated filaments (A extending offshore begin to develop in July or August and grow to lengths of order 200 km by September (Barton et al., 2001). Cape St. Vincent is a particular location for filaments and may favour flow separation from the shelf. Filaments quickly disappear when upwelling-favourable winds cease (October on average; Table 3). Monthly average values of the upwelling index (i.e., the flux s/qf) for the Vigo region are shown in Fig. 8 and Table 3. There is a clear annual cycle. However, inter-annual variability is large: the summer

Fig. 8. 1981 – 1999 upwelling at 42jN, 9jW, m3 s
1(100 m)
1. Data from NOAA Pacific Fisheries Environmental Laboratory. The range is for monthly values during 19 years.

J.M. Huthnance et al. / Journal of Marine Systems 32 (2002) 107–137

integral of s/qf (summing contiguous positive monthly values) ranges from 119 in 1999 to 273 in 1998; only in June to September is the standard deviation less than the index magnitude to make the sign of the annual cycle probable (corresponding to upwelling in these months). In May 1999, and as early as March to May in 1998, currents showed strong upwelling events (up to 0.35 m/s, Fig. 9; see also Vitorino et al., 2002a). These events occurred during typical winter conditions, when the mid-shelf was nearly homogeneous (except for surface stratification due to riverine outflow). They then have an important role in promoting re-stratification of mid-shelf waters near the bed, by means of onshore transport of cold water in the bottom layer. Current meters deployed in May 1987 showed that in response to sustained northerly winds established that month, an equatorward flow appeared first over the shelf and rapidly migrated to the upper slope region (Vitorino, 1989). On the other hand, 1997 upwelling only started in late June. The ends of the upwelling season are very variable; for example, in 1998, upwelling conditions were sustained by predominantly northerly winds until early December. Upwelling breakdown is poorly observed but may be rapid, related to change of wind regime and perhaps seasonal deepening of the thermocline. Summer upwelling was successfully modelled by Stevens et al. (2000), also showing enhanced upwelling off Cape Finisterre and Cabo da Roca. Detailed measurements of a filament were made in August 1998 when coastal upwelling occurred consistently; surface temperatures were < 15 jC near the coast, f 20 jC offshore (Barton et al., 2001). Initially, the wind weakened, the area of a ‘‘decaying’’ filament decreased and drifters deployed near the shelf break converged slowly towards the filament’s southern boundary, characterised by a strong temperature front which sharpened; sinking velocities f 16 m/day were estimated from the convergence. After a return to upwelling-favourable northerly winds, the drifters moved offshore. One drifter went far offshore; the others recirculated cyclonically, completing a circle in about 25 days. Speeds were typically 0.05 – 0.25 m/s. Only one drifter ultimately went far, going south in the autumn (as northerly winds persisted until December 1998). Cross-sectional currents (ADCP) and hydrography showed generally weak offshore flow limited to a thin surface layer and a double cold core,

119

the filament having resulted from a merger of two. The offshore flow was strongest, 0.2 – 0.3 m/s, and deepest at the northern boundary. The centre of the filament below 25 m corresponded to upwelled oceanic water moving slowly towards the coast, with onshore transport about 0.1 Sv (not reliable or necessarily typical). Hence, the filament may be essentially a surface phenomenon (with satellite images and surface drifters giving an indication of the flow only in a superficial layer). Later in August 1998, a developing filament was observed, in a freshening wind. Currents were f 0.3 m/s above 50 m depth; the filament did not have shoreward flow. Both of these filaments had about 0.6 Sv offshore transport (considered a coincidence), more than just entrainment of the upwelling Ekman transport s/qf from the corresponding sector of shelf. There are other aspects of the form of upwelling. One MORENA CTD section at 10.2jW showed (filament?) salinity maxima at 41jN and 41.8jN near 100 m. Stronger seasonal stratification (larger deformation radius) may imply more work to raise the thermocline to the surface — also affected by thermocline depth and friction. A shallower thermocline implies strong current response to winds. Along-shelf flow to the south will tend to intensify until energy and momentum are lost (to filaments, eddies or friction). Depths from which water upwells are indicated by the temperature of upwelled water. Bottom temperatures at 154 m depth in 1997 (Fig. 10) show a correlation
0.6 with upwelling index at a lag of 2 –3 days; temperature minima lag peaks in upwelling index. Values, typically 12.5 jC during upwelling, indicate an origin from about 210 – 230 m depth. This may be the greatest depth for upwelling water; the bottom water at 154 m does not necessarily reach the surface. Other sub-inertial flow at mid-shelf off northern Portugal, May – October 1987, was characterised by near-barotropic fluctuations. Energetic current-spectra bands were at 4 –6 and 11 –15 days, as in wind stress spectra, which also had a peak around 3 days. For May to mid September 1987, with sustained upwelling-favourable winds, the 4– 6-day (anticyclonic) current fluctuations were coherent with the wind stress (Vitorino and Coelho, 1998). The 4 – 6-day windstress fluctuations can excite the first two coastaltrapped wave modes (phase speeds 4.5 –5, 2.2 m/s);

120

J.M. Huthnance et al. / Journal of Marine Systems 32 (2002) 107–137

J.M. Huthnance et al. / Journal of Marine Systems 32 (2002) 107–137

Fig. 10. (a) Upwelling index and (b) sea-bed temperature from an ADCP at 42j 40VN, 154 m depth, June – December 1997.

coherence analysis suggested that the 4 –6-day current fluctuations may have corresponded to a second-mode coastal-trapped wave excited by the local wind. 2.4. Winter regime There is generally a northward warm flow along the upper slope (Section 2.2), sometimes 0.15 m/s or more; its establishment has been observed with a ‘‘nose’’ going north at this speed (Haynes and Barton, 1990). Drifters entrained in September 1986 in the warm tongue travelled north at 0.25 m/s over a period of a month or more. Current meters moored 750 m over the slope registered northward flow between 0.10 and 0.40 m/s at depths from 50 to 600 m during a period of calm wind. Simultaneous observations showed a high salinity anomaly centred near 70 m depth moving northwards in association with the warm flow at 0.27 m/s.

121

Surface drifters released near the shelf edge in ARCANE tended to remain trapped in this region in winter. However, the flow may have a branch around Galicia Bank as suggested by float tracks. The surface temperature field (in remote-sensed images) often shows a much broader (hundreds of kilometers) northward extension of warm water, probably affected by winds. The external baroclinic pressure field is possibly forcing the northward flow as a decaying branch of the Azores front approaches the continental slope near Cabo da Roca. Although winds favour downwelling on average, they are most variable in winter (Table 3) and always favour upwelling at some time in any month; conditions may switch from down- to upwelling and vice versa during the month. (Consequently, there is a complex transition from summer upwelling to the winter regime). Current records on the shelf in February to May 1998 show that for most of the time the wind was from the south driving a relatively small northward current (Fig. 9). However, when the wind reversed at times in March to May 1998, there was a strong southward flow with speeds O(0.3 m/s), upwelling towards the shore at the bottom and offshore flow in the surface waters. These data show transition from winter to summer; mid-shelf current measurements off northern Portugal during November 1996 to January 1997 show transition from upwelling to winter; both sets are discussed by Vitorino et al. (2002a). Conclusions for the winter regime (Vitorino et al., 2002a) are that the mid-shelf water column off northern Portugal is nearly homogenous, although OMEX cruise, RRS Charles Darwin 110, January 1998, recorded a remaining thermocline from summer 1997. Stratification occurs near the surface in cooler, fresher river plumes (especially the Douro plume). Northerly winds can displace the Douro plume offshore; southerly winds trap the Douro plume at the coast. Near-bottom stratification occurs during strong upwelling events. Downwelling-favourable winds are dominant and lead to sustained periods (1– 2 months) of poleward flow over the mid-shelf (0.1 –0.2 m/s).

Fig. 9. (a) Coastal wind stress at Porto (41jN from Portuguese Meteorological Office), (b) coastal sea level adjusted for atmospheric pressure, Viana do Castelo (41j41VN), (c) temperature and currents at depths 26, 51, 71, 81 m in water depth 86 m at 41j19VN, 8j59VW, (d) nephelometry at 81 m, same location, (e) wave shear velocity from wave buoy data at the same location: (a) to (c) are low-pass filtered (cut-off period 28 h); year is 1998.

122

J.M. Huthnance et al. / Journal of Marine Systems 32 (2002) 107–137

Strong mid-shelf sub-inertial currents (to 0.3 m/s) respond within about 12 h to both northerly and southerly wind-stress events, and comprise: a surface Ekman layer extending down to f 25 –30 m during the strongest wind-stress events; inviscid interior flow aligned north – south along the local topography; a bottom Ekman layer thicker than 15 m during periods of strong interior flow. Several energetic bands are present in both the wind stress and current spectra for the November –January and the February –May observations, notably at periods 3– 6 days for which coherence reveals wind-forcing of the currents. However, the along-shore momentum balance suggests that local wind-stress forcing alone is not able to explain the along-shore flow evolution. Other mechanisms such as incident remotely forced waves must also play a role. Regarding higher frequencies and sediment suspension, winds also cause near-inertial oscillations; in total, wind events can give currents >0.4 m/s and/or long waves sufficient for sediment resuspension on the shelf in depths to 200 m (Huthnance et al., 2002; Vitorino et al., 2002b). Energetic wave conditions are common along the northern Portuguese coast in winter; resuspension criteria for the fine sediments at mid-shelf are frequently exceeded. During major storms, waves also lead to development of a bottom nepheloid layer 10– 30 m thick. Resuspension criteria are also exceeded during periods of strong sub-inertial flow, such as the upwelling event in mid-March 1998. Measurements of particulates around Setubal canyon in February 1998 suggest upwelling flow onto the shelf at the head, possibly aided by a standing internal tide (TRANSCAN, 1999), but denser shelf water flow into the canyon around the sides was indicated by pCO2 levels.

