Any ecological geography of the ocean must be sensitive to the physical processes that determine when and where phytoplankton growth will occur, and which taxa will parti- cipate. From Sverdrup, we may assume that what we need to analyze are seasonal changes in (i) the light field and (ii) stratification within the water column. Apparently, these are the consequences of solar irradiance and of wind stress at the sea surface; however, both are strongly modified by continental geography and the distribution of land masses, and each also responds characteristically to latitude. It will be useful at this point to review the consequences of latitude for irradiance at the sea surface and on the processes that are induced by wind stress at the sea surface. The first are easily stated, the second are much more complex.
The regional, seasonal characteristics of irradiance at the sea surface are readily encap- sulated in two relatively simple propositions: (i) in low latitudes the seasonal cycle has a smaller amplitude than the diel cycle, whereas the annual cycle exceeds the diel at
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high latitudes (Hastenrath, 1985) and (ii) the irradiance field is dominated by a simple and predictable meridional gradient, which migrates with the seasons and is modi- fied by regionally variable cloudiness and atmospheric clarity that respond to processes above adjacent continents. At this point in the discussion, that is all that needs to be said.
Prominent in any textbook on biological oceanography are the ecological consequences of the balance between turbulence and mixing caused by wind stress at the surface and stratification induced by solar heating. That this balance is relevant only at higher latitudes is rarely mentioned, and the student may be left to discover that the physics required by the Sverdrup model is quite different at low latitudes. There, seasonal changes in depth of the surface mixed layer are due not to seasonal changes in turbulence induced by varying local wind stress but rather, through geostrophic adjustment, to changes in circulation that may occur in distant regions (Philander, 1985; Katz, 1987; Hastenrath and Merle, 1987). To understand this it is necessary to start with the essential, but often overlooked, difference between the effects of wind stress at the sea surface at high and low latitudes.
Simply stated, wind stress is translated mostly into potential energy at high latitudes, but into kinetic energy at low latitudes. This is not the place for a formal discussion of this effect, for which a good modern source is Tomczak and Godfrey (1994), but, simply stated, flow is induced by the pressure-gradient force at the sea surface, and because geostrophic balance is maintained between this gradient and the Coriolis parameter induced by Earth’s rotation (f, with dimensions of 1/sec), the slope of the sea surface required to produce motion diminishes equatorward as a direct function of latitude, and without discontinuity (Lighthill, 1969; Philander, 1985). It is perhaps characteristic of the evolution of oceanographic theory that this was not understood until the 1960s.
Consequently, seasonal moderation of wind stress over the North Atlantic in summer has a negligible effect on the flow of the Gulf Stream, whereas seasonal reversal of monsoon winds in the Arabian Sea reverses the Somali Current to almost 1000 m depth in just a few weeks. Only after reversal of the North Atlantic westerly winds for about 10 years would the same effect be produced on the Gulf Stream. At the extreme, geostrophic balance ceases to function at less than about 2 from the equator, where motion is determined by nonrotational fluid dynamics; the most striking manifestation of this effect is the undercurrent flowing eastward in the pycnocline below the equator.
That seasonal changes in wind stress should quickly modify the flow in major barotropic currents at low latitudes, such as that along the coast of Somalia and Arabia, is a fundamental characteristic of tropical oceans; moreover, since the spatial scale of the adjustment of barotropic currents is similar at all latitudes, seasonal geostrophic adjust- ments in the mixed-layer depths can be identified across whole ocean basins and have important consequences when considering the Sverdrup model. This, of course, is usually invoked where it is changes in local wind mixing and irradiance that force changes in mixed-layer depth, although it has also been used to interpret phytoplankton seasonality in lower latitudes (Obataet al., 1996). This investigation of the interaction between rel- ative changes in mixed-layer depth and Zcr and the concomitant increase or decrease of surface chlorophyll used Levitus mixed-layer depth (Zm), CZCS surface chlorophyll, and the surface light field. The results suggested that whenever and wherever Zm shoals up through Zcr, whether by changes in wind stress or by geostrophic adjustment, a bloom follows. The reverse effect was also simulated: whenever and wherever Zmbecame deeper than Zcr, a decrease in surface chlorophyll was produced in the following month.
