Editorial Type:
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Article Type:
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On Geometrical Aspects of Interior Ocean Mixing

Trevor J. McDougall School of Mathematics and Statistics, University of New South Wales, Sydney, New South Wales, Australia

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Sjoerd Groeskamp Institute for Marine and Atmospheric Studies, University of Tasmania, and CSIRO Marine and Atmospheric Research, Castray Esplanade, Hobart, Tasmania, Australia

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Stephen M. Griffies NOAA/Geophysical Fluid Dynamics Laboratory, Princeton, New Jersey

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Print Publication:
01 Aug 2014
DOI:
https://doi.org/10.1175/JPO-D-13-0270.1
Page(s):
2164–2175
Full access

Abstract

The small-slope approximation to the full three-dimensional diffusion tensor of epineutral diffusion gives exactly the same tracer flux as the commonly used projected nonorthogonal diffusive flux of layered ocean models and of theoretical studies. The epineutral diffusion achieved by this small-slope approximation is not exactly in the direction of the correct epineutral tracer gradient. That is, the use of the small-slope approximation leads to a very small flux of tracer in a direction in which there is no epineutral gradient of tracer. For (the tracer) temperature or salinity, the difference between the correct epineutral gradient and the small-slope approximation to it is proportional to neutral helicity. The authors also make the point that small-scale turbulent mixing processes act to diffuse tracers isotropically (i.e., the same in each spatial direction) and hence it is strictly a misnomer to call this process “dianeutral diffusion” or “vertical diffusion.” This realization also has implications for the diffusion tensor.

Corresponding author address: Trevor J. McDougall, School of Mathematics and Statistics, University of New South Wales, NSW 2052, Australia. E-mail: trevor.mcdougall@unsw.edu.au

Abstract

The small-slope approximation to the full three-dimensional diffusion tensor of epineutral diffusion gives exactly the same tracer flux as the commonly used projected nonorthogonal diffusive flux of layered ocean models and of theoretical studies. The epineutral diffusion achieved by this small-slope approximation is not exactly in the direction of the correct epineutral tracer gradient. That is, the use of the small-slope approximation leads to a very small flux of tracer in a direction in which there is no epineutral gradient of tracer. For (the tracer) temperature or salinity, the difference between the correct epineutral gradient and the small-slope approximation to it is proportional to neutral helicity. The authors also make the point that small-scale turbulent mixing processes act to diffuse tracers isotropically (i.e., the same in each spatial direction) and hence it is strictly a misnomer to call this process “dianeutral diffusion” or “vertical diffusion.” This realization also has implications for the diffusion tensor.

Corresponding author address: Trevor J. McDougall, School of Mathematics and Statistics, University of New South Wales, NSW 2052, Australia. E-mail: trevor.mcdougall@unsw.edu.au
Keywords: Models and modeling; Grid systems; Isopycnal coordinates; Parameterization; Subgrid-scale processes

1. Introduction

The idea that ocean properties are advected and mixed predominantly along “isopycnal” surfaces dates back at least to Iselin (1939), who noted the similarity between the salinity–temperature diagram of vertical casts in the center of a subtropical gyre and that plotted from data in the winter surface mixed layer in the whole hemisphere. This concept of “isopycnal mixing” has become part of the underpinning assumptions of physical oceanography. Papers by Veronis (1975), Solomon (1971), Redi (1982), and McDougall and Church (1986) pointed out that ocean models needed to rotate their diffusion tensor to be aligned with the locally referenced potential density surface in order to avoid the ill effects of having density mixed horizontally at fixed depth.

a. A physical argument for orienting lateral mixing along neutral tangent planes

While the strong lateral mixing achieved by the energetic mesoscale eddies is widely believed to be oriented along isopycnals (McDougall 1987a), we are aware of only one convincing argument that supports this assumption; the argument has been made in section 7.2 of Griffies (2004) and in section 2 of McDougall and Jackett (2005) and is explained here with the aid of Fig. 1. The argument involves the rather small amount of dissipation of turbulent kinetic energy that is measured in the ocean interior.

Fig. 1.
Fig. 1.

Sketch of a central seawater parcel being moved adiabatically and without change in its salinity to either the right or the left of its original position in a direction that is not neutral. When the parcel is then released it feels a vertical buoyant force and begins to move vertically (upward on the left and downward on the right) toward its original “isopycnal.”

Citation: Journal of Physical Oceanography 44, 8; 10.1175/JPO-D-13-0270.1

We begin by initially adopting the counterargument, so that we take the lateral mesoscale dispersion to occur along a surface that differs in slope from the neutral tangent plane. Individual fluid parcels are then transported above and below the neutral tangent plane and would need to subsequently rise or sink in order to regain a vertical position of neutral buoyancy. This situation is illustrated in Fig. 1 where a central seawater parcel is moved adiabatically and without exchange of salinity in a nonneutral direction to either the left or right and then released. The fluid parcel then feels a vertical buoyant force (upward if displaced to the left and downward if displaced to the right) and it begins to move back to its original “density” surface.

This vertical motion would either

  1. involve no small-scale turbulent mixing, in which case the combined process is adiabatic and isohaline and so is equivalent to epineutral dispersion, or

  2. the sinking and rising parcels would mix and entrain in a plumelike fashion with the ocean environment, and therefore experience irreversible mixing.

If the second case were to happen, the dissipation of mechanical energy associated with the dianeutral mixing would be observed. But in fact mechanical energy dissipation in the main thermocline is consistent with a dianeutral diffusivity of only (MacKinnon et al. 2013). This relatively small value of the dianeutral (vertical) diffusivity has been confirmed by purposely released tracer experiments (e.g., Ledwell et al. 2011).

McDougall and Jackett (2005) showed that even if all the observed dissipation of turbulent kinetic energy were due to the second case above of nonneutral lateral mixing (implying no contribution from breaking internal waves to ), the maximum tangent of the angle between the mixing plane and the neutral tangent plane is of order . Since the dominant cause of the observed dissipation of turbulent kinetic energy in the ocean interior is likely the breaking of internal gravity waves (MacKinnon et al. 2013), no matter whether they are forced by winds, by internal tides, or by lee waves above bottom topography, the allowable angle between the plane of mesoscale mixing and the neutral tangent plane is much less than .