Fig. 11. Budget scheme.

shore; i = 1, 2, 3, 4). We assume a steady state. If qS is the inf low flux at the southern end, and otherwise qN, qi are all respective outward fluxes, then water conservation gives qS ¼ qN þ Ri qi

ð1Þ

Let CS, CN and Ci be the respective concentrations (Ci is for inf lows to the box from adjacent waters or the coast), Cˆ u (CS + CN)/2 the average concentration along the length for outf lows, qiV the exchange transport across each side of the box (equal amounts in and out). Then constituent conservation gives qS CS ¼ qN CN þ Ri qi ðCi for inf lows; Cˆ for outf lowsÞ þ Ri qi VðCˆ
Ci Þ

ð2Þ

2.(2)
(CS + CN).(1) is ˆ ðCS
CN ÞqS ¼ qN ðCN
CS Þ þ 2Rinf lows qi ðCi
CÞ

3. Fluxes and budgets As a check on the validity of available data for testing models, we now consider the budgeting process for flux estimates in a box, followed by discussion of the choice of boxes. 3.1. Basis of calculation Suppose a box (Fig. 11) has south (S) and north (N) ends and four sides (top, bottom, inshore, off-

þ 2Ri qi VðCˆ
Ci Þ i.e., ðCS
CN ÞðqS þ qN Þ=2 ¼ Rinf lows Aqi AðCˆ
Ci Þ þ ½Ri qi VCi V u
Ri LWi Ki @C=@ni 

ð3Þ

where CiV u Cˆ
Ci, L is the box length, Wi the width/ depth of the side i, Ki the diffusion coefficient across

J.M. Huthnance et al. / Journal of Marine Systems 32 (2002) 107–137

123

side i and ni the outward normal to side i. Eq. (3) shows that because fluxes of water add to zero, an arbitrary change of ‘‘zero’’ for constituent concentrations has no effect; only concentration differences are significant. Estimates of these terms thus depend on (a) measured mean concentration differences for the chosen constituents, and (b) averaged statistics for the measured currents from all sources. All the information on currents can be ‘‘pooled’’ to estimate the mean and exchange flows. We also need a relation between the exchange velocities making up qiV and the diffusion coefficient Ki, viz., an offshore scale for the concentration difference (Section 4.3) or a direct estimate of lateral Ki from floats and drogues (Section 5.2).

determined concentration differences. The north and south boundaries should ideally enclose but avoid filaments. The intensity of OMEX-II CTD stations in 42– 43jN suggests that this may be an appropriate range. We partition the cross-section laterally at the 200 m depth contour and 10j20VW to give shelf, slope and ocean sectors. No budgets were attempted in the ocean west of 10j20VW or below the partition at 1600 or 1900 m, for lack of sufficient hydrography and current measurements.

3.2. Choice of boxes

4.1. Measured mean concentration differences

We do not have the data to calculate fluxes qiVCiV as detailed area- and time-integrals. Instead, we consider the appropriate time and space scales (box sides) over which to take the averaged statistics required by Section 3.1. The scales must be large enough for sufficient data to give confidence in the averages. On the other hand, we have to respect the obvious difference between summer upwelling and winter downwelling regimes. The biogeochemical interests of OMEX also emphasise the difference between regimes on the shelf, the slope and further offshore. Partitions at levels 100 m (salinity maximum), 450 m (salinity minimum) and 1000 m (MSOW salinity maximum), with an additional layer 1000– 1600 m, have DS/Dz = 0, minimising vertical exchange of salinity. This simplifies flux budgeting, inference of lateral exchange being separated from uncertainties in vertical exchange, but has disadvantages too: no estimates of vertical exchange coefficients; boxes do not identify with water types. For example, MSOW is divided between boxes above and below its salinity maximum; such partitioning fails to capture its northward decrease of salinity in one box-layer. An ‘‘alternative’’ scheme centres boxes around the main water types, with bounding levels at 100, 250, 650, 1400, 1900 m. Then the box in 650– 1400 m captures the MSOW, and salinity gradients are maxima at 250, 650, 1400 m. Along the slope, the boxes need not fully span the utilised current data, which are averaged along the slope in any case, but should be chosen for well-

We consider the constituents salinity and oxygen; nutrient budgeting is doubtful (lacking number and accuracy of data). Salinity and oxygen data in 42 – 43jN, east of 11jW, were sought from OMEX, ‘‘Galicia’’ and MORENA, further requiring that the hydrographic stations within one cruise were evenly spread over the budgeting area and covered the continental shelf, slope and nearby abyssal plain. Only four data sets qualified: OMEX, 20 June to 20 August 1997 and 24 December 1997 to 15 March 1998; ‘‘Galicia,’’ September 1986; MORENA I, May 1993. If data were restricted to any one set or year, the resolution discussed in Section 3.2 would leave gaps (notably currents near the surface and over the shelf in summer). Therefore, we have ‘‘pooled’’ these data to estimate the budget terms (Eq. (3)). Average concentration differences from OMEX (June – August 1997) and ‘‘Galicia’’ cruises represent Summer and averages of MORENA I and OMEX (December 1997 to March 1998) represent Winter. For meridional gradients (only), extra weight (  3) was given to the MORENA I data in proportion to their extra extent. Inter-annual variability raises the issue of consistency of data between years, addressed by error estimates derived from variability of the ‘‘pooled’’ data, noting that only spatial differences in constituents are relevant. The resulting seasonal values are climatological, for no particular year. Characteristic values for boxes and their interfaces were determined by multiple linear regression in space (with respect to x, y, z). The estimates’ accuracy

4. Methods and budget calculation

124

J.M. Huthnance et al. / Journal of Marine Systems 32 (2002) 107–137

depends on the number and distribution of measurements, variability on the scale of the boxes and smaller, and applicability of the linear regression. We estimate accuracy of order 0.02 for salinity. MORENA data are more extensive north – south so that meridional gradients are more reliable. When comparing summer 1997 with winter 1997/ 1998 (for the ‘‘alternative’’ boxes, Fig. 12), the most striking differences are found in the upper 250 m. As expected, the upper 100 m is colder in winter than in

summer. The freshest water in this surface layer is found over the continental shelf, with lowest values in winter, probably due to seasonal river runoff. Further offshore, over the slope and beyond, the surface salinity is highest in winter. In 100– 250 m, salinity and temperature are highest in winter, owing to the winter poleward flow. Below 250 m in summer, the highest temperatures and salinities are against the continental slope; in winter, this on-/offshore gradient is weakened or even reversed.

Fig. 12. Distribution of mean temperature and salinity in boxes in the OMEX-II area between 42jN and 43jN for the summer of 1997 and the winter of 1997/1998. The boxes between 0 and 100 m are characteristic for the surface water and the boxes between 100 and 200 m contain Central Water from the upper part of the permanent thermocline. The boxes between 250 and 650 m, 650 and 1400 m and 1400 and 1900 m, respectively, represent the layer with the sub-surface salinity minimum, the MSOW layer and the LSW layer.

J.M. Huthnance et al. / Journal of Marine Systems 32 (2002) 107–137

125

distance (104/90 km) from 42j to 43jN. An average of all mean along-slope velocity estimates in the box for the corresponding season (i.e., summer or winter) is multiplied by the cross-slope sectional area of the box (Table 4). Rinf lowsAqiA(Cˆ
Ci). Precipitation –evaporation is estimated over the sea as 0.3 m/year (Times, 1992), which we attribute to 6 winter months (November to April only) and apply to the top shelf and slope boxes. River runoff (to the top shelf box only) is estimated from known monthly mean values for the Douro from ´ gua, 2000). These the years 1985 –1993 (Instituto da A are scaled up by two estimated factors: (i) 1.93 to add the Ave, Ca´vado, Lima and Minho inputs — the factor is based on comparison with Douro discharges for the common periods of available data in 1985 – 1990 ´ gua, 2000); (ii) 14.7/13.9 for the ratio (Instituto da A (total catchment feeding 40 –43jN/combined catchment for rivers with data) in 1  1j area units from Times (1992). A third of this runoff is then attributed to 42– 43jN via a coastal plume; November to April is regarded as winter, May to October as summer. An alternative estimate of runoff, based on the Rias

4.2. Averaged statistics for measured currents To form the best estimates of along-slope mean current and cross-slope current variance for Eq. (3), data as in Section 2.2 were used (Table 2). Monthly mean currents, stability (the ratio of vector to arithmetic mean speeds) and low-frequency variance (ellipses) were derived. The along-slope direction is defined as the major axis of the variance ellipse after removing tides by low-pass filtering ( f 30 h Butterworth twopass). (Many of the records have gaps and other problems. In ARCANE, the records at 1000 m near 42j07VN are the shortest, f 170 days, and values correspondingly less reliable. Here, the along-slope direction was taken as N10jW; a seventh-order Butterworth filter, cut-off frequency at 1/40 h was used.) 4.3. Budget calculation Estimates of the terms of Eq. (3) were calculated as follows. (CS
CN)( qS + qN)/2. The south-to-north concentration gradient as in Section 4.1 is multiplied by the

Table 4 Box and current statistics (from all available current meter data, by box) for budgeting Cross-slope

Summer

Winter

Area

Max width

hvi

AuVA

No.

hvi

AuVA

No.

