lit review7

OSX 4016: Literature Review

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Canyons: Sewage Pipes of the Shelf?

Meghan Rochford

Abstract

Shelf seas are regions of great importance both physically and biogeochemically. Many physical

processes affect these regions, such as topography, wind stress, tidal currents and stratification. In

temperate regions topography is a controlling factor of shelf seas, it causes an along-slope geostrophic

flow to form, which cannot cross isobaths. Biogeochemically shelf seas are hotspots for phytoplankton

growth and nutrient recycling. The Celtic Sea is a broad shelf sea known for its strong tidal forcing and

seasonal stratification. During spring and summer it becomes stratified due an increase in surface

temperatures, allowing the thermocline to shoal, bringing nutrients up from the deeper ocean,

causing the spring bloom. The conditions which allow for dense water cascades to occur have been

documented in the Celtic Sea. Submarine canyons are areas of enhanced cross-shelf exchange and are

important topographic features which have been observed to cause enhanced

upwelling/downwelling, mixing due to internal tides, and act as conduits for dense water cascades. In

this review, the difference between long and short canyons will be discussed with reference to

changing Rossby numbers and the effect it has on the flow regime of a canyon.

Shelf Seas

The coastal ocean and open ocean each have diverse physical and biological processes. The area

where they meet, the shelf-edge, has its own unique processes. Exchange between the open ocean

and shelf seas have important implications for shelf-sea currents, flushing and the supply of nutrients

(Huthnance et al. 2009), which in turn, have implications for phytoplankton production (Rees et al.

1999). It is also believed that shelf processes exercise some control on open ocean circulation in ocean

basins and mixing over slopes, which are known to contribute to the main oceanic density structure

(Munk & Wunsch 1998). Topography is an important controlling factor in shelf seas, as it constrains

large-scale flows (geostrophic) from crossing the slope (Huthnance et al. 2009), causing unique

exchange processes at the shelf-edge. This geostrophic constraint is not as severe at the equator, and

is more relaxed, especially in Ekman layers, due to friction (Huthnance 1995). The transport of

nutrients and carbon between shelf seas and the open ocean has important significance to the

nutrient and carbon cycles, however, it has not been adequately quantified (Huthnance et al. 2009).


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Physical Processes

It has been documented many times that along-isobath flow along the continental slope is primarily

steered by steep topography (Huthnance 1995; Allen 2004), meaning that cross-slope flow is hindered,

and the exchange between coastal and open ocean waters is limited (Allen & Durrieu de Madron

2009). This means that a homogeneous, geostrophic flow has no divergence, and therefore moves as

rigid columns of water, unable to change its length (Taylor 1923). Because of the rigidity of the water,

it cannot change depth, confining it between isobaths. In the same way as above, a stratified flow is

limited to a depth with no flow, and anything above that depth can flow across isobaths. Therefore, a

shelf-break current, from the surface to a depth, will block even a stratified geostrophic flow from

crossing isobaths (Allen & Durrieu de Madron 2009). Deep ocean shelf exchange (DOSE) therefore

occurs when there are ageostrophic flow dynamics. This occurs in the presence of large frictional

processes, time dependence, or advection (Allen 2004). Fig. 1 shows the main environmental physical

processes associated with shelf seas: wind stress, tidal currents, stratification, and shelf geometry

(Estrade et al. 2008).

Fig. 1. Summary of the physical processes which occur in a shelf sea (Huthnance et al. 2009).


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The Ekman (1905) paper is the cornerstone concept of upwelling processes. Ekman (1905) states that

wind-forced oceanic mass transport occurs to the right of wind in a shallow layer of water. Northerly

winds drive surface Ekman transport offshore along a western boundary. These conditions are

equatorward at an eastern boundary, and polewards at a western boundary (in the northern

hemisphere)(see Fig. 1 and Huthnance 1995). This causes upwelling of deep, cold water, from a depth

no deeper than 200-300 metres (Pond & Pickard 1983). Sea surface temperatures (SST) taken off the

Northwest African coast shows that there is a temperature minimum close to the coast, showing this

cold water upwelling (Fig. 2)(Estrade et al. 2008).

Fig. 2. Seven-day composite of MODIS SST

images off the northwest African coast.

Shows a temperature minimum close to the

coast (Estrade et al. 2008).

Estrade et al. (2008) used a two-dimensional model to study Ekman’s 1905 theory further, showing

that upwelling occurs offshore of a ‘kinematic barrier’ to the cross-shelf flow (dependant on the level

of stratification), resulting from the upper and lower Ekman layers merging. Their results showed that

90% of Ekman transport upwells for h/D (h being depth and D being the Ekman depth) between 1.25

and 0.5 (alongshore wind) (Fig. 3). Wenju & Lennon (1988) suggested that the seasonal upwelling

studied in the Taiwan Strait was dependant on surface wind stress, driving Ekman pumping.


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Fig. 3. Conceptual schematic of the mechanism of upwelling separation from the coast (Estrade et al. 2008).