Nycander (2011) has examined mixing in the ocean along inclined planes and has concluded that an exchange between potential and kinetic energies is required to move seawater parcels along a neutral tangent plane. He also concluded that such energetic arguments do not shed light on the question of which mixing direction is preferred by the energetic ocean mesoscale eddies. Nycander’s results confirm the long-standing practice (since the 1980s) of defining the neutral direction using parcel movement arguments in terms of the lack of vertical buoyant restoring forces, rather than in terms of the changes in gravitational potential energy.

We conclude that the only evidence to support the notion that strong lateral mixing is directed along neutral tangent planes is the measured smallness of the dissipation of turbulent kinetic energy in the ocean interior, coupled with the arguments of McDougall and Jackett (2005) and Griffies (2004). Furthermore, this evidence from measured dissipation rates is in accordance with the interpretation of tracer distributions as per isopycnal water mass analysis, with this framework originating from Iselin (1939). The relatively small dissipation of mechanical energy in the ocean interior represents a key distinction from the troposphere, where radiative damping leads to relatively large levels of dissipation and associated diabatic mixing. Consequently, oceanographers have a fundamental reason to be concerned about details of how mixing is oriented. This concern motivates our examination of the geometry of ocean interior mixing.

b. Specification of the neutral tangent plane

To provide a mathematical foundation for later discussions, we present a physical and mathematical review of a neutral tangent plane. Physically, the neutral tangent plane is that plane in space in which a seawater parcel can be moved an infinitesimal distance without being subject to a vertical buoyant restoring force. That is, the neutral tangent plane is the plane of neutral or zero buoyancy.

As a thought experiment, consider the seawater parcel at a point in the ocean and enclose it in an insulating plastic bag. Upon moving to a new location a small distance away, the parcel will experience an increment in pressure . Its in situ density will thus change by , where is the adiabatic and isohaline compressibility. At the same new location the seawater environment surrounding the enclosed parcel has an Absolute Salinity that is different than that at the original location by , a Conservative Temperature difference of , and a density difference of , where and are the appropriate saline contraction and thermal expansion coefficients, defined with respect to Absolute Salinity and Conservative Temperature (IOC et al. 2010; McDougall and Barker 2011). If at the new location the displaced parcel does not feel a buoyant (Archimedean) force, its density must be equal to that of the environment at its new location . That is,
e1
Hence, along an infinitesimal neutral trajectory in the neutral tangent plane, the variations of and of the ocean must obey
e2

This thought experiment is typical of our thinking about turbulent fluxes. We imagine the adiabatic and isohaline movement of fluid parcels and then let these parcels mix molecularly with their surroundings. Central to this way of thinking about turbulent fluxes are the following two properties of the tracer that is being mixed:

  1. It must be a “potential” fluid property, for otherwise its value will change during the displacement even though the displacement is done without exchange of heat or mass, and

  2. it should be close to being a “conservative” fluid property so that when it does mix intimately (i.e., molecularly) with its surroundings, we can be sure that there is negligible production or destruction of the property.

The present paper is concerned with the mixing of tracers by both epineutral mixing and by small-scale mixing processes. The processes we examine are different from the separation of diffusion from advection (or the separation between symmetric and antisymmetric diffusion) that are crucial in the temporal residual mean theory and its parameterization (McDougall and McIntosh 2001; Gent et al. 1995; Griffies 1998). Rather, our focus concerns two basic geometric aspects of ocean interior mixing.

In sections 2 and 3, we show that the use of the projected nonorthogonal coordinate system gives the same diffusive tracer fluxes as the small-slope approximation to the Redi (1982) diffusion tensor. Thereafter, we show that this tracer flux has a component that is nonzero in a direction in which there is no tracer gradient. We show that for temperature and salinity, this unphysical aspect of the small-slope diffusive flux is proportional to neutral helicity. In section 4, we turn from epineutral diffusion to consider the mixing achieved by small-scale mixing processes such as breaking internal gravity waves. As a three-dimensional isotropic diffusion process, small-scale mixing should not be included in ocean mixing parameterizations as a one-dimensional “dianeutral” or “vertical” diffusion. This realization leads to a simplification of the Redi (1982) diffusion tensor.

2. The small-slope approximation to epineutral diffusion

a. The exact epineutral gradient

Using standard three-dimensional geometry (rather than the projected nonorthogonal approach considered in section 2b), we know that the gradient of a scalar in a surface, which in our case is the neutral tangent plane, is
e3ab
In these expressions, is the unit normal to the neutral tangent plane:
e4
The vector is parallel to the unit normal vector , but has unit length in the vertical direction (parallel to gravity), that is,
e5
We define the projected horizontal gradient operator in section 2b, which summarizes the projected nonorthogonal approach. The slope of the neutral tangent plane is discussed in the appendix, where we find the following expression in terms of the horizontal () and vertical gradients of Absolute Salinity and Conservative Temperature:
e6
The last parts of Eq. (4) express in terms of the spatial gradient of locally referenced potential density and of neutral density (Jackett and McDougall 1997). Both and have a positive (i.e., upward) vertical component. Equation (6) has introduced the x and y components and of the vector slope , and we will also use the shorthand notation for the inner product .
The epineutral gradient [Eq. (3)] can be written in tensor form in terms of the Cartesian gradient components as
e7
which introduces the Redi (1982) tensor for epineutral diffusion; the full Redi diffusion tensor also includes small-scale turbulent mixing that we deliberately exclude from discussion until section 4.
The parcel-based definition of the neutral tangent plane, , of Eq. (2), can now be expressed in terms of the (exact) epineutral gradients of and as
e8ab

b. The projected nonorthogonal version of the epineutral gradient

The epineutral gradient that is used in many theoretical oceanographic studies (e.g., McDougall and Jackett 1988; McDougall 1988, 1995) and in layered ocean models (Bleck 1978a,b) is based on the projected nonorthogonal coordinate system first introduced by Starr (1945), which is widely used in geophysical fluid theory and modeling. In this coordinate system the vertical coordinate is strictly vertical (parallel to the effective gravitational force); planes of constant latitude and longitude are strictly vertical cones and planes, respectively; and tracer gradients are calculated with respect to an undulating coordinate surface defined by the layer coordinate. This procedure for computing a tracer gradient is illustrated in Fig. 2. Note particularly that the lateral gradient of a tracer in the undulating isopycnal or neutral density surface is calculated using the difference in tracer values between points b and a on this surface (see Fig. 2b), but the distance increment in the denominator of the projected nonorthogonal gradient is given by , that is, the distance is measured at constant height. Hence, the projected nonorthogonal gradient in a general undulating surface is given by
e9
Importantly, horizontal distances are measured between vertical surfaces of constant latitude x and longitude y, while values of the scalar property are evaluated on the surface (e.g., an isopycnal surface or, in the case of , a neutral tangent plane). Note that has no vertical component; it is not directed along the surface, but rather it points in the horizontal direction and thus is perpendicular to gravity.
Fig. 2.
Fig. 2.