Box (original) Shelf, 0 – 100 m Shelf, 100 – 200 m Slope, 0 – 100 m Slope, 100 – 450 m Slope, 450 – 1000 m Slope, 1000 – 1600 m width at base

2.66 0.66 6.83 23.28 34.24 33.20

33.7 19.4 68.3 68.3 64.5 59.8 50.3


24
10 3
6 21 17

17

5 1 9 37 35 31

41 27 43 17 10 29

15 3 20 19 27 25

7 5 15 26 31 25

Box (alternative) Shelf, 0 – 100 m Shelf, 100 – 200 m Slope, 0 – 100 m Slope, 100 – 250 m Slope, 250 – 650 m Slope, 650 – 1400 m Slope, 1400 – 1900 m width at base

2.66 0.66 6.83 10.21 25.79 44.33 24.33

33.7 19.4 68.3 68.3 67.1 63.0 53.9 43.2


24
10 3
21 5 18 0

5 1 9 13 31 58 7

41 27 43 38 1 21 27

15 3 19 20 12 29 21

7 6 15 15 23 40 3

13 18 18 15

17 13 14 20 16 5

Box dimensions in kilometers, currents in millimeters per second; notation as in Section 3.1. ‘‘No.’’ is the number of current series used; alongslope box length is 1j or 104/90 km.

126

J.M. Huthnance et al. / Journal of Marine Systems 32 (2002) 107–137

´ lvarezBaixas drainage basin only, is given by A Salgado et al. (submitted for publication) as 204 m3 s
1 for mid-May to mid-October. Cˆ is 35.8 and Ci is 0 for salinity; for oxygen, Cˆ and Ci are supposed to be essentially equal for the small inputs of water. We have neglected other net inflows to the boxes, assuming boxes extensive enough to have compensating outflows through the same face. RiqiVCiV u
RiLWiKiDC/Dni. The left-hand expression is used for the lateral faces of a box. An average of all cross-slope velocity variability estimates in boxes adjacent to the face (Table 4) is multiplied by the area of the face. The right-hand expression could have been used with observational estimates of effective lateral diffusivity: Ki = K. The ratio of estimates from the two approaches is K/ (lAuVA), where DC/Dni is evaluated as a difference CiV over an offshore distance l, i.e., half the box width. For typical K = 560 m2 s
1 (Section 5.2), l = 26 km and AuVA = 19 mm/s (Table 4), the ratio is 1.13, suggesting encouraging consistency. For the top and bottom faces of a box, the expression
RiLWiKiDC/Dni is used; there are no estimates of vertical velocity variability ( qiV). Where C is consistently a maximum or minimum at the face (e.g., salinity at depths 100, 450, 1000 m for the ‘‘original’’ boxes, Section 3.2), zero flux is imposed. Otherwise, box dimensions (Table 4) are used together with a large value 10
3 m2 s
1 for KV (Sections 2.1 and 2.2); DC/Dni is estimated from 2 [C(face)
C (interior)]/(box depth). At faces between boxes ‘‘A’’, ‘‘B’’, fluxes so calculated may not be equal and opposite between the boxes as required to conserve salt or oxygen. Hence, a ‘‘balanced’’ flux was derived: balanced flux ðA to BÞ ¼ ½original flux ðA to BÞ
flux ðB to AÞ=2:

ð4Þ

For salinity, a ‘‘mean’’ calculation was also carried out; water volume fluxes and salinity differences were each estimated as the mean of summer and winter values. Salinity differences for this ‘‘mean’’ calculation over the slope below 250 m were generally consistent (within a factor 2) with differences derived from the wider-ranging hydrography in Van Aken (2000b, 2001).

4.4. Calculated fluxes for budget Results are shown in Table 5. For each box, a closed budget would show ‘‘S –N’’ equal to ‘‘sum other.’’ All cases (summer/winter/‘‘mean’’; salinity/ oxygen; original/alternative boxes) have large discrepancies, which are not systematically improved for the ‘‘balanced’’ flux estimates. However, comparisons between the flux estimates are instructive. Consider initially the ‘‘mean’’ salinity with the alternative boxes. In this scheme, the lower three ‘‘slope’’ boxes do not share lateral faces with any other box; the inshore face is the sea floor. Therefore, only the vertical fluxes are ‘‘balanced.’’ The ‘‘balanced’’ values differ little from the ‘‘sum other’’ values, despite large vertical differences in salinity through the Mediterranean water. This is because vertical fluxes make relatively small contributions to ‘‘sum other,’’ despite using a large vertical diffusivity, 10
3 m2 s
1. Estimated precipitation– evaporation over the sea is also small in the salinity budget: 1.3, 2.6 m3 s
1 respectively to the top shelf and slope boxes. However, river runoff as estimated is important to the top shelf (coastal) box, contributing 4.9, 14.9, 9.9 m3 s
1 in summer, winter, ‘‘mean,’’ respectively; the alternative ´ lvarez-Salgado et al. (submitted for pubestimate of A lication) contributes 7.2 m3 s
1 in summer. Variability and uncertain attribution of runoff to 42– 43jN may be sources of discrepancies in this coastal box (only). The comparability of the independent estimates K and lAuVA of the lateral exchange coefficient (Section 4.3) suggests that this aspect of the lateral fluxes is reasonably well estimated. This leaves the estimates of throughbox transport and horizontal gradients of salinity as the likely causes of large discrepancies. At the MSOW level 650– 1400 m, these are in effect the terms in the calculations of Daniault et al. (1994) and Huthnance (1995), which are compared with the present ‘‘mean’’ salinity figures in Table 6. The present values of both transport q and north – south salinity difference are relatively small and more than account for the budget discrepancies; much smaller q and salinity difference at the adjacent levels above and below MSOW exaggerate the discrepancy. However, there is reasonable comparability for the top two slope boxes. In comparison with summer, winter salinity fluxes show: extra northwards flux above 250 m over the shelf and slope, contributed both by water volume

J.M. Huthnance et al. / Journal of Marine Systems 32 (2002) 107–137

127

Table 5 Estimated flux contributions to budgets in boxes, (a) salinity, (b) oxygen (a) Salinity fluxes

Summer

Winter

S–N (m** 3/s)

Sum other (m** 3/s)

Balanced (m** 3/s)

Box (original) Shelf, 0 – 100 m Shelf, 100 – 200 m Slope, 0 – 100 m Slope, 100 – 450 m Slope, 450 – 1000 m Slope, 1000 – 1600 m


1.9
0.4 0.5
3.9
1.4
14.7


6.6 1.6
4.8
11.4 34.1 99.4


1.9 5.8
0.2
7.2 34.1 99.4

Box (alternative) Shelf, 0 – 100 m Shelf, 100 – 200 m Slope, 0 – 100 m Slope, 100 – 250 m Slope, 250 – 650 m Slope, 650 – 1400 m Slope, 1400 – 1900 m


1.9
0.4 0.5
8.6 1.5 14.4 0


6.6 1.6
4.8 8.6 15.1 64.1 13.6


1.9 0.4
0.2 3.2 11.8 69.3 18

(b) Oxygen fluxes

Summer S–N (kmol/s)

Mean

S–N (m** 3/s)

Sum other (m** 3/s)

Balanced (m** 3/s)

S–N (m** 3/s)

Sum other (m** 3/s)

Balanced (m** 3/s)

11.1 1 18.8 11.5 4.8 57.8

1.1 0 18.2
1 148.5 58.4

1.7 14.5 22.3 17.7 148.5 58.4

1.5 0.3 6.8 3.6 3.2 13


2.7 0.8 5.2
6.1 82.6 89.2

0 9.9 9.9 5 82.5 89.2

11.1 1 18.8 17.8 0.3 46.5 28.2

1.1 0 17.6 2.2 27.1 129.5 79.6

1.7 4.9 22.3 8.2 23 134.1 84.8

1.5 0.3 6.8 3.7 0.9 29.4 5.3


2.4 0.8 5.2 4.6 24.7 95.9 44.4

0.2 2.6 10.1 4.7 21 100.8 49.1

Winter Sum other (kmol/s)

Balanced (kmol/s)

S–N (kmol/s)

Sum other (kmol/s)

Balanced (kmol/s)

Box (original) Shelf, 0 – 100 m Shelf, 100 – 200 m Slope, 0 – 100 m Slope, 100 – 450 m Slope, 450 – 1000 m Slope, 1000 – 1600 m


0.45 0.17 0.02 0.08 0.14 0.56


2.07
0.04 2.1
0.38 0.45
2.62


1.91 0.27 2.28
0.07 0.44
2.62


0.35
0.09
0.38
0.44
0.31
2.89

0.14
1 2.47
0.44
1.75
4.43


0.31
1.59 2.01
0.9
1.61
4.43

Box (alternative) Shelf, 0 – 100 m Shelf, 100 – 200 m Slope, 0 – 100 m Slope, 100 – 250 m Slope, 250 – 650 m Slope, 650 – 1400 m Slope, 1400 – 1900 m


0.45 0.17 0.02 0.26
0.25 0.16 0


2.07
0.04 2.1
0.44
2.78
2.1
0.7


1.91
0.2 2.3
0.5
2.96
2.56
0.91


0.35
0.09
0.38
0.16
0.04
1.95
3.22

0.17
1 2.44 1.99 0.3
3.05
3.24


0.28
2.01 2.2 0.77
0.37
3.63
3.58

S – N denotes the left hand side of Eq. (3), i.e., the net south – north flux into the box. ‘‘Sum other’’ denotes the sum of all the terms on the right hand side of Eq. (3), i.e., the sum of the other fluxes out of the box. ‘‘Balanced’’ also denotes the right hand side of Eq. (3), using the balanced vertical flux calculation (Eq. (4)). Units are flow rate into/out of box, multiplied by salinity (difference) in dimensionless whole-number units (e.g., typical salinity value is 0.0358).

flow and salinity differences; extra freshwater inflows (reduced salinity) in the top shelf box; generally larger transports and salinity differences below 250 m. The oxygen fluxes show many similar features. ‘‘Sum other’’ and ‘‘balanced’’ at the MSOW level are negative, reflecting reduced oxygen content relative to

above, below and offshore. Again, these terms are much larger than the S – N flux estimates. Negative flux values in 0– 100 m over the shelf in summer reflect depressed oxygen concentrations there and suggest a deficit (otherwise unquantified) from the warm riverine input.