Tidal energy, especially internal tides, and subsequently internal waves are important physical

processes needed for mixing in the water column. At the upper shelf slope and shelf edge there is

enhanced internal mixing due to the formation of non-linear internal waves (often tidal) (Fig. 4). These

waves can then break into internal solitons, causing significant fluxes across the shelf edge (Inall et al.

2001). This is all made possible by the relatively steep slope. These solitons cause an increase in

current shear, leading to increased internal turbulence and mixing. This leads to a dissipation of the

energy from the internal tide (Sharples et al. 2007).


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Fig. 4. Schematic showing the formation

and dissipation of internal waves at the

shelf edge. A) During off-shelf ebb flow

the thermocline is distorted, forming a

depression over the shelf edge. B) As the

ebb tidal flow decreases shorter internal

waves, with increased amplitude. C) As

the waves flow onto the shelf, they are

further shortened and steepened, causing

a higher shear and an increase in vertical

mixing (Sharples et al. 2007).

There are different levels of stratification experienced throughout shelf seas (Fig. 1); for some, like the

Irish Sea, seasonal stratification is common. This is when an increased heat flux in spring and summer

out-competes mixing due to wind and tides, and causes the formation of a warmer surface layer, while

the water column remains vertically homogeneous in winter. In other regions the water column

remains vertically homogeneous throughout the year, caused by higher levels of turbulent energy in

the system, due to wind and tidal mixing (van Aken 1986).

Tidal mixing fronts are boundaries between stratified and well-mixed seas (Fig. 1). They occur in

regions with high tidal dissipation adjacent to regions of large seasonal heat exchange (Simpson &

Bowers 1981). The average temperature of the stratified region rises more slowly than the mixed, due

to the heat being kept at the surface. Therefore, the surface layer of the stratified regime has a higher

temperature than that of the mixed regime. This causes a density difference between the two regimes,

with the mixed side having a lower depth-mean density. At the bottom layer depth the density

increases from the mixed regime to the stratified. Overall, this causes a pressure slope at depth and

at the surface from the mixed to the stratified regime. This produces a pressure gradient force (PGF)

driven flow, which is deflected by the Coriolis force, balancing as a geostrophic flow. This flow is

perpendicular to the PGF; therefore, there is an along front flow parallel to the mixing front flow.


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Biogeochemical Processes

Primary production in temperate regions is characterised by seasonal intermittency, and starts with a

spring bloom. This occurs when the light-determined critical depth descends into the mixed layer

depth, allowing for further algal growth. As heating increases and wind stress decreases the mixed

layer depth shoals. The timing of the spring bloom varies locally and inter-annually. For example, the

bloom may occur earlier if there are larger freshwater inputs, causing a shallower surface layer, or it

may be later if there is an increase in suspended sediments, reducing light penetration, and thus

phytoplankton growth. It ends once the initial near surface nutrient concentration becomes limited

(Huthnance et al. 2009). Summer growth depends on upwelling of biologically fixed regenerated

nitrogen. It is exchanged through the thermocline by turbulence from winds, waves and internal

waves. The presence of a subsurface chlorophyll maximum occurs when the surface water becomes

nutrient limited, forcing the advection of nutrients from the bottom layer through the thermocline,

where it is immediately consumed by phytoplankton.

The role of shelf seas in the carbon cycle is of significant importance, yet they are sometimes over-

looked in global estimates of CO2 uptake and production. Using the North Sea as an example, the

difference in CO2 uptake will be discussed. The North Sea can be divided into two regimes: in the north

the water column is stratified, while in the south the water column is shallower, and therefore

vertically homogeneous (mixed) (Fig. 6). In the northern North Sea the uptake of dissolved inorganic

carbon (DIC) occurs in the surface mixed layer, organic material sinks into the bottom layer, where

respiration takes place. This allows the surface layer to have a low concentration of carbon, allowing

for the continual uptake of DIC. This allows the northern part of the sea to act as a sink (Thomas et al.

2004). In the southern North Sea production and respiration occur in the same ‘compartment’. This

means that the uptake and release of DIC is in equilibrium with the atmosphere, causing a higher

concentration of DIC in the water column, causing it to act as a source (Prowe et al. 2009).


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Fig. 6. Schematic of the North Sea: The difference in regimes in the north (sink) and south (source).

The Northwest European Shelf

The Northwest European Shelf encompasses the Hebrides and Malin shelves, the English Channel, the

Irish Sea, the Celtic Sea, and Irish Shelf (Fig. 6). It is approximately 2000 kilometres long, from the

Amorican Shelf in the south, up to the North Sea in the north (46-60°N), at the eastern boundary of

the North Atlantic Ocean. Due to the presence of the British Isles, the region experiences a range of

complex topography.

The region experiences strong tidal forcing at the ocean boundary. The English Channel, Irish Sea and

Bristol Channel are characterised by strong tidal responses, with the largest ranges occurring on the

eastern side of the basins. The largest tidal ranges (>8 metres at M2 tides) have been recorded near

the port of St. Malo, at the head of the Bay of Seine, and Avonmouth, Bristol (Simpson 1998).