(a) A three-dimensional perspective of the projected nonorthogonal coordinate system that is commonly used in layered ocean models. (b) The projected nonorthogonal gradient of a property in the neutral tangent plane, in the limit as the distances tend to zero, is equal to the tracer difference between points a and b in the above figure, divided by the exactly horizontal distance or .

Citation: Journal of Physical Oceanography 44, 8; 10.1175/JPO-D-13-0270.1

The projected nonorthogonal epineutral tracer gradient can be written in terms of the regular Cartesian gradients (e.g., is the exactly horizontal gradient of ) as (from McDougall and Jackett 1988; McDougall 1995; Griffies 2004)
e10ab
using as defined in Eq. (5). Note that and are each horizontal two-dimensional vectors. In the appendix we prove that the neutral tangent definition of Eq. (8) can be expressed in terms of the corresponding projected nonorthogonal gradients as
e11ab

c. Equivalence of the small-slope approximation and the projected coordinate approach

The three-dimensional tracer gradient based on the projected nonorthogonal approach is
e12
which is the flux in the neutral tangent plane (this property can be checked by showing that the scalar product of Eq. (12) with is zero) whose horizontal component is . The use of the projected nonorthogonal gradient in an ocean model or in a theoretical study is equivalent to using the three-dimensional flux (12) in Cartesian coordinates.
This three-dimensional flux, Eq. (12), can be expressed in tensor form in Eq. (13) below, which we note is the same as the commonly used small-slope approximation [first introduced by Gent and McWilliams (1990)] to the exact epineutral gradient . That is, we have found that the use of the projected nonorthogonal coordinate framework gives the same three-dimensional tracer gradient as the small-slope approximation to the exact epineutral gradient , since
e13

This equivalence between the projected nonorthogonal coordinate version of an epineutral gradient and the small-slope approximation of epineutral diffusion was built into the small-slope approximation of Gent and McWilliams (1990), as pointed out in section 6.3 of Griffies and Greatbatch (2012). In section 3, we compare the exact epineutral tracer gradient of Eqs. (3) and (7) with the approximate small-slope form of Eqs. (12) and (13). We emphasize [from Eq. (13)] that this comparison is equivalent to comparing the projected nonorthogonal version of the epineutral gradient with its exact version .

d. The two versions of the epineutral gradient are not parallel

One might have hoped that the two forms of the epineutral gradient and would be parallel and differ by the factor , which is the ratio of the lateral distances measured in the neutral tangent plane to those measured exactly horizontally in the direction. However, this is not the case. We expose the issues by asking whether the direction of the isoline of constant tracer within the neutral tangent plane as determined by using the projected nonorthogonal coordinate system, namely , is in fact an isoline of . That is, we ask whether is zero. Using Eqs. (10b) and (13) we find that
e14
This result shows that unless either or are zero, the direction of the constant tracer calculated using the projected coordinate approach, or equivalently by using the small-angle approximation, is incorrect. In other words, in order for and to be parallel, either or must be zero, that is, the two-dimensional gradients and must either be parallel or perpendicular. In the general situation when and are neither parallel nor perpendicular, we note that while the two gradients and are both directed in the neutral tangent plane, they are not parallel. This misalignment means that when using to parameterize eddy fluxes, there will be a component of the tracer flux in a direction in which there is in fact no tracer gradient. In addition, even when and are parallel (say because ), these gradients differ in magnitude, but this difference is of little concern.

3. Comparing the two versions of the epineutral gradient

a. Deriving equations relating the two epineutral gradients

To help understand the relationships between the two gradients and , we begin by finding an expression for the exact epineutral gradient in terms of the projected nonorthogonal gradients and . Using Eqs. (3a) and (10b) we find
e15
Using Eqs. (13) and (15) we can relate the following components of to those of . The magnitude of the vertical component of is
e16
the horizontal component of is
e17
and the magnitude of in the direction (where ) is
e18
while the magnitude of in the direction is
e19
Another very useful combination that we can derive from Eqs. (13) and (15) is the difference between and [recall that the factor is the ratio of the lateral distances measured along the neutral tangent plane versus horizontally for displacements in the horizontal direction ]:
e20
The last part of this equation is obtained by first noticing that is (i) perpendicular to and (ii) is a horizontal vector, so it must be in the direction.
Equation (20) is a convenient starting point to derive an expression for the cross product of and (which would be zero if these three-dimensional gradients were parallel):
e21
This result confirms that found in Eq. (14). Namely, in order for the exact and the approximate versions of the two epineutral tracer gradients, and , to be parallel, either or must be zero. That is, the two-dimensional gradients and must either be parallel or perpendicular.

b. Visualizing the difference between the two epineutral gradients

The geometry of the horizontal components of the projected and exact forms of the epineutral gradients, and , is sketched in Fig. 3, using what we have learnt from the above Eqs. (18)–(20). The slope of the neutral tangent plane is taken to be directed due south (having the Southern Ocean in mind). Because both and lie in the neutral tangent plane (i.e., both and are zero), it is sufficient to examine the horizontal components of and in order to understand their differences.

Fig. 3.
Fig. 3.

(a) Sketch of the horizontal components of the projected nonorthogonal epineutral tracer gradient and the exact epineutral tracer gradient . The horizontal components of these gradients are and , respectively. The components of these horizontal fluxes in the direction are equal [see Eq. (19)], while the components in the direction are in the ratio [see Eq. (18)]. The vector points in the direction and is equal to [see Eq. (20)]. Panel (a) has been drawn with . (b) Illustrates the special case when both and point due north (or due south) and these gradients are in the ratio .