128

J.M. Huthnance et al. / Journal of Marine Systems 32 (2002) 107–137

Table 6 Estimated terms in the Mediterranean water box (650 – 1400 m)

Present Daniault et al. (1994) Huthnance (1995)

q, Sv

SS – SN

qV, Sv

SV

0.86 2.66 q

0.034 0.078 0.063

1.88 1.32 0.16q

0.048 0.157 0.4

Italicised terms are inferred by forcing a closed budget. For Huthnance (1995), there is no identification with the offshore extent of the present slope box. Daniault et al. (1994) used various figures; for the comparison here, we have used their smaller flow speed, 0.06 m/s, south – north salinity gradient as given, their average inferred value, K = 500 m2 s
1, together with the box scale (l = 63/2 km) for qV and hence SVl
2 = 1.58  10
10 m
2, easily within their wide range of utilised observed values ASxxA.

Neither the ‘‘original’’ nor the ‘‘alternative’’ box system gives a consistently better budget. To summarise, the attempted budgets show evidence of riverine inputs (reduced salinity and oxygen), MSOW characteristics (increased salinity, reduced oxygen) and underestimation of S –N flow and gradients. Further discussion is given in Section 6.

5. Independent estimates of process contributions We consider possible alternative process-based estimates to lateral transport, exchange velocities or effective diffusion, to support and add interest to the budgeting. 5.1. Upwelling and downwelling Monthly average values of the upwelling index UI = s/qf for the Vigo region (Fig. 8, Table 3) indicate a typical summer value 0.5 m2 s
1 for the exchange transport q. A similar value (opposite sign) holds for winter downwelling. Several months in spring and autumn have much smaller typical exchange as a monthly mean. However, there are large variations of the upwelling index within months, shown by ‘‘sdd’’ and ‘‘sd6h’’ values in Table 3. Thus, ‘‘mabsd’’ in Table 3 shows average daily absolute values that are larger than the monthly means (even after adding the latter’s variability). In particular, exchange in winter approaches 1 m2 s
1 in association with stronger winds. Slight minima f 0.4 m2 s
1 at either end of the upwelling season are also indicated.

‘‘Mabsd’’ in Table 3 are based on daily values, an inherently arbitrary choice. However, the majority of variability in the 6-hourly values, ‘‘sd6h,’’ is present in the daily values ‘‘sdd.’’ The exchange-flow duration necessary (before reversal) for irreversible constituent transport depends on the particular constituent’s diffusive or non-conservative behaviour. By taking daily values, we include transfers on time scales of days during which constituents (e.g., phytoplankton) may undergo substantial changes. A long return flow time scale in the filament current system allows faster-sinking material to fall over the slope and deep ocean rather than on the shelf. Along-slope transport grows as the integral fm UIdt prior to limitation by friction; about 8.4 m2 s
1 (0.084 m/s in 100 m water depth) at 42jN for each day with strong UI = 1 m2 s
1. 5.2. Filaments and eddies Lagrangian statistics from OMEX drifters provide estimates of dispersion rates (Table 7) in August 1998, in an upwelling filament (Section 2.3) and in winter 1998/1999. In summer, the filaments induce enhanced, roughly isotropic dispersion. In winter, dispersion is less, especially across-shore. The summer – winter difference is hinted at by slightly larger values of AuVA in summer over the shelf (Table 4); over the slope, AuVA is, if anything, larger in winter. Autumn 1986 drifters gave intermediate values (Haynes and Barton, 1991). The drifters and filaments are both near-surface. Similar overall figures were obtained from the ARCANE floats at 450 and 1000 dbar. (Zonal and meridional components were estimated as K f tL(uV2) from the integral Lagrangian time scale, tL (first zero of the autocorrelation function) and corresponding velocity component variance (uV2). Our ‘‘OMEX box’’ area may show a local minimum in values of K.) Table 7 Lateral effective diffusivity values, m2 s
1

OMEX

ARCANE

Period

Across-slope

Along-slope

Summer, 1998 Winter, 1998/99 Autumn, 1986 Overall 42 – 43jN

870 190 340 560 400 – 900

1070 330 250 850 1300 – 1500

‘‘Overall’’ uses all the data as if contemporaneous.

J.M. Huthnance et al. / Journal of Marine Systems 32 (2002) 107–137

For 15 –25jN off northwest Africa, Kostianoy and Zatsepin (1996) estimated a velocity 0.01 m/s (or transport 0.6 m2 s
1) of water exchange, based on filament statistics (60 filaments per year in this latitude range, each of average width 30 km, active life 5 days) and in situ measurements for one filament (velocity 0.4 m/s through 60 m depth). Their corresponding on/offshore effective diffusivity was 550 m2 s
1. In this case, the estimated transport is less than the estimated mean Ekman transport 1 –2 m2 s
1, whereas offshore transport (about 0.6 Sv) in the 1998 filaments measured in OMEX exceeded the corresponding Ekman transport (Section 2.3). Filament statistics for northwest Iberia have been derived from satellite images. In 1993– 1999, the average numbers of imaged filament-days/year attributed to ranges 42 –43jN and 41 –44jN are 141 and 258, respectively (with some ambiguity from nonzero filament width). The greater prevalence (filament
days/year per 1j) in 42 –43jN than in 41 – 44 jN may be due partly to the ambiguity in attribution and partly to the prevalence of upwelling off Cape Finisterre, as found in a previous satellite survey for 1982 – 1990 (Haynes et al., 1993). Mean filament width is about 30 F 9 km; the ‘‘mature’’ and ‘‘developing’’ filaments observed in 1998 had a width of 34.4, 18.4 km, respectively. The images do not indicate filament depth, and only very rarely can features be tracked to estimate currents in a filament (the filament ‘‘nose’’ is generally found to move offshore more slowly than the current within the filament, as water is mixed or subducted at the nose). Hence, we have no basis for changing the 1998 estimate 0.6 Sv transport for the imaged filaments. Thus, we calculate an average offshore transport in filaments as about 0.6 Sv for 100 days/year per 1j latitude, i.e., 1.5 m2 s
1. Eddies carry Mediterranean or slope waters offslope to the ocean; such eddy transports may be important locally, at Cape St. Vincent where Meddies are regularly shed and Cape Ortegal to which the warm winter slope current regularly extends and tends to flow off-shelf; the quantity may be estimated very approximately. Pingree and Le Cann (1993) indicate the volume of one example Meddy leaving the southern Portuguese slope as approximately 500 m depth by 17 km radius, i.e., 454 km3, and that perhaps five per year may be shed in this area

129

(Setubal/Cape St. Vincent), i.e., an average rate 0.07 Sv. Bower et al. (1997) found similar eddy sizes (radii 10 –20 km) but estimated 15 –20 per year shed near Cape St. Vincent and Estremadura Promontory. For sites north of 40jN, the number of eddies are uncertain and exchanges via eddies’ outer circulation and any interacting cyclones (Pingree and Le Cann, 1993) are unquantified. Around Cape Ortegal and in the Bay of Biscay, Pingree and Le Cann (1992a,b) discuss slope water eddies. Two examples had estimated volumes 400 km3 (F90a) and 1570 km3 (X91 using diameter 50 km and depth 800 m). In the years 1979– 1991 typically one eddy is observed, but in years of strong warm winter flow along the slope, there may be more, e.g., three eddies were observed in 1990. We have obtained statistics of eddy manifestations in satellite images for the area 40.5 – 45.5jN out to 13jW for 1993 – 1999. They were rarely apparent in four of these years (at most 20 eddy-days each) but 1996, 1997 and 1998 had totals of 132, 111 and 252 apparent eddy-days. These numbers are underestimates; eddies are missed if obscured by cloud or lacking a sea-surface temperature signal. The eddies had mean diameter f 52 F 22 km. Some tended to be associated with filaments; most of the others were clustered north of Cape Finisterre –Cape Ortegal and so tend to confirm the earlier findings. Also allowing for the western slope and Galicia Bank, we thus base our transport estimate on another five eddies per year around NW Iberia, with volume of order 500 km3 each. Then for f 1000 km of Iberian and Biscay slope, 20 eddies of 500 km3 each year amounts to 0.3 Sv or 0.3 m2 s
1, a modest contribution to the overall exchange. 5.3. Slope current Cross-slope exchange takes place in the bottom boundary (Ekman) layer under the along-slope flow. In addition to the wind-forced contribution (Section 5.1), we estimate a contribution from forcing by the meridional oceanic density gradient and corresponding pressure field. If friction balances this forcing, then the baroclinic zonal eastward flow from the ocean is returned by the bottom Ekman layer. The exchange transport is gHs/8f (Huthnance, 1995), if a uniform meridional density gradient qy is assumed throughout the ocean depth H so that the steric slope s is Hq
1 qy .