The M2 tide enduces frictional stresses at the seabed, over much of the region, with a maximum stress

of 0.25 Pa, which is the equivalent of a sea surface wind stress of 13 ms-1. In the Irish Sea, extreme

stresses have been recorded at 4 Pa, which is on the scale of hurricane force wind stresses. Due to

this, the European Shelf is believed to cause approzimately 12.5% of the global tidal energy loss

(Simpson & Bowers 1981).


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Fig. 6. Map of the Northwest

European Shelf, with the locations

of the Celtic Sea, Irish Sea, Irish

Shelf, English Channel, the

Hebrides and Malin Shelf (200

metre depth contours, red lines

separating shelves) (Huthnance et

al. 2009).

The area is subject to strong seasonal forcing, with surface heating and cooling changing the structure

of the water column. In regions with strong tidal flows (the English Channel and eastern Irish Sea) the

water column is continually mixed, while regions such as the Celtic Sea and Hebridean Shelf

experiences stong seasonal stratification. In the North Channel, between the Irish Sea and the Malin

Shelf there is complete vertical homogeneity to a depth of 200 metres, due to a strong tidal current

of 1.5 ms-1 at spring tides (Simpson 1998).

The Celtic Sea

The Celtic Sea is a 500 kilometre wide (approximately), 100-200 metre deep shelf sea with a highly

dynamical environment (Fig. 6)(Huthnance et al. 2009; Green et al. 2008). It has a large tidal energy

input originating from the Atlantic Ocean. It is characterised by strong tidal currents, which are known

to be the dominant source of mechanical energy (Simpson 1998). The area is subject to strong

seasonal variations in surface heating and cooling. Freshwater supply to the area is limited, meaning

stratification is dominated by temperature (Green et al. 2008). This stratification becomes established

over summer, where buoyancy input out-competes stirring by wind and tidal stirring.

At the shelf edge, internal tides generate internal waves due to the forcing of the barotropic tide. It

causes the water column up and down the steep slope, generating waves (Fig. 4)(Sharples et al. 2007).

This also produces a baroclinic energy flux (Green et al. 2008). These internal waves cause mixing and


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diffusion across the thermocline, bringing cooler, deeper water to the near surface. This water is then

exposed by wind mixing (Fig. 1). These processes lead to a large ocean-shelf exchange of 3 m2s-1

(Huthnance et al. 2009).

North Atlantic water forms a poleward slope current, that is warm and saline, flowing along the

continental slope from Portugal, past Ireland to (Cooper 1952). This barotropic current is centred at

approximately 500 metres on the slope (Cooper 1949; Pingree & Le Cann 1989; Huthnance et al. 2009).

The depth of the slope current is suggested to be forced by the dynamic height of warmer subtropical

waters (Huthnance 1984). Below this current is the bottom Ekman layer, where the current is reduced

to zero, due to friction. Off-shore Ekman transport in the region is believed to be of the order of 1

m2s-1 (Huthnance 1995).

In the Celtic Sea, low-frequency circulation is generally weak, except at the upper slope, and when

channelled through topographic features (e.g. canyons). This localised exchange is equal to the slope

current transport (of order of 1 Sv.) (Huthnance et al. 2009). The discontinuous coastline allows wind-

driven flow eastward through the English Channel (0.1 Sv.: Prandle et al. (1996)) and northward

through the Irish Sea (0.1 Sv.: Knight & Howarth (1999)). At the shelf edge, tidal currents exceed 0.5

ms-1, with tidally reflected flow reaching 0.1 ms-1 (Huthnance et al. 2009). Internal tides are particularly

strong in the region of the shelf edge. At spring tide they have wavelengths >50 metres, and propagate

as decaying waves both off- and on-slope. This means that off-shore exchange occurs in wave-form of

approximately 1.3 m2s-1 (Huthnance et al. 2001).

The enhanced mixing that is associated with internal waves, causes a flux of nutrients and chlorophyll

over the thermocline, associated with a cooler band of water (Sharples et al. 2007; Green et al. 2008;

Huthnance et al. 2009). This flux of nutrients allows the subsurface chlorophyll maximum to be

sustained. Primary production is therefore localised. It has also been documented that there is an

increase in chlorophyll near seamounts and banks, possibly due to intensified mixing (Green et al.

2008).


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Fig. 7. Satellite imagery showing the average sea surface temperatures (left), and chlorophyll concentrations

(right) in the Celtic Sea, between 3rd July and 19th July 2004 (Green et al. 2008).

Historical surveys (Cooper 1949) of the Celtic Sea show that there are cascading-favourable conditions

over the shelf in winter, and traces of cascades at the continental slope. In simplest terms cascading

can be split into two locations. Both the locations in Fig. 8 have the same initial salinity, temperature

and density. During the onset of winter, the water columns are cooled, with location A cooling faster,

due to the reduced depth. This causes higher density water further up the slope, causing it to ‘cascade’

down past location B. When looking at an actual continental slope, the same concept can be used. In

Fig. 9 there are three locations; the shelf, the shelf break and the slope. Both locations A and B cool

faster than C, with A cooling faster than B, causing both to cascade down the slope due to higher

densities (Cooper 1949). Ivanov et al. (2004) also determined that there were two distinct areas of

dense water cascading in the Celtic Sea: in the South Celtic Sea and over the adjoining Armorican Shelf.