Citation: Journal of Physical Oceanography 44, 8; 10.1175/JPO-D-13-0270.1

From Eq. (19), we see that the components of and in the direction are equal, while from Eq. (18) the components of and in the direction are in the ratio , that is, . From Eq. (20), the difference between the three-dimensional gradients and is exactly the horizontal vector , and this vector is shown in Fig. 3a. This vector points in the direction , and its magnitude is . In the special case where (illustrated in Fig. 3b), points due north and , showing that the projected nonorthogonal version of the epineutral tracer gradient is now parallel to the exact epineutral gradient , but that these gradients differ in magnitude by the expected factor . This special case of being zero provides excellent motivation for emphasizing the comparison between and in Eq. (20) and Fig. 3a in the general case when is not zero.

This geometric discussion serves to illustrate that the relevant comparison between the exact and the small-slope approximated versions of the epineutral tracer gradient is between and [or equivalently between and ]. In the diffusion tensor form, the relevant difference between these fluxes is
e22

The fact that the use of the small-slope approximation leads to gradients in the and directions that differ in relative magnitude (compared to the corresponding components of ) by the factor has been pointed out in section 14.1.4.3 of Griffies [2004; see his Eq. (14.26) in which this result is couched in terms of an amplified isopycnal diffusivity by the factor in the direction]. However, the fact that the small-slope approximation involves the flux of tracer in a direction in which there is no actual tracer gradient has, to our knowledge, not been noticed heretofore.

c. How different are the two epineutral gradients?

Since the horizontal component of is , it would seem that the relative error in neglecting the difference between and would scale [using Eq. (20)] as
e23
However, the relative error is actually significantly less than this scaling suggests because the magnitude of in the ocean is usually much less than . Here we develop expressions for and make the connection to neutral helicity.
Since , is also equal to . Using the expression Eq. (A5) for the neutral tangent plane slope, we find
e24
Consider now the case where the tracer is Conservative Temperature (we could equally well have chosen Absolute Salinity for this purpose since their epineutral gradients are parallel), so that Eq. (24) becomes
e25
To within the Boussinesq approximation we may take and to be equal to the corresponding gradients in an isobaric surface, and , so that
e26
where the last three parts of this equation have used the results of McDougall and Jackett (1988, 2007) and section 3.13 of IOC et al. (2010) that relate these various triple scalar products to neutral helicity , defined as , where is the thermobaric parameter (McDougall 1987b) that expresses the nonlinear dependence of specific volume on both Conservative Temperature and pressure; a nonlinear property that does not concern us in this section of the paper.

The triple scalar product of neutral helicity has arisen in the context of

  1. the ill-defined nature of neutral surfaces and the empty nature of ocean hydrographic data in space (McDougall and Jackett 1988, 2007);

  2. the mean vertical downwelling advection achieved by the helical nature of neutral trajectories (Klocker and McDougall 2010);

  3. the close connection between and the spiraling of epineutral contours when the ocean is not motionless (Zika et al. 2010); and now

  4. the difference between the projected nonorthogonal epineutral gradient of , , and its exact epineutral gradient counterpart .

The quantification of the magnitude of neutral helicity in the ocean is far from complete, yet the above studies have shown that the angle between and is quite small in most of the ocean, even in the Southern Ocean where the influence of processes related to neutral helicity seem to be the largest (Klocker and McDougall 2010). These results for neutral helicity were obtained in both a smooth ocean atlas and in an ocean model without eddies, but neutral helicity has not yet been examined in the context of a high-resolution eddy-permitting ocean model. Nevertheless, if we take these results of McDougall and Jackett (2007) and Klocker and McDougall (2010) at face value, then we would conclude that the relative magnitude of the error in the tracer flux due to using the small-slope approximation may be no more than a few percent of the magnitude estimated by the scale analysis of Eq. (23), that is, no more than a few percent of . It must be said that the reason why neutral helicity is as small as it is in the ocean is far from clear; see McDougall and Jackett (2007) and Klocker and McDougall (2010) regarding neutral helicity and the consequent requirement that the ocean is “thin” in space.

Specifically, if we let be the (small) angle between the two-dimensional gradients and , then from Eq. (23) we see that the relative error in neglecting the difference between and is actually
e27

The error made by neglecting this fraction of the epineutral tracer gradient can be compared with the very small errors that are made by taking Conservative Temperature to be 100% conservative. From McDougall (2003) and Graham and McDougall (2013), we know that Conservative Temperature is approximately two orders of magnitude more conservative than is potential temperature [see Figs. 2a and 4 of McDougall (2003) and Fig. 8 of Graham and McDougall (2013)], while from Fig. 11a of Graham and McDougall (2013) and appendix A.14 of IOC et al. (2010) we see that the epineutral gradient of potential temperature is often about 1% different than that of Conservative Temperature . This 1% difference is due to the nonconservative nature of potential temperature. We conclude that the relative error in using the epineutral gradient of Conservative Temperature is the product of these two factors of , namely, . This can be compared with the relative error, Eq. (27), involved in fluxing a tracer in a direction in which there is no gradient. For an epineutral slope of as large as this relative error is , which is smaller than the relative error in using Conservative Temperature by the factor , this factor representing the influence of neutral helicity. Since we expect to be no larger than 0.05 (McDougall and Jackett 2007), we conclude that this effect is absolutely tiny, even when compared with the use of Conservative Temperature, which itself is an improvement by two orders of magnitude on oceanographic practice under the International Equation of State of Seawater—1980 (EOS-80).

d. The component of the small-slope gradient in a direction in which there is no actual gradient

The component of the small-slope approximated gradient in the direction in three-dimensional space in which there is no tracer gradient is . The magnitude of this component can be calculated with the use of Eq. (21) as follows:
e28
where the approximations are (i) so that , and (ii) that is close to being parallel to so that the sine of the angle between these two-dimensional vectors is small.

We conclude that the relative magnitude of the component of in the direction in which there is no tracer gradient is the same as Eq. (27), namely, . In other words, the salient approximation in using the projected nonorthogonal approach to calculating tracer gradients manifests itself in the direction in space in which there is no tracer gradient. This result can be understood from Eq. (20) and Fig. 3a; for the usual situation where is close to being parallel to , the error vector of Eq. (20), , is close to being aligned in the direction , in which there is no epineutral tracer gradient.

e. Incorporating the full epineutral gradient into ocean models

It is possible for a layered ocean model to incorporate the full epineutral tracer flux while retaining the projected nonorthogonal nature of its coordinate system. This correction can be achieved by replacing the existing horizontal component of the projected epineutral flux with the horizontal component of the full epineutral flux [from Eq. (15)]:
e29
The projected nonorthogonal coordinate system of the layered model will treat this gradient as the horizontal component of its projected epineutral gradient whose three-dimensional expression has the additional exactly vertical component whose magnitude is the scalar product of with Eq. (29), namely, , exactly as it should be, from Eq. (15). So this approach provides a straightforward fix for layered ocean models. For Cartesian coordinate models, the solution is to simply use the full diffusion tensor [Eq. (7)] rather than its small-slope approximation [Eq. (13)].