130

J.M. Huthnance et al. / Journal of Marine Systems 32 (2002) 107–137

For the Iberian context, Frouin et al. (1990) used hydrography in 38 – 42jN to estimate an onshore baroclinic transport 1 m2 s
1 in the top 250 m. (Arhan et al. (1994) also estimated zonal transports from hydrography in the northeast Atlantic, but our latitude range is not explicit in their results; their baroclinic transport is hard to infer.) Hence, 1 m2 s
1 in a downwelling sense is the best estimate for cross-slope exchange induced by the oceanic density field. An estimate for total offshore transport from the slope current (from all causes) derives from Jorge da Silva’s (1996) geostrophic calculations of its transport from hydrographic sections. Table 1 shows the transport decreasing northwards by 3.7 Sv/1000 km or 3.7 m2 s
1 (obtained by linear regression). This large value applies well offshore, at the outer edge of the slope current, and might be affected by branching of the flow around Galicia Bank (Section 2.2).

6. Discussion 6.1. Context description

In a two-layer model of an internal soliton, lowerlayer transport is cn (to a lowest order –linear-approximation; Inall et al., 2001). Here c=[ gDqH1H 2/ q(H1 + H2)]1/2 is the long wave speed, on upper and lower layer depths H1, H2 with density difference Dq, and n is the interface displacement. Taking gDq/ q = 10
2 m s
2, H1 = 40 m, H2 = 100 m and n = 20 m for 10% of the time (cf. Section 2.2) gives an average exchange transport f 1 m2 s
1 in summer when the waves occur.

Water masses on large scales (>100 km) are well described in Van Aken (2000a,b, 2001) and summarised in Section 2.1. However, to form useful local budgets for distinct summer and winter regimes with spatial resolution distinguishing the shelf, slope and different water masses, requirements are more stringent. There are fewer data to average in estimating for any one such budget; cruise data from just two summers and two winters qualified for use. Differences over short distances are small, hence, sensitive to measurement accuracy and variability. Variability in time and space is relatively important in local budgets; for many boxes, widely different estimates of gradients were averaged. Similarly, Sections 2.2 –2.4 and 5 reflect a good overall description and documentation of the main flow features and processes: a poleward slope current, upwelling and filaments, eddies, tides, wind-forced current variability, lateral exchange and mixing. Summer and winter regimes are described. Again, budgets respecting different areas and regimes pose more stringent requirements. Despite the ability to use measurements outside 42– 43jN, there are few current meter records from the upper slope, or from the shelf, especially in summer. Floats and drifters provide extra information for exchanges rather than local box/seasonal average currents.

5.5. Fronts

6.2. Budgeting

Fronts occur at the offshore boundary of any river plume or upwelled water having characteristics different from the adjacent surface waters. Up/downwelling exchanges are estimated in Sections 5.1 and 5.2. For river plumes, experiments and theory suggest lateral detrainment of about one third of the plume transport near the nose (Stern et al., 1982). Plume transport attributable to 42– 43jN is estimated as roughly 0.1 Sv, from (a) (attributable runoff < 500 m3 s
1; Sections 4.3 and 4.4)  (ambient salinity 35.8/plume salinity deficit 0.2), (b) along-shelf transport estimate q f 0.1 Sv in winter for the shelf 0 –100 m box (Table 4). Hence, cross-frontal exchange for the plume appears to be 0.3 m2 s
1 or less.

Measurement-based flux budgeting (Sections 3 and 4) has not given close balances in general (with little to choose between the original and alternative boxdepth partitions). However, the budget attempts are informative. For our boxes’ depth/width ratio, precipitation –evaporation and vertical flux between boxes are relatively small (this conclusion confers little merit on the original box scheme’s property of zero vertical flux between boxes). Therefore, the main contributions are from horizontal transports and gradients, plus riverine runoff to the top shelf box. Distributions of the budget discrepancies indicate possible error sources. Thus, coastal boxes suggest that runoff may have been underestimated, especially

5.4. Internal (tidal) solitons

J.M. Huthnance et al. / Journal of Marine Systems 32 (2002) 107–137

in winter. Values would probably be increased if runoff were attributed more locally, rather than a portion of the runoff coming to the coastal sector 40 – 43jN; runoff from rivers north of the Douro exceeds their share based on catchment area; the local ´ lvarez-Salgado et al. (submitted for estimate of A publication) is f 50% greater. Our use of climatological runoff, rather than the particular years of utilised marine hydrographic data, may also be a factor. S– N salinity flux estimates are consistently less than the sum of the other fluxes; underestimation was also suggested by comparison with previous MSOW budgets (Section 4.4). The S – N salinity difference is subject to local variation; individual cruises even showed reversal, perhaps from temporal aliasing or from branching of the flow (e.g., around Galicia Bank). The mean currents hvi (Table 4) used for the budgeting are of order 0.02 m/s in general, and 0.04 m/s only near the surface in winter; these are much less than some values in Section 2.2, but consistent with ARCANE float statistics. Differences from Section 2.2 values arise because the poleward slope current and MSOW ‘‘core’’ are close against the slope; the measurements (out to 10j20VW) were not usually so close to the steep slope. Net transport across 10j20VW into or out of the along-slope flow is not determined: Section 5.3 (Table 1) and Maze´ et al. (1997; Section 2.2) suggest export from the slope current; Frouin et al. (1990) suggest a small import in the upper layers. These uncertain budget elements (distribution of runoff; inter-annual and within-season variability; variable S – N salinity differences, current distribution within a box) illustrate the inevitable compromises between resolution and good statistics in a measurement-based budget. Uncertainties are both from insufficient data within boxes (suggesting use of larger boxes) and from unresolved structure within boxes (suggesting use of smaller boxes). There is no obvious objective methodology for box choice; features, areas and seasons of interest also enter the choice. This reinforces our initial proposal that models (e.g., Stevens et al., 2000; Coelho et al., 2002) should be used for finer resolution in space and time; they intrinsically close budgets. Nevertheless, the present attempt helps to define budgetary elements and some transports that a more detailed model should match. It introduces necessary

131

caution in using observations to test models, by checking their validity when incomplete. The boxes themselves may be suitable domains for model – observation comparison. Model output should be averaged, in space and time in the same way as the observations, to make comparison between equivalent variables. The budgeting helps to reconcile the large differences between transport estimates (Frouin et al., 1990; Jorge da Silva, 1996); also Daniault et al. (1994), Maze´ et al. (1997). In the latter two, one of the main lateral transport terms is estimated essentially by difference and is thereby sensitive to uncertainties, as is clear from the range of values in different sections in Daniault et al. (1994). However, their average inferred value K = 500 m2 s
1 is closely confirmed here. To compare ‘‘mean’’ poleward transports, the present estimate between the coast and 10j20VW is 1.6 Sv. Scaling up to achieve an average balance between the S– N and other ‘‘mean’’ salinity fluxes gives 8 Sv (probably an upper bound; S –N salinity differences might have been scaled up alternatively). Between the coast and 10.5jW, Maze´ et al. (1997) have 9 Sv decreasing to 4 Sv between 37.2j and 43jN. Jorge da Silva’s (1996) values are summarised in Table 1. Overall, these transports are roughly consistent. Frouin et al. (1990) estimate only the winter slope current above 300 m, increasing from f 0.3 Sv to f 0.6 Sv in 38– 42jN. Our winter value above 250 m is 0.8 Sv (alternative boxes, no scaling; for these upper boxes in winter, S – N approximately balances the other salinity fluxes; there is little case for scaling). Thus, the large transports of Jorge da Silva (1996) and Maze´ et al. (1997) are reconcilable with (i) Table 4 if it underrepresents the strongest MSOW flow, as our budget imbalance suggests, and (ii) small values in Frouin et al. (1990), which only cover the top 300 m. 6.3. Flux estimates Effective vertical diffusivity, KV, has been estimated in various contexts. The more direct calculations from FLY dissipation measurements (Sherwin et al., 2002) are typically < 10
4 m2 s
1 except in particular contexts: 3  10
4 m2 s
1 within the summer thermocline due to internal tides and waves (2  10
3 m2 s
1 further south at 41jN in MORENA; Jeans and Sherwin, 2001); 10
2 m2 s
1 near the surface in