Fig. 8. Simplised diagram showing the basis of water

cascading at two locations: A and B (twice as deep)

(Cooper 1949).

Fig. 9. Simplised diagram of dense water cascading at

the continental slope (Cooper 1949).


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Canyons

Submarine canyons are abundant features along continental shelves to deep ocean basins. Shepard &

Dill (1966) mapped just 96 major submarine canyons globally. Using modern satellite bathymetry data,

more than 660 submarine canyons have been mapped globally (Fig. 10)(De Leo et al. 2010). They are

areas of enhanced cross-shelf exchange, primarily because they are regions of large Rossby numbers.

In such regions, the effects of planetary rotation are secondary to the effects of advection of

momentum. Because canyons are much smaller than the surrounding slope, the Rossby number gets

increasingly larger (Wåhlin 2002; Allen 2004).

Fig. 10. Global distribution of submarine canyons: red (circles) named, white un-named, and yellow and orange

are from studies (De Leo et al. 2010).

Canyons are features with complex topography, hydrography, flow, and sediment transport and

accumulation. Because of these complexities, they are known for their distinctive characteristics, such

as accelerated currents, enhanced mixing, and dense water cascades. These can be forced by

topographic or climatic forcings (Klinck 1996; Mulder et al. 2012). Submarine canyons are important

features for physical and biogeochemical reasons. They have been observed to be hotspots for

enhanced upwelling and downwelling, allowing for an increased exchange between shelf waters and

open ocean (Allen & Durrieu de Madron 2009; Allen & Hickey 2010).

Geostrophic flows cannot cross isobaths, restricting cross-shelf flows, forcing along-slope continental

slope flows. Due to geostrophy, this causes an across-slope pressure gradient (Allen & Durrieu de

Madron 2009). In a submarine canyon, flow cannot be along-slope due to the restrictions of the

canyon walls. This means that the Coriolis force cannot balance the pressure gradient force allowing

for flow down the pressure gradient (Freeland & Denman 1982). Therefore, flow is dominated by the


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pressure gradient at the canyon rim. There is no direct impact on the near surface flow. Canyon flow

can be broken down into two types (Allen & Durrieu de Madron 2009):

1. A wind-driven shelf-break or slope current, with the strongest effect felt at the canyon rim;

2. On-shelf deep water formation with an equally strong cross-slope pressure gradient.

However, this water cascades deep into the canyon, making it independent of the wind-driven

flows.

A typical upwelling or downwelling event can be divided into three main phases (Allen & Durrieu de

Madron 2009):

1. An initial transient phase;

2. A near steady advection-dominated phase;

3. A relaxation phase.

The first phase, the initial transient phase is a time-dependent response, as the shelf-break flow

increases. It is generally quite strong and occurs quickly, normally within an inertial period (Allen &

Durrieu de Madron 2009). If there is a steady wind, causing the along-current to continue, density

advection within the canyon reduces the time dependent upwelling after about five days (She & Klinck

2000). It is essentially linear, with similar responses for both positive and negative flows (see Fig. 11).

The second phase, the advection-dominated phase, is not linear, and therefore more complicated. It

is dependent on the canyon topography and flow strength. This phase occurs when the shelf-break

flow is reasonably steady. In this phase, upwelling is generally stronger than downwelling (She & Klinck

2000). Upwelling is driven by negative flows (Fig. 11), thus opposing the shelf waves and arresting

them, leading to strong across isobaths flow. Downwelling is driven by positive flows, moving in the

same direction as wave propagation, allowing along-isobath flows to be established around the

canyon and onto the shelf (Allen & Durrieu de Madron 2009).

Oscillatory flows over the canyon have been suggested to create mean flow over the canyon, due to

the asymmetry of upwelling and downwelling. In the positive phase, the flow leaves the canyon via

the downstream wall, having diverged from the upstream wall (Fig. 11). In the negative phase, flow

follows the upstream wall into the canyon, and leaves via the downstream wall (Allen & Durrieu de

Madron 2009).

In the final phase, the relaxation phase, shelf-break flow reduces (Allen & Durrieu de Madron 2009).

Hickey (1997) suggested that upwelled water leaves the canyon laterally in this phase, rather than

horizontally.


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Fig. 11. Schematic showing the flow of

water through a submarine canyon. The

negative phase (upwelling conditions) is

flow in the opposite direction of Kelvin

wave propagation, while the positive

phase (downwelling conditions) is in the

same direction as wave propagation

(Edited: Allen & Durrieu de Madron

2009).