This same suggestion of replacing the horizontal component of the projected epineutral gradient with Eq. (29) applies to theoretical studies that use the projected nonorthogonal coordinate system in order to accurately represent epineutral diffusion.

4. The isotropic diffusivity of small-scale mixing processes

The so-called vertical diffusivity or dianeutral diffusivity used in physical oceanography is a diffusion coefficient that parameterizes mixing from a variety of small-scale mixing processes, with the primary mixing associated with breaking internal gravity waves. The turbulent cascade in breaking internal waves begins at vertical scales of order 1 m and proceeds downscale to the molecular dissipation scale of a few millimeters. The turbulent fluxes achieved by this turbulent cascade do not occur only in the vertical direction or the dianeutral direction, but rather the diffusion occurs almost isotropically (Osborn 1980; Gargett et al. 1984). Hence, the isotropic diffusivity of small-scale turbulent mixing processes causes fluxes per unit area in x, y, z Cartesian coordinates at the rate
e30
This small-scale isotropic turbulent diffusion can be added to the full diffusion tensor of epineutral mixing, Eq. (7) [with ], to obtain
e31
where is the epineutral diffusion coefficient. The small-slope approximation to Eq. (31) (now with ) becomes
e32
The small-slope approximation by itself does not change the contribution of the dianeutral diffusivity in going from Eqs. (31) to (32). An isolated blob of tracer will still diffuse spherically in a stationary ocean with no mesoscale diffusion. In practice one might drop the two occurrences of from the terms in the (1, 1) and (2, 2) positions in the tensor and then the small-scale turbulent diffusion would act one-dimensionally in the vertical direction, the same as in the usual small-slope approximation. But there is no need to get rid of these horizontal diffusion terms from Eqs. (31) or (32); these terms are not part of a small-slope approximation.

The above treatment of dianeutral diffusion in Eqs. (30)–(32) is different from what has been done previously in ocean modeling (e.g., Redi 1982; Hirst et al. 1996; Gent and McWilliams 1990). In the traditional treatment of dianeutral diffusion, mixing is taken to occur exactly in the dianeutral direction with a zero component along the neutral tangent plane. However, this orientation is not in accord with how the ocean works. Instead, a blob of dye is diffused by small-scale turbulent mixing in a spherical manner. In contrast, as illustrated in Fig. 4, the dianeutral diffusion tensor of Redi (1982) diffuses this blob only in the dianeutral direction. Hence, if dianeutral diffusion is allowed to operate in the absence of epineutral diffusion, an initial blob of tracer will be diffused dianeutrally into a thin line in space, as in Fig. 4. This one-dimensional dianeutral diffusion is not how small-scale turbulent mixing works; rather the diffusion should be in three dimensions. Likewise, in the now common small-slope approximation of the full Redi diffusion tensor, the dianeutral diffusion becomes purely vertical diffusion, which is again a purely one-dimensional diffusion. In this case, a small spherical blob of dye, when acted upon by only this small-slope approximated dianeutral diffusion will spread vertically in a thin vertical line, and it will never be any wider in the x and y directions than its initial size in these directions.

Fig. 4.
Fig. 4.

Sketch showing how, in the absence of epineutral diffusion, the so-called dianeutral diffusion of the Redi (1982) diffusion tensor diffuses a small initial blob of tracer (the black dot) only in the dianeutral direction, leading to a line of tracer in the direction normal to a neutral tangent plane. In the small-slope approximation to the Redi diffusion tensor, the so-called dianeutral diffusion diffuses tracer only in the vertical direction, leading to a vertical line of tracer. However, in reality, small-scale mixing processes actually diffuse properties isotropically so that a small initial blob of tracer will spread spherically in the absence of epineutral diffusion.

Citation: Journal of Physical Oceanography 44, 8; 10.1175/JPO-D-13-0270.1

We contend that an ocean model should be able to increase the dianeutral diffusivity and decrease the epineutral diffusivity (even to zero) and achieve three-dimensional spherical diffusion. For this purpose, Eq. (31) is the proper form of the full diffusion tensor and Eq. (32) is its small-slope approximation. The existing Redi (1982) and Gent and McWilliams (1990) versions of these tensors are not appropriate for this purpose.

5. Conclusions

The points raised in this note are of a conceptual nature and heretofore incompletely explored in the literature. We anticipate that their importance for ocean modeling practice, though yet to be tested, may in fact be negligible. That is, the niceties exposed in this paper may have little impact on large-scale simulation integrity, particularly compared to uncertainties associated with other aspects of diffusive closures (e.g., values for the diffusivities). We are nonetheless compelled to raise the ideas here to clarify basic notions regarding geometric aspects of ocean interior mixing parameterized by diffusion. In this way this note complements the warnings of Young (2012) regarding the unconventional (but not incorrect) use of projected nonorthogonal coordinates in atmospheric and oceanic science.

Specifically, the key points made in this note are the following:

  • The epineutral diffusion achieved by the small-slope approximation is not down the correct epineutral tracer gradient . That is, under the small-slope approximation, there is a small gradient of tracer in a direction in which there is no actual epineutral gradient of tracer. This is also an undesirable property of the projected nonorthogonal coordinate system that is used in layered ocean models and in theoretical oceanographic studies.

  • The difference between the correct epineutral tracer gradient and the small-slope approximation to it, , is explained geometrically (see Fig. 3). The relevant difference between the tracer gradients, , is equal to [from Eq. (20)] and points in the direction of the thermal wind (here is the square of the slope of the neutral tangent plane).

  • For (the tracer) Conservative Temperature, the difference between the correct (Redi) epineutral gradient and the small-slope approximation to it is proportional to neutral helicity and to the square of the slope of the neutral tangent plane.