132

J.M. Huthnance et al. / Journal of Marine Systems 32 (2002) 107–137

winter storms (the largest values). Less direct inference of 10
3 m2 s
1 or more from the salinity field (Section 2.1) and detailed CTD profiles (ARCANE, Section 2.2) suggest larger values close to the steep slope (generally out of the range measured by FLY). Here, larger KV = Ce/N2 (Section 2.2.3) may be favoured by (i) near-bed shear (observed, see Section 2.2.3, increasing the estimate of dissipation e) and (ii) reduced stratification N2. Model calculations show effective KV values of order 10
4 m2 s
1 in the thermocline, but up to 10
2 m2 s
1 under surface winds and in shear above the bed (Davies et al., 2002; Section 2.2.3), thus also suggesting larger KV against the slope. For on/off-shelf exchanges we have several quite consistent estimates. Thus, K = 560 m2 s
1 (OMEX drifters average, Table 7) is consistent with (i) the ARCANE floats’ value, 400 – 900 m2 s
1, on a slightly larger length scale, (ii) 500 m2 s
1 (Daniault et al., 1994), an average inferred from MSOW salinity gradients, (iii) our on/off-shelf velocity variability 19 mm/s (average AuVA in Table 4, assuming as a length scale our average cross-slope box half-width l = 26 km). Depth-specific values AuVA can be used to estimate exchange transports, e.g., 3.1 m2 s
1 at the 200 m depth contour. This gross rate, 3.1 m2 s
1, exchanges the shelf – sea cross-section (3.32 km2; Table 4) in about 12 days. Not all water everywhere on the shelf will be exchanged in this time; some areas are probably less accessible to exchange. Nevertheless, this is a short time relative to seasonal cycles or the estimate f 9 months for the very broad North Sea, for example, see Huthnance (1997). Several processes contribute significantly to this exchange (Section 5): winds favouring up/downwelling (nearly 1 m2 s
1, especially in winter), filaments (about 1.5 m2 s
1 on average, much more locally), secondary circulation in the slope current from meridional oceanic density gradients (1 m2 s
1), internal tidal solitons (1 m2 s
1). Smaller contributions may be from plume fronts, 0.3 m2 s
1 or less, and eddies, estimated here as of order 0.3 m2 s
1. Another estimate of eddy exchange might be derived from the fates of floats, indicating (e.g.,) the fraction of slope water going to Meddies, as discussed in Le Cann et al. (2002b, in preparation). Barotropic tidal exchanges of order 0.05 m/s current  200 m depth = 10 m2 s
1 are large, but their large scale and reversal every 6 h inhibit

lasting exchange. The separate process estimates are not precise enough to expect addition to the gross value, 3.1 m2 s
1. Indeed, they do not apply equally everywhere; the full value of the slope current secondary circulation is probably off-shelf from the shelf break, whereas the internal tidal solitons are on-shelf from the shelf break. Where processes do occur together, their exchanges may not be additive, especially in respect of wind forcing, filaments and slope currents. Nevertheless, these process estimates amount to further support for the gross exchange rate. 6.4. Nutrient supply We estimate the annual supply of nutrient-rich water for near-surface new production, of interest in the OMEX context. Nutrients are supplied by (a) upwelling and (b) upward diffusion. For (a), we use the daily values of upwelling index when positive in 1993 – 1999. Implicitly, we assume: that upwelled water is nutrient-rich, not depleted from previous upwelling; that the nutrients supplied by upwelled water are consumed before any subsequent downwelling; enough light to allow production through the winter. These assumptions all improve on longer time scales, so the daily based estimates ‘‘sumup’’ in Table 3 are upper bounds. Distributed across 33.7 km shelf width, the annual total 8788 UI days amounts to 225 m ‘‘equivalent depth.’’ For (b), we suppose KV < 10
4 m2 s
1 for the whole year over a depth range 50 m below the surface mixed layer (e.g., in and below the summer thermocline). This upper bound is 63 m ‘‘equivalent depth’’; it is probably less in the ocean where Sherwin et al. (2002) found eddy diffusion about one third of values over the shelf. Thus, our estimate of ‘‘equivalent depth’’ (a + b) is 200– 300 m (maximum). A comparative estimate over Goban Spur is about 100 m (Huthnance et al., 2001). The great majority off Iberia is contributed by upwellingfavourable winds, which vary from year-to-year (Sections 2.3 and 5.1); for comparison with 225 m average ‘‘equivalent depth,’’ values in 1997, 1998, 1999 were 181, 249, 210 m, respectively. 6.5. Variability Our flux and budget estimates are climatological; they do not describe variability between years or

J.M. Huthnance et al. / Journal of Marine Systems 32 (2002) 107–137

within the ‘‘summer’’ and ‘‘winter’’ seasons. We briefly discuss these aspects. In relation to on-/off-shelf fluxes, the largest estimated contributors (secondary circulation with the slope current, upwelling, filaments and internal solitons) all show substantial variability. However, the first two have prospects of good estimation from large-scale variables: the oceanic density field and wind forcing. Meridional density gradients and winds are integrated over the order of 1000 km to force the slope current (Huthnance, 1984), reducing sensitivity to variations on short space scales (e.g., variable latitude of branches of the Azores front). The density field on 1000 km scales is more stable (Section 2.1, the widescale inter-annual variation in ENACW properties does not appear to affect our local budgets which are insensitive to an overall change of salinity). Nevertheless, Coelho et al. (2002; Section 2.2.1) find benefit in using actual rather than climatological density in modelling the slope current. Most MSOW variability relates to branching and eddy shedding, i.e., short spatial scales. Below this, the large-scale oceanic density field is established on time scales of many years (even AOU ageing in southward ENACW flow implies some years from its origin). Short-period current fluctuations associated with the large-scale wind and atmospheric pressure fields are possible. The upwelling index (UI; Sections 2.3 and 5.1; Table 3) is a good indicator of wind variability. Thus, measures of integrated upwelling (crossshelf exchange transport) vary by a factor of 2 between extreme years (1998 and 1999 being strong and weak, respectively). However, month-to-month variability is much greater (even week-to-week variability outside the upwelling season), being smoothed by taking the annual integral. At any one time, direct wind-forced exchange can be estimated from the current upwelling index. Filaments are stochastic; associated exchange depends on time and location relative to filament occurrence. The 1993 –1999 statistics (Section 5.2) show a factor of 2 between the years with most and least imaged filament-days; one might estimate a similar factor between the associated total exchanges (i.e., from 1 to 2 m2 s
1 on average). Internal solitons relate to the seasonal stratification, but we do not have the evidence to quantify variability except for comparison between OMEX and MOR-

133

ENA with different years and locations (Section 2.2.3; Sherwin et al., 2002; Jeans and Sherwin, 2001). Models (e.g., Davies et al., 2002) suggest much variability on short time and space scales associated with evolving stratification; for example, the degree of upwelling or downwelling. Eddies have highly variable statistics (Section 5.2) but appear to contribute only modest exchange. For other fluxes in box budgets, the main uncertainty in the (typically underestimated) south – north flux relates to S – N gradients, well-known on the large-scale (Section 2.1) but variable to the extent of sign reversal because of branching and eddies at the box scale. Quantification is difficult; it might be possible in a future model assimilating sufficient near-real-time hydrography, or over long periods by a model with the right branching and eddy statistics. Vertical mixing appears to be small overall for our boxes; it is extremely variable with a range 10
5 to 10
2 m2 s
1, higher values occurring near the surface in storms, in internal waves (i.e., in ‘‘events’’) and perhaps close to the slope. Runoff is important to the coastal box; Douro values for the years 1985 – 1990 vary modestly in the range 147 –212 (  103 m3 s
1 days), but range from 1 to 82 for individual months in that period; month-to-month variation is probably most relevant for the plume in the box of 100 km scale. Quantification needs a coastal model with (available) daily runoff input. For the ‘‘equivalent depth’’ of nutrient supply to new production (Section 6.4) a factor 1.4 spans the 1997, 1998, 1999 values (integral of positive daily UI for the year). (These years’ integrated upwelling values span a factor 2. Moreover, the two measures are poorly correlated).

7. Conclusions We have a good description of the broad-scale mean hydrography. Features include Mediterranean Sea Outflow water (MSOW) close to the slope, with high salinity decreasing polewards, albeit only slowly off Galicia. Above, there is a salinity minimum near 450 m depth; in the adjacent ocean, this is in southward flow, but with little mean flow against the slope. Above the salinity minimum, close to the slope, is a persistent but varying poleward flow. In winter, this

134

J.M. Huthnance et al. / Journal of Marine Systems 32 (2002) 107–137

slope current may surface; in summer, it may be only below 200 m (but at times as shallow as 100 m depth) with upwelling and southward flow near the surface. There is large-scale inter-annual variability and branching of these MSOW and slope flows. The circulation also shows large variability on shorter time scales: seasonal (upwelling), sub-tidal/synoptic (changes in wind forcing, eddies), tidal and higher frequencies. Upwelling varies inter-annually; some spring, autumn and winter months may have net upwelling conditions; only the months June to September reliably have mean upwelling-favourable conditions. An overall seasonal pattern in currents has been established, but with large year-to-year variations. The large-scale currents also show (i) large cross-shelf variance and offslope flow off southern Portugal and NW Spain, especially at the Mediterranean water levels, and (ii) low cross-shelf variance in 42– 43jN. Shelf currents respond to direct forcing by winds, including those favouring up/downwelling, with common spectral peaks at periods of a few days. For budgeting fluxes, ‘‘the devil is in the detail.’’ Flux variability exceeds mean values and causes problems. Length scales of 50 km for filaments and eddies (for example) are of the same scale as the budgeted OMEX area; synoptic surveys can give reversed concentration gradients. Careful consideration is needed of the box size and time period over which averages are taken, to provide stable flux estimates for a budget balance and yet resolution for valid use as a model test. In practice, the available data did not allow such a compromise; uncertainties remain both from insufficient data and from unresolved structure. Observations do not yield water fluxes directly but have helped understanding of the main controlling processes. The budget attempt provides evidence for model validation. Precipitation –evaporation over the sea and vertical exchange were small contributors to the budgets of the chosen boxes. The main fluxes are from river runoff in winter, lateral transports and constituent gradients. Various transport estimates can be reconciled, and tend to support a value f 4 Sv at 42– 43jN (MSOW and above; east of 10.5jW), decreasing polewards. Less than 1 Sv of this appears to be in the slope current above the salinity minimum.