Upwelling

Submarine canyons are important mechanisms for coastal upwelling, with a high concentration of

zooplankton seen around them (Allen et al. 2001). However, there is a difference in upwelling between

short canyons and long canyons. Short canyons are those that the head of the canyon reaches the

continental slope, long before it reaches the coastline, for example, Astoria (off the west coast of the

USA) and Barkley canyons (west of Vancouver Island)(Hickey 1997; Allen 2000). In a long canyon, the

head of the canyon does not reach the continental slope before the coastline, but rather it extends

far into the coastal region, usually into an estuary (Hickey 1995; Allen 2000). Examples of long canyons

are Juan de Fuca, Mackenzie and Monterey canyons (Waterhouse et al. 2009). Flow in short canyons

has been well studied and documented, while long canyons remain largely unstudied.

In short canyons, as a geostrophic flow passes over the canyon, water is driven up the canyon (Fig.

12). This occurs due to a pressure gradient imbalance caused by restrictions in the topography

(Freeland & Denman 1982). This imbalance is what causes enhanced mixing and upwelling (Hickey

1995). Water columns, originating upstream of the canyon, flow over the top, becoming stretched.

This is due to an increase in bottom depth downstream of the canyon rim. This stretching creates a

cyclonic vorticity in the flow (Hickey 1997; Allen 2004). This has been linked to flow separation at the

mouth of the canyon, which is then advected into the canyon. The flow then turns towards the canyon

head and is advected onto the shelf. Due to the vortex stretching, a cyclonic eddy is formed from the

shelf break down to a depth in the canyon mouth (She & Klinck 2000). Flow above the canyon (<100

metres) does not feel the effects of the canyon, except for a possible elevation of isopycnals (Hickey

1997; Allen 2004; Waterhouse et al. 2009).


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Fig. 12. Schematic showing the processes which lead to upwelling within a short submarine canyon (Allen &

Hickey 2010).

In long canyons, upwelling has been observed to happen throughout the canyon, in all regions (Allen

et al. 2001; Waterhouse et al. 2009). Flow dynamics in long canyons are dependent on Rossby

numbers, time-dependence and advection. Some of the flow characteristics are similar to that of short

canyons, while some are different (Skliris et al. 2001). Waterhouse et al. (2009) used modelling to

determine the effects of changing Rossby numbers on a long canyon, such as the Juan de Fuca.

When low Rossby numbers were input into the model it was determined that there were two

characteristic stages of flow, due to the restriction of isobath convergence (Waterhouse et al. 2009):

1. The generation of vorticity through isopycnals stretching on the upstream side of the canyon,

upwelling occurring downstream of the mouth rim, a slow cyclonic flow within the canyon

walls, and a slowing of flow on the shelf, upstream of the canyon.

2. The formation of an eddy at the canyon mouth, continuation of the flow cyclonic flow in the

earlier stage, as well as the continuation of upwelling at the downstream rim of the mouth. In

this stage the vorticity of the canyon mouth eddy was dependent on stratification. However,

the upstream flow within the canyon was not dependent on stratification.

When a moderate Rossby number was used, it was observed that the pattern of upwelling was

different between short and long canyons. In long canyons, upwelling occurred at the mouth of the

canyon, while in short canyons it occurs through the head of the canyon and at the downstream rim.

At high Rossby numbers, there were similar upwelling processes between the short and long canyons.

Both showed advective regimes, with upwelling at the head of the canyons. As the Rossby number


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decreases in short canyons, upwelling tends to occur in the first phase. Therefore, it was determined

that upwelling was dependent on flow strength and a high Rossby number. In long canyons, upwelling

occurred during all Rossby number simulations, even during quasi-ready conditions. This is due to high

isobath convergence over the canyon (Waterhouse et al. 2009; Allen 2000).

Dense Water Cascading

As discussed previously, dense water cascades form on the shelf, where water is cooled and cascades

down the slope (Fig. 9). Submarine canyons have been documented to be conduits for this process

(Allen & Durrieu de Madron 2009). Dense water (DW) cascading contributes to the ventilation of

intermediate and deep water of the open ocean, which has a substantial impact on biogeochemical

cycles. The effect of canyons on DW cascades varies with the length, width (Wåhlin 2004) and

orientation (Chapman 2000) of the canyon, as well as the topographic features present (Wåhlin et al.

2008). In a uniformly sloping shelf with a canyon cutting into it, a portion of DW will cascade into the

canyon, forming a plume, flowing offshore along the canyon axis, to the right side of the canyon. The

formation of eddies, due to a density front, have been documented to slump into the canyon,

disrupting this DW plume (Chapman & Gawarkiewicz 1995). Wåhlin (2002) looked at the steering

influence of canyons on DW cascades. Dependent on the magnitude of the along-slope current, DW

was observed to cascade through the canyon, with an accumulation of this DW in the canyon.

Wåhlin (2004) looked at the influence of length and width of canyons on DW cascading. It was

determined that the transport capacity of deep channels was larger than that of shallow channels.

When gently sloping topography was used in the model, there was a maximum downward flow

through a wide canyon (>10 kilometres), however, steeper regions were the most active, when the

canyon was a few kilometres wide. Wåhlin et al. (2008) looked at the influence of the overall shape of

the canyon (v-shaped) with respect to different flow regimes and topographic features. They

determined that small scale topography has a bigger influence on mixing than large scale topography.