  • While it is uncomfortable to realize that the small-angle approximation to epineutral diffusion (and equivalently, the use of the projected nonorthogonal coordinate system) gives rise to an epineutral property flux in a direction in which there is actually no epineutral tracer gradient, it must be said that the fraction of the epineutral flux in this direction is very small and is negligible for all foreseeable applications. For example, we have shown that for an epineutral slope as large as , the error in using Conservative Temperature (which is not in fact 100% conservative) is greater than the error discussed in this paper of having a lateral tracer flux in a direction in which there is no gradient.

  • Small-scale mixing processes act to diffuse tracers isotropically (i.e., directionally uniformly in space); hence, it is a misnomer to call this process dianeutral diffusion. This realization simplifies the diffusion tensor that is used in ocean models. This is illustrated by a thought experiment in which a tiny blob of tracer is diffused by small-scale mixing processes (see Fig. 4). The blob should diffuse spherically, whereas the full Redi (1982) tensor has it diffusing as a line (a pencil) normal to the neutral tangent plane, and the small-slope approximation has this line being vertical.

Acknowledgments

The authors thank the anonymous reviewers for help in clarifying elements of this note. SMG wishes to thank Peter Gent for discussions on matters related to this work during July 2012.

APPENDIX

The Slope of the Neutral Tangent Plane

Here we first prove that even though the two-dimensional gradients and point in different directions, the projected gradients of Absolute Salinity and Conservative Temperature, and , are parallel and the neutrality condition applies to the projected gradients in the neutral tangent plane.

The neutral tangent plane, as discussed in sections 1 and 2, has the property that the gradients of Absolute Salinity and Conservative Temperature in this plane cause equal and opposite contributions to the gradient of locally referenced potential density, that is,
ea1
This balancing of the gradients of and through and is called the “neutrality condition” and Eqs. (8) and (A1) show that this neutrality condition applies to the three-dimensional gradients and . Using Eq. (15) we can express and in terms of the corresponding gradients and in the projected nonorthogonal coordinate system, and we form the linear combination that we know to be zero:
ea2
where we have used the shorthand notation . This vector Eq. (A2) can only be satisfied if both and , which reduces to the requirement that . Hence, we have found that the neutrality condition Eq. (A1) implies that
eA3ab
a result previously found by Griffies and Greatbatch (2012). That is, even though the two-dimensional gradients and are neither equal nor parallel (and the same comment applies to the corresponding gradients of Absolute Salinity), the projected two-dimensional epineutral gradients and are parallel, and they obey the neutrality condition , just as their three-dimensional cousins and also obey the neutrality condition .
Now we find the expression for the slope of the neutral tangent plane , with our discussion complementing that found in section 6.5 of Griffies (2004). The projected nonorthogonal gradient of a scalar in a neutral tangent plane is related to that at constant height by Eq. (10a), namely, , where is the exactly vertical gradient at constant longitude and latitude and all of and are exactly two-dimensional gradients. Applying this equation to Absolute Salinity and Conservative Temperature, and taking the linear combination with the thermal expansion coefficient and the saline contraction coefficient, gives
ea4
and from Eq. (A3) we know that the left-hand side of this equation is zero so the slope of the neutral tangent plane is (where is the buoyancy frequency)
ea5
This equation serves to define the x and y components and of the vector slope . For later use we use the shorthand notation for the inner product .
It can also be shown that the two-dimensional projected gradient of pressure in the neutral tangent plane can be expressed as
ea6

REFERENCES

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    • Search Google Scholar
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    • Search Google Scholar
    • Export Citation

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    • Search Google Scholar
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  • McDougall, T. J., and J. A. Church, 1986: Pitfalls with the numerical representation of isopycnal and diapycnal mixing. J. Phys. Oceanogr., 16, 196–199, doi:10.1175/1520-0485(1986)016<0196:PWTNRO>2.0.CO;2.

    • Search Google Scholar
    • Export Citation

  • McDougall, T. J., and D. R. Jackett, 1988: On the helical nature of neutral trajectories in the ocean. Prog. Oceanogr., 20, 153–183, doi:10.1016/0079-6611(88)90001-8.

    • Search Google Scholar
    • Export Citation

  • McDougall, T. J., and P. C. McIntosh, 2001: The temporal-residual-mean velocity. Part II: Isopycnal interpretation and the tracer and momentum equations. J. Phys. Oceanogr., 31, 1222–1246, doi:10.1175/1520-0485(2001)031<1222:TTRMVP>2.0.CO;2.

    • Search Google Scholar
    • Export Citation

  • McDougall, T. J., and D. R. Jackett, 2005: The material derivative of neutral density. J. Mar. Res., 63, 159–185, doi:10.1357/0022240053693734.

    • Search Google Scholar
    • Export Citation

  • McDougall, T. J., and D. R. Jackett, 2007: The thinness of the ocean in space and the implications for mean diapycnal advection. J. Phys. Oceanogr., 37, 1714–1732, doi:10.1175/JPO3114.1.

    • Search Google Scholar
    • Export Citation

  • McDougall, T. J., and P. M. Barker, 2011: Getting started with TEOS-10 and the Gibbs Seawater (GSW) Oceanographic Toolbox. SCOR/IAPSO WG127 Rep., 31 pp. [Available online at http://www.teos-10.org/pubs/Getting_Started.pdf.]

  • Nycander, J., 2011: Energy conversion, mixing energy, and neutral surfaces with a nonlinear equation of state. J. Phys. Oceanogr., 41, 28–41, doi:10.1175/2010JPO4250.1.

    • Search Google Scholar
    • Export Citation

  • Osborn, T. R., 1980: Estimates of the local rate of vertical diffusion from dissipation measurements. J. Phys. Oceanogr., 10, 83–89, doi:10.1175/1520-0485(1980)010<0083:EOTLRO>2.0.CO;2.

    • Search Google Scholar
    • Export Citation

  • Redi, M. H., 1982: Oceanic isopycnal mixing by coordinate rotation. J. Phys. Oceanogr., 12, 1154–1158, doi:10.1175/1520-0485(1982)012<1154:OIMBCR>2.0.CO;2.

    • Search Google Scholar
    • Export Citation

  • Solomon, H., 1971: On the representation of isentropic mixing in ocean models. J. Phys. Oceanogr., 1, 233–234, doi:10.1175/1520-0485(1971)001<0233:OTROIM>2.0.CO;2.