Also consequent upon the equatorward vanishing of the Coriolis parameter are changes in the scale taken by internal wave trains, implicated in wind mixing of the surface layers of the ocean. The period of these waves changes from only 12 hours at the poles to much longer at the equator (Garrett, 2003). Their propagation into the interior of the ocean
Stratification and Irradiance: The Consequences of Latitude 57
induces a cascade to smaller scales and to ever-smaller eddying and, hence, mixing. This process becomes less energetic as a function of latitude, with a slower cascade toward the equator. Thus, wind-induced turbulence and mixing is—other things being equal—
a direct function of latitude and another reason for the physical stability of tropical oceans.
The final consequence of latitudinal change in the Coriolis parameter is that the values taken by the Rossby radii of deformation are latitude-dependent (Emery et al., 1984;
Houryet al., 1987). The external radius (Re) determines the length scales for barotropic (geostrophically balanced) phenomena; it is the length scale over which gravitational phenomena balance the tendency of f to deform a surface and may be computed as (gH/f), where H is bottom depth and g is acceleration due to gravity. The internal radius (Ri) also depends on stratification and determines the horizontal dimensions of quasigeostrophic features—mesoscale eddies and Rossby waves. It may be approximated as (1/first eigenvalue N)×f and, because f vanishes at the equator, Ri here takes a value of infinity. At 50 latitude, Ri takes values of only 25 km whereas at 5 latitude, values are around 300 km. For this reason, we may expect larger and fewer mesoscale eddies as we approach the equator.
There are also latitudinal differences in the density gradients (and thus in their resis- tance to wind mixing) between the summer pycnocline of midlatitudes and the permanent shallow pycnocline of low latitudes; the latter is steeper and has greater resistance to mixing, thus enhancing the effects discussed earlier. Furthermore, the heat balance at the surface of low-latitude oceans is such that there is a mean downward heat flux across the sea surface that is balanced by horizontal transport to higher latitudes; seasonal changes in incoming irradiance here are sufficiently small that at no season is there sufficient loss of heat to produce convective deepening of the permanent tropical thermocline. Finally, there is often an excess of precipitation over evaporation at the surface of the tropical ocean, so mixed-layer water at low latitudes is characteristically not only warmer but also a little fresher than deeper water. The tropical Ekman layer, therefore, lies above a pycn- ocline having greater stability and resistance to mixing than elsewhere; this is expressed as a relatively high subsurface maximum of the Brunt-Väisälä buoyancy frequency (N, in cycles hr−1; see later discussion).
Finally, consider also the effect of the seasonal increase in the westward stress of the trade winds across the tropical ocean that occurs within about 15–20 latitude of the equator. It is this stress that maintains an upward slope of the sea surface and downward slope of the pycnocline to the west in each ocean. The adjustment time of an ocean basin to wind stress is related to the time taken for planetary waves to propagate across the ocean, so an equilibrium seasonal response cannot occur within a single season across the great width of the tropical Pacific (15,000 km). Additionally, over this great distance, wind stress in the west is not in phase with that in the east, so the response of the ocean cannot be simple; consequently, it is only when wind stress changes for longer periods, on the interannual El Niño–Southern Oscillation scale, that an equilibrium response in Pacific circulation can occur. However, the tropical Atlantic (just 5000 km wide at the equator) responds to the seasonal cycle of trade wind stress with a seasonal basin-wide geostrophic adjustment of mixed-layer depth that maintains equilibrium.
The aspects of ocean physics discussed in this section serve to distinguish very clearly the characteristics of low- and high-latitude oceans as biological habitats. But they do not suggest the discontinuities we might like to have in order to partition the apparent continuum forced by the continuous vanishing of the Coriolis parameter equatorward.
However, if we look more closely at one feature from the previous discussion—relative pycnocline stability, the property that most closely controls resistance of the water column to vertical mixing—then I believe that the required discontinuities can be demonstrated.
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