Chapman (2000) looked at three different canyon orientations: normal, diagonal and parallel to the

shelf. It was found that little DW enters the normal and diagonal canyons. This is due to the fact that

along-slope flow follows isobaths, and cannot flow down into the canyon. In regards to the parallel

canyon, there was a higher portion of the flow in this canyon. The amount of DW in the canyon was

dependent on the rate of flow over the canyon. A slower along-slope flow meant that there was more

DW in the canyon, due to the reduced speed and therefore increased meandering. It was also found

that DW cascading has shown to induce localised upwelling of deep water onto the shelf (Kämpf 2005).


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Internal Tides and Mixing

Upwelling and dense water cascading are both processes of advection, which cause the movement of

water through the canyon. Another process of transport is the mixing of deep water, due to tides.

Submarine canyons act as conduits for deep ocean water further onto the shelf then if it were of

uniform length. This means that the head of canyons is an area of enhanced tidal mixing (Allen &

Durrieu de Madron 2009).

Long canyons (as discussed above) are particularly strong areas of enhanced tidal mixing. This is due

to them stretching from the shelf break, along the slope and up to large estuaries. Deep ocean water

circulates in these estuaries, and with the axis and head of the canyon being subjected to large tidal

currents, mixing is enhanced. Long canyons are therefore areas of strong nitrate concentrations (Allen

& Durrieu de Madron 2009).

In long canyons that do not reach into large estuaries, tidal currents do not penetrate into the canyon,

but rather across the canyon, parallel to the shelf-break (Allen & Durrieu de Madron 2009). It has been

documented that many of these canyons are areas of very high internal tidal energy, such as

Hydrography Canyon (Wunsch & Webb 1979) and Monterey Canyon (Kunze et al. 2002). They are

areas of enhanced tidal and internal wave energy due to focusing (Wunsch & Webb 1979) or due to

them being large regions of critical slope (Hotchkiss & Wunsch 1982). It has also been suggested that

small scale roughness causes enhanced mixing (Wåhlin et al. 2008) and internal tidal energy (Kunze et

al. 2002). Diffusivity values in such canyons are very large at the canyon axis (0.05 m2s-1), with the

canyon rim have values just a factor of ten smaller, recorded in the Monterey Canyon.

Conclusion

Shelf seas are important regions for both physical and biogeochemical reasons. They are regions

where the exchange of open ocean and shelf water occurs. At the shelf break, a poleward (in the

northern hemisphere, eastern side of basin) along-isobath slope current flows as rigid water columns,

unable to change its depth (Taylor 1923). Because of this rigidity, exchange between the open ocean

and shelf sea is limited. Wind stress drives Ekman transport offshore, allowing for deeper water to be

upwelled close to the coast, bringing with it nutrient rich water (Ekman 1905). Shelf seas are areas of

high tidal energy. This energy enhances mixing at the shelf edge due to the formation of internal waves

(Green et al. 2008), which have been observed to move onshore. Shelf seas, such as the Irish Sea,

become seasonally stratified during the spring and summer months, due to an increase in buoyancy,

which out-competes stirring by wind and tides (van Aken 1986). Tidal mixing fronts form where

stratified water meets mixed water. During spring, the water column becomes stratified due to an

increase in temperature at the surface. As the light determined critical depth descends below the


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thermocline (due to an increase in day length) more nutrients become available, leading to the spring

bloom. It varies locally and inter-annually due to freshwater and sediment inputs.

The Northwest European Shelf is characterised by its complex topography, due to the presence of the

British Isles. It is approximately 2000 kilometres long, from Portugal to Norway. The area is known for

its strong tidal forcing, with the M2 tide causing an increase in frictional stress at the seabed, causing

much mixing. The region is subject to strong seasonal forcings, being heated in spring and summer,

and cooled in winter. The Celtic Sea is a region located within the Northwest European Shelf. It is 500

kilometres wide and 100-200 metres deep. It has a large tidal energy range, originating from the North

Atlantic, which is the dominant mechanical energy input (Simpson 1998). Stratification in the region

is dominated by temperature, which is established during summer. A barotropic poleward current is

centred around 500 metres, made up of North Atlantic water (Cooper 1949). Ocean-shelf exchange in

the region is approximately 3 m2s-1(Huthnance et al. 2009). The region has also been documented as

having favourable conditions for dense water cascades, in winter (Cooper 1949).

Submarine canyons are regions of enhanced upwelling and downwelling on the continental shelf.

There have been more than 660 mapped globally using satellite bathymetry (De Leo et al. 2010). They

are characterised as regions with complex topography, flow and hydrography, and large Rossby

numbers. They are known to have two forms of water transport: advection (upwelling/downwelling

and dense water cascades) and mixing (internal tides). Upwelling/downwelling has three phases: the

first is an initial transient phase when shelf break flow increases, the second is an advection-

dominated phase during steady state, and the third is a relaxation phase, when flow begins to slow.