    • Search Google Scholar
    • Export Citation

  • Starr, V. P., 1945: A quasi-Lagrangian system of hydrodynamical equations. J. Meteor., 2, 227–237, doi:10.1175/1520-0469(1945)002<0227:AQLSOH>2.0.CO;2.

    • Search Google Scholar
    • Export Citation

  • Veronis, G., 1975: The role of models in tracer studies. Numerical Models of Ocean Circulation, National Academy of Science, 133–146.

  • Young, W. R., 2012: An exact thickness-weighted average formulation of the Boussinesq equations. J. Phys. Oceanogr., 42, 692–707, doi:10.1175/JPO-D-11-0102.1.

    • Search Google Scholar
    • Export Citation

  • Zika, J. D., T. J. McDougall, and B. M. Sloyan, 2010: A tracer-contour inverse method for estimating ocean circulation and mixing. J. Phys. Oceanogr., 40, 26–47, doi:10.1175/2009JPO4208.1.

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    • Export Citation
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Journal of Physical Oceanography

  • Bleck, R., 1978a: Finite difference equations in generalized vertical coordinates. Part I: Total energy conservation. Contrib. Atmos. Phys., 51, 360–372.

    • Search Google Scholar
    • Export Citation

  • Bleck, R., 1978b: Finite difference equations in generalized vertical coordinates. Part II: Potential vorticity conservation. Contrib. Atmos. Phys., 52, 95–105.

    • Search Google Scholar
    • Export Citation

  • Gargett, A. E., T. R. Osborn, and P. W. Nasymyth, 1984: Local isotropy and the decay of turbulence in a stratified fluid. J. Fluid Mech., 144, 231–280, doi:10.1017/S0022112084001592.

    • Search Google Scholar
    • Export Citation

  • Gent, P. R., and J. C. McWilliams, 1990: Isopycnal mixing in ocean circulation models. J. Phys. Oceanogr., 20, 150–155, doi:10.1175/1520-0485(1990)020<0150:IMIOCM>2.0.CO;2.

    • Search Google Scholar
    • Export Citation

  • Gent, P. R., J. Willebrand, T. J. McDougall, and J. C. McWilliams, 1995: Parameterizing eddy-induced tracer transports in ocean circulation models. J. Phys. Oceanogr., 25, 463–474, doi:10.1175/1520-0485(1995)025<0463:PEITTI>2.0.CO;2.

    • Search Google Scholar
    • Export Citation

  • Graham, F. S., and T. J. McDougall, 2013: Quantifying the nonconservative production of Conservative Temperature, potential temperature, and entropy. J. Phys. Oceanogr., 43, 838–862, doi:10.1175/JPO-D-11-0188.1.

    • Search Google Scholar
    • Export Citation

  • Griffies, S. M., 1998: The Gent–McWilliams skew flux. J. Phys. Oceanogr., 28, 831–841, doi:10.1175/1520-0485(1998)028<0831:TGMSF>2.0.CO;2.

    • Search Google Scholar
    • Export Citation

  • Griffies, S. M., 2004: Fundamentals of Ocean Climate Models. Princeton University Press, 518 pp.

  • Griffies, S. M., and R. J. Greatbatch, 2012: Physical processes that impact the evolution of global mean sea level in ocean climate models. Ocean Modell., 51, 37–72, doi:10.1016/j.ocemod.2012.04.003.

    • Search Google Scholar
    • Export Citation

  • Hirst, A. C., D. R. Jackett, and T. J. McDougall, 1996: The meridional overturning cells of a world ocean model in neutral density coordinates. J. Phys. Oceanogr., 26, 775–791, doi:10.1175/1520-0485(1996)026<0775:TMOCOA>2.0.CO;2.

    • Search Google Scholar
    • Export Citation

  • IOC, SCOR and IAPSO, 2010: The International Thermodynamic Equation of Seawater—2010: Calculation and use of thermodynamic properties. UNESCO Intergovernmental Oceanographic Commission, Manuals and Guides 56, 196 pp. [Available online at http://www.teos-10.org/pubs/TEOS-10_Manual.pdf.]

  • Iselin, C. O’D., 1939: The influence of vertical and lateral turbulence on the characteristics of the waters at mid-depths. Eos, Trans. Amer. Geophys. Union, 20, 414–417.

    • Search Google Scholar
    • Export Citation

  • Jackett, D. R., and T. J. McDougall, 1997: A neutral density variable for the world’s oceans. J. Phys. Oceanogr., 27, 237–263, doi:10.1175/1520-0485(1997)027<0237:ANDVFT>2.0.CO;2.

    • Search Google Scholar
    • Export Citation

  • Klocker, A., and T. J. McDougall, 2010: Quantifying the consequences of the ill-defined nature of neutral surfaces. J. Phys. Oceanogr., 40, 1866–1880, doi:10.1175/2009JPO4212.1.

    • Search Google Scholar
    • Export Citation

  • Ledwell, J., L. St. Laurent, J. B. Girton, and J. M. Toole, 2011: Diapycnal mixing in the Antarctic Circumpolar Current. J. Phys. Oceanogr., 41, 241–246, doi:10.1175/2010JPO4557.1.

    • Search Google Scholar
    • Export Citation

  • MacKinnon, J., L. St. Laurent, and A. C. Naveira Garabato, 2013: Diapycnal mixing processes in the ocean interior. Ocean Circulation and Climate, 2nd ed. G. Siedler et al., Eds., Academic Press, 159–183.

  • McDougall, T. J., 1987a: Neutral surfaces. J. Phys. Oceanogr., 17, 1950–1964, doi:10.1175/1520-0485(1987)017<1950:NS>2.0.CO;2.

  • McDougall, T. J., 1987b: Thermobaricity, cabbeling, and water-mass conversion. J. Geophys. Res., 92, 5448–5464, doi:10.1029/JC092iC05p05448.

    • Search Google Scholar
    • Export Citation

  • McDougall, T. J., 1988: Neutral-surface potential vorticity. Prog. Oceanogr., 20, 185–221, doi:10.1016/0079-6611(88)90002-X.

  • McDougall, T. J., 1995: The influence of ocean mixing on the absolute velocity vector. J. Phys. Oceanogr., 25, 705–725, doi:10.1175/1520-0485(1995)025<0705:TIOOMO>2.0.CO;2.