Models constrained by data (Coelho et al., 2002) should be used for finer resolution in space and time for the eventual best quantification of fluxes and formation of closed budgets. Vertical diffusivity, KV, is typically estimated as less than 10
4 m2 s
1 except in particular contexts: 3  10
4 m2 s
1 within the summer thermocline due to internal tides and waves; 10
2 m2 s
1 near the surface in winter storms (the largest values). Less direct estimates and models suggest large values 10
3 m2 s
1 or more close to the steep slope. Effective cross-shelf diffusivity, K, is rather consistently estimated at about 500 m2 s
1 and corresponds with cross-shelf current variability AuVA f 19 mm/s or ocean-shelf exchange 3.1 m2 s
1. The main contributing processes (about 1 m2 s
1 each on average; not simply additive) are winds favouring up/downwelling, filaments (about 1.5 m2 s
1), secondary circulation in the poleward slope current, internal tidal solitons. Estimated exchange in shed eddies is only of order 0.3 m2 s
1. Exchange flows in filaments after upwelling are variable but appear to give more cross-margin exchange than the Ekman transport forcing upwelling. Filaments transport water off-slope and may also bring water from large distances onto the shelf. The rate 3.1 m2 s
1 exchanges the volume of shelf water with the ocean in about 12 days. The ‘‘equivalent depth’’ of nutrient-rich water upwelled for new production, spread across the width of the shelf, has an estimated upper bound 225 m annually. This value is less if upwelled water is depleted from previous upwelling or its nutrients are not all used up. An upper bound on diffusive supply to near-surface water is 63 m; probably less in the adjacent ocean. To put OMEX fluxes into a wider context, e.g., comparison with Goban Spur or global perspectives, fluxes for the whole Iberian margin may have to be taken into account, due to large inferred off-slope fluxes off both southern Iberia and the northwest corner. Notation AOU Apparent Oxygen Utilisation ARCANE France – USA NE Atlantic project, 1996– 1999 (Le Cann et al., 2002a,b, in preparation) B Buoyancy loss of turbulence

J.M. Huthnance et al. / Journal of Marine Systems 32 (2002) 107–137

BODC British Oceanographic Data Centre Bord-Est French hydrography and current measurements NE Atlantic boundary, 1988 – 1989 c (long) wave speed CS, CN, Ci, Cˆ, CiV Constituent concentrations: S, N ends, side i inflow, outflow (CS + CN)/2, Cˆ
Ci ENACW Eastern North Atlantic Central Water f Coriolis parameter: twice the vertical component of Earth’s rotation g Acceleration due to gravity ‘‘Galicia’’ Instituto de Investigacio´n Marin˜as (Vigo, Spain) research cruise programme along the margin off Galicia, 1975 to 1991 H, H1, H2 Ocean or layer depths as defined locally Ifm Institut fu¨r Meereskunde, Universita¨t Kiel IH Marinha-Instituto Hidrografico, Lisboa, Portugal ISOW Iceland Scotland Overflow Water JGOFS Joint Global Ocean Flux Study, International Geosphere-Biosphere Programme K, Ki, KV Diffusivities: lateral, across budget box side i, vertical L, l Budget box length, offshore scale f Wi/2 LDW Lower Deep Water LO Ozmidov length scale, e1/2N
3/2 (Ozmidov, 1965) LSW Labrador Sea Water LT Thorpe scale of density inversions (Thorpe, 1977) Meddy Mediterranean Water eddy MORENA Multidisciplinary Oceanographic Research in the Eastern boundary of the North Atlantic; EU project MAS2-CT93-0065 MSOW Mediterranean Sea Outflow Water N Buoyancy frequency: N 2 =
gqz/q0 ni Outward normal to budget box side i NEADW Northeast Atlantic Deep Water OMEX Ocean Margin Exchange; EU project MAS3CT97-0076 P Turbulence production rate qS, qN, qi, qiV volume inflow at S end, outflow at N end and side i, exchange across side i RAFOS Neutrally buoyant floats positioned relative to moored acoustic sources RAYO Puertos del Estado (Spain) marine physical environment network Rf Flux Richardson Number B/P

135

S s

salinity steric slope, Hq
1qy, where H is the full ocean layer depth to which qy pertains SEFOS Shelf Edge Fisheries and Oceanography Study; EU project AIR2-CT93-1105 SOMA September/October – March/April seasonal variation, (Pingree et al., 1999) Sv Sverdrup u 106 m3 s
1 t time UI Upwelling index (see Table 3) u onshore or eastward velocity component v along-shore or northward velocity component Wi Width/depth of budget box side i WOCE World Ocean Circulation Experiment x cross-shelf or eastward coordinate y along-shelf or northward coordinate z vertical coordinate C Rf /(1
Rf) DS, Dq difference (in salinity S, density q) e turbulent dissipation H potential temperature q density s wind stress n interface displacement Subscripts denote differentiation, h. . .i — ensemble mean, (V)—deviation from the mean Acknowledgements This work was supported by the EU through the MAST programme, contract MAS3-CT97-0076 (Ocean Margin Exchange — OMEX). We thank Toby Sherwin for helpful discussions and early sight of Sherwin et al. (2002), and referees for prompting improvements to the text.

References ´ lvarez-Salgado, X.A., Gago, J., Mı´guez, B.M., Gilcoto, M., Pe´rez, A F.F., 2000. Surface waters of the NW Iberian margin: upwelling on the shelf versus outwelling of upwelled waters from the Rias Baixas. Estuarine, Coastal and Shelf Science 51, 821 – 837. ´ lvarez-Salgado, X.A., Beloso, S., Joint, I., Nogueira, E., Chou, L., A Pe´rez, F.F., Groom, S., Cabanas, J.M., Rees, A.P., Elskens, M., 2002. New production of the NW Iberian shelf during the upwelling season. Deep-Sea Research submitted for publication. Arhan, M., Colin de Verdie`re, A., Me´mery, L., 1994. The eastern

136

J.M. Huthnance et al. / Journal of Marine Systems 32 (2002) 107–137

boundary of the subtropical North Atlantic. Journal of Physical Oceanography 24, 1295 – 1316. Barton, E.D., Inall, M.E., Sherwin, T.J., Torres, R., 2001. Vertical structure, turbulent mixing and fluxes during Lagrangian observations of an upwelling filament system off Northwest Iberia. Progress in Oceanography 51, 249 – 267. Biscaye, P.E., Flagg, C.N., Falkowski, P.G., 1994. The Shelf Edge Exchange Processes experiment, SEEP-II: and introduction to hypotheses, results and conclusions. Deep-Sea Research Part II 41, 231 – 252. Bower, A.S., Armi, L., Ambar, I., 1997. Lagrangian observations of Meddy formation during a Mediterranean undercurrent seeding experiment. Journal of Physical Oceanography 27, 2545 – 2575. Brink, K.H., Cowles, T.J., 1991. The coastal transition zone program. Journal of Geophysical Research 96, 14637 – 14647. Cherubin, L., Carton, X., Paillet, J., Morel, Y., Serpette, A., 2000. Instability of the Mediterranean water undercurrents southwest of Portugal: effects of baroclinicity and of topography. Oceanologica Acta 23, 551 – 573. Coelho, H.S., Neves, R.J.J., White, M., Leita˜o, P.C., Santos, A.J.P., 2002. A model for the Iberian coastal ocean circulation. Journal of Marine Systems 32 (this issue). Daniault, N., Maze´, J.P., Arhan, M., 1994. Circulation and mixing of Mediterranean water west of the Iberian peninsula. Deep-Sea Research 41, 1685 – 1714. Davies, A.M., Xing, J., Huthnance, J.M., Hall, P., Thomsen, L., 2002. Models of near-bed dynamics and sediment movement at the Iberian margin. Progress in Oceanography in press. Dubert, J.P., 1998. Dynamique du systeme de courants vers le pole au voisinage de la pente continentale a l’Ouest et au nord de la Peninsule Iberique. PhD Thesis. Universite´ de Bretagne Occidentale. Ferron, B., Mercier, H., Speer, K., Gargett, A., Polzin, K., 1998. Mixing in the Romanche fracture zone. Journal of Physical Oceanography 28, 1929 – 1945. Frouin, R., Fiuza, A.F.G., Ambar, I., Boyd, T.J., 1990. Observations of a poleward surface current off the coast of Portugal and Spain during winter. Journal of Geophysical Research 95, 679 – 691. Haynes, R., Barton, E.D., 1990. A poleward flow along the Atlantic coast of the Iberian Peninsula. Journal of Geophysical Research 95, 11425 – 11441. Haynes, R., Barton, E.D., 1991. Lagrangian observations in the Iberian coastal transition zone. Journal of Geophysical Research 96, 14731 – 14741. Haynes, R., Barton, E.D., Pilling, I., 1993. Development, persistence and variability of upwelling filaments off the Atlantic Coast of the Iberian Peninsula. Journal of Geophysical Research 98, 22681 – 22692. Huthnance, J.M., 1984. Slope currents and ‘‘JEBAR’’. Journal of Physical Oceanography 14, 795 – 810. Huthnance, J.M., 1995. Circulation, exchange and water masses at the ocean margin: the role of physical processes at the shelf edge. Progress in Oceanography 35, 353 – 431. Huthnance, J.M., 1997. North Sea interaction with the North Atlantic Ocean. Deutsche Hydrographische Zeitschrift 49, 153 – 162. Huthnance, J.M., Coelho, H., Griffiths, C.R., Knight, P.J., Rees, A.P., Sinha, B., Vangriesheim, A., White, M., Chatwin, P.G.,