Upwelling occurs predominantly in the first two phases (Allen & Durrieu de Madron 2009).

In short canyons, there is a pressure gradient imbalanced caused by topographic restrictions, which

allows for upwelling to occur. Cyclonic vorticity is a dominate feature in short canyons (Hickey 1997;

Allen 2004). In long canyons, which protrude onto the shelf, reaching a large estuary, different Rossby

numbers allow for different flow regimes. When a moderate Rossby number is considered upwelling

occurs at the mouth of the canyon, whereas in short canyons it occurs at the head and downstream

of the rim. When a high Rossby number is considered the flow regime is similar to that of the short

canyon. Unlike short canyons, upwelling/downwelling can occur during any of the three phases in a

long canyon (Waterhouse et al. 2009).

Dense water cascading has been observed to occur through canyons. The magnitude of the cascade is

dependent on the width, length (Wåhlin 2004), orientation (Chapman 2000) and topographic features

(Wåhlin et al. 2008) of the canyon. Canyons which are orientated parallel to the coastline are believed

to allow higher volumes of water to pass through them, forming a plume. The speed of the along-

slope current also affects the amount of dense water cascading through the canyon. Slower currents


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will allow for more to cascade through the canyon, while faster currents will flow directly passed it

(Chapman 2000).

Mixing, due to internal waves and tidal energy, has also been suggested as a mechanism for water

transport through a canyon. In long canyons, where the head reaches into an estuary, there are large

tidal currents, allowing for enhanced mixing, bringing a high level of nutrients to the region. In long

canyons which do not extend up to the coastline internal tides are the dominant mixing force, due to

focusing and a large critical slope. In these canyons, diffusivity values are of order 0.05 m2s-1 at the

axis (Allen & Durrieu de Madron 2009).

References

Allen, S.E., 2000. On subinertial flow in submarine canyons: Effect of geometry. Journal of Geophysical Research, 105, pp.1285–1297.

Allen, S.E. et al., 2001. Physical and biological processes over a submarine canyon during an upwelling event. Canadian Journal of Fisheries and Aquatic Sciences, 58, pp.671–684.

Allen, S.E., 2004. Restrictions on deep flow across the shelf-break and the role of submarine canyons in facilitating such flow. Surveys in Geophysics, 25, pp.221–247.

Allen, S.E. & Durrieu de Madron, X., 2009. A review of the role of submarine canyons in deep-ocean exchange with the shelf. Ocean Science Discussions, 6(2), pp.1369–1406.

Allen, S.E. & Hickey, B.M., 2010. Dynamics of advection-driven upwelling over a shelf break submarine canyon. Journal of Geophysical Research: Oceans, 115(March), pp.1–20.

Chapman, D.C., 2000. The influence of an alongshelf current on the formation and offshore transport of dense water from a coastal polynya. Journal of Geophysical Research, 105(C10), p.24007.

Chapman, D. C., & Gawarkiewicz, G., 1995. Offshore transport of dense shelf water in the presence of a submarine canyon. Journal of Geophysical Research: Oceans (1978–2012), 100(C7), 13373-13387.

Cooper, L.H.N., 1949. Cascading over the continental slope of water from the Celtic Sea. , pp.719–750.

Cooper, L.H.N., 1952. The physical and chemical oceanography of the waters bathing the continental slope of the Celtic Sea.

Ekman, V. W., 1905. On the influence of the earth's rotation on ocean currents, Arkiv. Mat, Astron, Fysik, 2, pp.1–52.

Estrade, P. et al., 2008. Cross-shelf structure of coastal upwelling: A two — dimensional extension of Ekman’s theory and a mechanism for inner shelf upwelling shut down. Journal of Marine Research, 66, pp.589–616.

Freeland, H.J. & Denman, K.L., 1982. A topographically controlled upwelling center off southern Vancouver Island. Journal of Marine Research, 40(4), pp.1069–1093.

Green, J. a. M. et al., 2008. Internal waves, baroclinic energy fluxes and mixing at the European shelf edge. Continental Shelf Research, 28, pp.937–950.

Hickey, B.M., 1995. Coastal Submarine Canyons. Topographic Effects in the Ocean. SOEST Special Publications, pp.95–110.


19

Hickey, B.M., 1997. The Response of a Steep-Sided, Narrow Canyon to Time-Variable Wind Forcing. Journal of Physical Oceanography, 27, pp.697–726.

Hotchkiss, F.S. & Wunsch, C., 1982. Internal waves in Hudson Canyon with possible geological implications. Deep Sea Research Part A. Oceanographic Research Papers, 29(4), pp.415–442.

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(95), pp.353–431.

Huthnance, J.M. et al., 2001. Physical structures, advection and mixing in the region of Goban Spur. Deep-Sea Research Part II: Topical Studies in Oceanography, 48(14-15), pp.2979–3021.

Huthnance, J.M., 1984. Slope Currents and “JEBAR.” American Meteorological Society, 14, pp.795–810.

Huthnance, J.M., Holt, J.T. & Wakelin, S.L., 2009. Deep ocean exchange with west-European shelf seas. Ocean Science, 5, pp.621–634.

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(13-14), pp.1449–1472.

Ivanov, V. V. et al., 2004. Cascades of dense water around the world ocean,

Kämpf, J., 2005. Cascading-driven upwelling in submarine canyons at high latitudes. Journal of Geophysical Research C: Oceans, 110, pp.1–10.

Klinck, J.M., 1996. Circulation near submarine canyons: A modeling study. Journal of Geophysical Research, 101(95), p.1211.

Knight, P.J. & Howarth, M.J., 1999. The flow through the north channel of the Irish Sea. Continental Shelf Research, 19(5), pp.693–716.

Kunze, E. et al., 2002. Internal Waves in Monterey Submarine Canyon. Journal of Physical Oceanography, 32(6), pp.1890–1913.

De Leo, F.C. et al., 2010. Submarine canyons: hotspots of benthic biomass and productivity in the deep sea. Proceedings. Biological sciences / The Royal Society, 277(May), pp.2783–2792.

Mulder, T. et al., 2012. Present deep-submarine canyons activity in the Bay of Biscay (NE Atlantic). Marine Geology, 295-298, pp.113–127.

Munk, W. & Wunsch, C., 1998. Abyssal recipes II: Energetics of tidal and wind mixing. Deep-Sea Research Part I: Oceanographic Research Papers, 45(1998), pp.1977–2010.

Pingree, R.D. & Le Cann, B., 1989. Celtic and Armorican slope and shelf residual currents. Progress in Oceanography, 23, pp.303–338.

Pond, S. & Pickard, G. L., 1983. Introductory Dynamical Oceanography, Pergamon Press, 329.

Prandle, D. et al., 1996. Combining modelling and monitoring to determine fluxes of water, dissolved and particulate metals through the Dover Strait. Continental Shelf Research, 16(2), pp.237–257.

Prowe, A.E.F. et al., 2009. Mechanisms controlling the air-sea CO2 flux in the North Sea. Continental Shelf Research, 29(15), pp.1801–1808.

Rees, A.P., Joint, I. & Donald, K.M., 1999. Early spring bloom phytoplankton-nutrient dynamics at the Celtic Sea shelf edge. Deep-Sea Research Part I: Oceanographic Research Papers, 46, pp.483–510.

Sharples, J. et al., 2007. Spring-neap modulation of internal tide mixing and vertical nitrate fluxes at a shelf edge in summer. Limnology and Oceanography, 52(5), pp.1735–1747.


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She, J. & Klinck, J.M., 2000. Flow near submarine canyons driven by constant winds. Journal of Geophysical Research, 105(C12), pp.28671–28694.

Shepard, F. P. & Dill, R. F., 1966. Submarine canyons and other sea valleys. Rand McNally.

Simpson, J. H., 1998. The Celtic Seas coastal segment. In: Brink, K. H. & Robinson, A. R. (Eds.), The Sea, 11, Wiley, New York, pp.659-698.

Simpson, J.H. & Bowers, D., 1981. Models of stratification and frontal movement in shelf seas. Deep Sea Research Part A. Oceanographic Research Papers, 28(7), pp.727–738.

Skliris, N. et al., 2001. Shelf-slope exchanges associated with a steep submarine canyon off Calvi (Corsica, NW Mediterranean Sea): A modeling approach. Journal of Geophysical Research, 106(C9), p.19883.

Taylor, G.I., 1923. Stability of a viscous liquid contained between two rotating cylinders. Proc. Roy. Soc. A, 223, pp.289–343.

Thomas, H. et al., 2004. Enhanced open ocean storage of CO2 from shelf sea pumping. Science (New York, N.Y.), 304(5673), pp.1005–1008.

van Aken, H.M., 1986. The onset of seasonal stratification in shelf seas due to differential advection in the presence of a salinity gradient. Continental Shelf Research, 5(4), pp.475–485.

Wåhlin, a. K., 2004. Downward channeling of dense water in topographic corrugations. Deep-Sea Research Part I: Oceanographic Research Papers, 51(4), pp.577–590.

Wåhlin, a. K. et al., 2008. Laboratory observations of enhanced entrainment in dense overflows in the presence of submarine canyons and ridges. Deep-Sea Research Part I: Oceanographic Research Papers, 55(6), pp.737–750.

Wåhlin, a. K., 2002. Topographic steering of dense currents with application to submarine canyons. Deep-Sea Research Part I: Oceanographic Research Papers, 49, pp.305–320.

Waterhouse, A.F., Allen, S.E. & Bowie, A.W., 2009. Upwelling flow dynamics in long canyons at low Rossby number. Journal of Geophysical Research: Oceans, 114(5), pp.1–18.

Wenju, C. & Lennon, G.W., 1988. Upwelling in the Taiwan Strait in response to wind stress, ocean circulation and topography. Estuarine, Coastal and Shelf Science, 26, pp.15–31.

Wunsch, C. & Webb, S., 1979. The Climatology of Deep Ocean Internal Waves. Journal of Physical Oceanography, 9(2), pp.235–243.

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