2001. Physical structures, advection and mixing in the region of Goban Spur. Deep-Sea Research, Part II 48, 2979 – 3021. Huthnance, J.M., Humphery, J.D., Knight, P.J., Chatwin, P.G., Thomsen, L., White, M., 2002. Near-bed turbulence measurements, stress estimates and sediment mobility at the continental shelf edge. Progress in Oceanography in press. Inall, M.E., Shapiro, G.I., Sherwin, T.J., 2001. Mass transport by non-linear internal waves on the Malin Shelf. Continental Shelf Research 21, 1449 – 1472. ´ gua, 2000. Web page , http://www.inag.pt/snirh. Instituto da A Jeans, D.R.G., Sherwin, T.J., 2001. The evolution and energetics of large amplitude nonlinear internal waves on the Portuguese shelf. Journal of Marine Research 59, 327 – 353. Jorge da Silva, A., 1996. The slope current off the west Iberian coast in autumn. ICES C.M. 1996/S:35 Kostianoy, A.G., Zatsepin, A.G., 1996. The West African coastal upwelling filaments and cross-frontal water exchange conditioned by them. Journal of Marine Systems 7, 349 – 359. Le Cann, B., Serpette, A., Colas, F., 2002a. Long-term Eulerian current observations over the northwest Iberian Peninsula slope and nearby ocean. In preparation. Le Cann, B., Serpette, A., Colas, F., Carton, X., 2002b. Lagrangian circulation near the eastern boundary of the mid-latitude North Atlantic Ocean. In preparation. Maze´, J.P., Arhan, M., Mercier, H., 1997. Volume budget of the eastern boundary layer off the Iberian Peninsula. Deep-Sea Research, Part I 44, 1543 – 1574. Monaco, A., Biscaye, P.E., Soyer, J., Pocklington, R., Heussner, S., 1990. Particle fluxes and ecosystem response on a continental margin: the 1985 – 1988 Mediterranean ECOMARGE experiment. Continental Shelf Research 10, 809 – 839. Ozmidov, R.V., 1965. On the turbulent exchange in a stably stratified ocean. Izvestiya, Atmospheric and Oceanic Physics 1, 853 – 860. Paillet, J., Le Cann, B., Serpette, A., Morel, Y., Carton, X., 1999. Real-time tracking of a Galician Meddy. Geophysical Research Letters 26 (13), 1877 – 1880. Paillet, J., Le Cann, B., Carton, X., Morel, Y., Serpette, A., 2002. Dynamics and evolution of a northern Meddy. Journal of Physical Oceanography, in press. Pingree, R.D., 1995. The droguing of Meddy Pinball and seeding with ALACE floats. Journal of the Marine Biological Association of the United Kingdom 75, 235 – 252. Pingree, R.D., Le Cann, B., 1990. Structure, strength and seasonality of the slope currents in the Bay of Biscay region. Journal of the Marine Biological Association of the United Kingdom 70, 857 – 885. Pingree, R.D., Le Cann, B., 1992a. Anticyclonic Eddy X91 in the southern Bay of Biscay, May 1991 to February 1992. Journal of Geophysical Research 97, 14353 – 14367. Pingree, R.D., Le Cann, B., 1992b. Three anticyclonic Slope Water Oceanic eDDIES (SWODDIES) in the southern Bay of Biscay in 1990. Deep-Sea Research 39, 1147 – 1175. Pingree, R.D., Le Cann, B., 1993. A shallow meddy (a Smeddy) from the secondary Mediterranean salinity maximum. Journal of Geophysical Research 98, 20169 – 20185. Pingree, R.D., Sinha, B., Griffiths, C.R., 1999. Seasonality of the

J.M. Huthnance et al. / Journal of Marine Systems 32 (2002) 107–137 European slope current (Goban Spur) and ocean margin exchange. Continental Shelf Research 19, 929 – 975. Quaresma, L.S., Coelho, E.F., Rosa, L., Beja, J., 2000. Internal tide on Nazare´ submarine canyon — head canyon measurements. In: Dias, J.A., Ferreira, O. (Eds.), Proceedings, 3rd Symposium on the Iberian Atlantic Continental Margin, pp. 165 – 167. Reid, J.L., Mantyla, A.W., 1976. The effect of the geostrophic flow upon coastal sea elevations in the northern Pacific Ocean. Journal of Geophysical Research 81, 3100 – 3110. Rodriguez, I., Alvarez, E., Krohn, J., Backhaus, J., 1991. A midscale tidal analysis of waters around the north-western corner of the Iberian Peninsula. In: Sanchez-Arcilla, A., et al., (Eds.), Computer Modelling in Ocean Engineering 91, Proceedings of the second International Conference on Computer Modelling in Ocean Engineering, Barcelona, 30 September – 4 October 1991. A.A. Balkema, Rotterdam, pp. 33 – 42. Shapiro, G.I., Maschanov, S.L., 1996. Spreading pattern and mesoscale structure of Mediterranean outflow in the Iberian Basin estimated from historical data. Journal of Marine Systems 7, 337 – 348. Sherwin, T.J., Inall, M.E., Torres, R.J., 2002. The seasonal and spatial variability of small scale turbulence at the Iberian margin. Journal of Marine Research 60, 73 – 100. Souza, A.J., Simpson, J.H., Harikrishnan, M., Malarkey, J., 2001. Flow structure and seasonality in the Hebridean slope current. Oceanologica Acta 24, S63 – S76. Stern, M.E., Whitehead, J.A., Hua, B.-L., 1982. The intrusion of a density current along the coast of a rotating fluid. Journal of Fluid Mechanics 123, 237 – 265. Stevens, I., Hamann, M., Johnson, J.A., Fiu´za, A.F.G., 2000. Comparisons between a fine resolution model and observations in the Iberian shelf-slope region. Journal of Marine Systems 26, 53 – 74. Thorpe, S.A., 1977. Turbulence and mixing in a Scottish loch. Philosophical Transactions of the Royal Society of London A286, 125 – 181. Times, 1992. The Times Atlas of the world. Times Books, HarperCollins, London, xlvii + 123 plates + 222 pp. TRANSCAN team, 1999. Dynamical mass transport in submarine

137

canyons (TRANSCAN 98/1 data report). Instituto Hidrogra´ficoPortugal, Rel.TF.OC-03/98. Van Aken, H.M., 2000a. The hydrography of the mid-latitude northeast Atlantic Ocean: I. The deep water masses. Deep-Sea Research, Part I 47, 757 – 788. Van Aken, H.M., 2000b. The hydrography of the mid-latitude northeast Atlantic Ocean: II. The intermediate water masses. DeepSea Research, Part I 47, 789 – 824. Van Aken, H.M., 2001. The hydrography of the mid-latitude northeast Atlantic Ocean: I. The thermocline water mass. Deep-Sea Research, Part I 48, 237 – 267. Van Haren, H., Maas, L., Zimmerman, J.T.F., Ridderinkhof, H., Malschaert, H., 1999. Strong inertial currents and marginal internal wave stability in the central North Sea. Geophysical Research Letters 26, 2993 – 2996. Vitorino, J., 1989. Circulacß a˜o residual ao largo da costa NW de Portugal durante a estacß a˜o de afloramento de 1987 (Residual circulation off the NW Portuguese coast during the 1987 upwelling season). Anais do Instituto Hidrografico 10, 25 – 37 in Portuguese. Vitorino, J., Coelho, E.F., 1998. Water mass dynamics at the Northwest Iberian Margin. Ocean Margin Exchange OMEX II — Phase II First Annual Science Report, pp. 149 – 175. Vitorino, J., Marreiros, M., 2002. Results from CECIR XVII cruise: March – April 1990. Instituto Hidrografico in preparation. Vitorino, J., Oliveira, A., Jouanneau, J.M., Drago, T., 2002a. Winter dynamics on the northern Portuguese shelf. Part 1: Physical processes. Progress in Oceanography (in press). Vitorino, J., Oliveira, A., Jouanneau, J.M., Drago, T., 2002b. Winter dynamics on the northern Portuguese shelf. Part 2: Bottom boundary layer and sediment dispersal. Progress in Oceanography (in press). Walsh, J.J., Biscaye, P.E., Csanady, G.T., 1988. The 1983 – 1984 Shelf Edge Exchange Processes (SEEP)-I experiment: hypotheses and highlights. Continental Shelf Research 8, 435 – 456. Wollast, R., Chou, L., 2001. The carbon cycle at the ocean margin in the northern Gulf of Biscay. Deep-Sea Research, Part II 48, 3265 – 3293.