    • Search Google Scholar
    • Export Citation

  • McDougall, T. J., 2003: Potential enthalpy: A conservative oceanic variable for evaluating heat content and heat fluxes. J. Phys. Oceanogr., 33, 945–963, doi:10.1175/1520-0485(2003)033<0945:PEACOV>2.0.CO;2.

    • Search Google Scholar
    • Export Citation

  • McDougall, T. J., and J. A. Church, 1986: Pitfalls with the numerical representation of isopycnal and diapycnal mixing. J. Phys. Oceanogr., 16, 196–199, doi:10.1175/1520-0485(1986)016<0196:PWTNRO>2.0.CO;2.

    • Search Google Scholar
    • Export Citation

  • McDougall, T. J., and D. R. Jackett, 1988: On the helical nature of neutral trajectories in the ocean. Prog. Oceanogr., 20, 153–183, doi:10.1016/0079-6611(88)90001-8.

    • Search Google Scholar
    • Export Citation

  • McDougall, T. J., and P. C. McIntosh, 2001: The temporal-residual-mean velocity. Part II: Isopycnal interpretation and the tracer and momentum equations. J. Phys. Oceanogr., 31, 1222–1246, doi:10.1175/1520-0485(2001)031<1222:TTRMVP>2.0.CO;2.

    • Search Google Scholar
    • Export Citation

  • McDougall, T. J., and D. R. Jackett, 2005: The material derivative of neutral density. J. Mar. Res., 63, 159–185, doi:10.1357/0022240053693734.

    • Search Google Scholar
    • Export Citation

  • McDougall, T. J., and D. R. Jackett, 2007: The thinness of the ocean in space and the implications for mean diapycnal advection. J. Phys. Oceanogr., 37, 1714–1732, doi:10.1175/JPO3114.1.

    • Search Google Scholar
    • Export Citation

  • McDougall, T. J., and P. M. Barker, 2011: Getting started with TEOS-10 and the Gibbs Seawater (GSW) Oceanographic Toolbox. SCOR/IAPSO WG127 Rep., 31 pp. [Available online at http://www.teos-10.org/pubs/Getting_Started.pdf.]

  • Nycander, J., 2011: Energy conversion, mixing energy, and neutral surfaces with a nonlinear equation of state. J. Phys. Oceanogr., 41, 28–41, doi:10.1175/2010JPO4250.1.

    • Search Google Scholar
    • Export Citation

  • Osborn, T. R., 1980: Estimates of the local rate of vertical diffusion from dissipation measurements. J. Phys. Oceanogr., 10, 83–89, doi:10.1175/1520-0485(1980)010<0083:EOTLRO>2.0.CO;2.

    • Search Google Scholar
    • Export Citation

  • Redi, M. H., 1982: Oceanic isopycnal mixing by coordinate rotation. J. Phys. Oceanogr., 12, 1154–1158, doi:10.1175/1520-0485(1982)012<1154:OIMBCR>2.0.CO;2.

    • Search Google Scholar
    • Export Citation

  • Solomon, H., 1971: On the representation of isentropic mixing in ocean models. J. Phys. Oceanogr., 1, 233–234, doi:10.1175/1520-0485(1971)001<0233:OTROIM>2.0.CO;2.

    • Search Google Scholar
    • Export Citation

  • Starr, V. P., 1945: A quasi-Lagrangian system of hydrodynamical equations. J. Meteor., 2, 227–237, doi:10.1175/1520-0469(1945)002<0227:AQLSOH>2.0.CO;2.

    • Search Google Scholar
    • Export Citation

  • Veronis, G., 1975: The role of models in tracer studies. Numerical Models of Ocean Circulation, National Academy of Science, 133–146.

  • Young, W. R., 2012: An exact thickness-weighted average formulation of the Boussinesq equations. J. Phys. Oceanogr., 42, 692–707, doi:10.1175/JPO-D-11-0102.1.

    • Search Google Scholar
    • Export Citation

  • Zika, J. D., T. J. McDougall, and B. M. Sloyan, 2010: A tracer-contour inverse method for estimating ocean circulation and mixing. J. Phys. Oceanogr., 40, 26–47, doi:10.1175/2009JPO4208.1.

    • Search Google Scholar
    • Export Citation
  • <sc>Fig</sc>. 1.
    View in gallery
    Fig. 1.

    Sketch of a central seawater parcel being moved adiabatically and without change in its salinity to either the right or the left of its original position in a direction that is not neutral. When the parcel is then released it feels a vertical buoyant force and begins to move vertically (upward on the left and downward on the right) toward its original “isopycnal.”

  • <sc>Fig</sc>. 2.
    View in gallery
    Fig. 2.

    (a) A three-dimensional perspective of the projected nonorthogonal coordinate system that is commonly used in layered ocean models. (b) The projected nonorthogonal gradient of a property in the neutral tangent plane, in the limit as the distances tend to zero, is equal to the tracer difference between points a and b in the above figure, divided by the exactly horizontal distance or .

  • <sc>Fig</sc>. 3.
    View in gallery
    Fig. 3.

    (a) Sketch of the horizontal components of the projected nonorthogonal epineutral tracer gradient and the exact epineutral tracer gradient . The horizontal components of these gradients are and , respectively. The components of these horizontal fluxes in the direction are equal [see Eq. (19)], while the components in the direction are in the ratio [see Eq. (18)]. The vector points in the direction and is equal to [see Eq. (20)]. Panel (a) has been drawn with . (b) Illustrates the special case when both and point due north (or due south) and these gradients are in the ratio .

  • <sc>Fig</sc>. 4.
    View in gallery
    Fig. 4.

    Sketch showing how, in the absence of epineutral diffusion, the so-called dianeutral diffusion of the Redi (1982) diffusion tensor diffuses a small initial blob of tracer (the black dot) only in the dianeutral direction, leading to a line of tracer in the direction normal to a neutral tangent plane. In the small-slope approximation to the Redi diffusion tensor, the so-called dianeutral diffusion diffuses tracer only in the vertical direction, leading to a vertical line of tracer. However, in reality, small-scale mixing processes actually diffuse properties isotropically so that a small initial blob of tracer will spread spherically in the absence of epineutral diffusion.

Fig. 1.

Sketch of a central seawater parcel being moved adiabatically and without change in its salinity to either the right or the left of its original position in a direction that is not neutral. When the parcel is then released it feels a vertical buoyant force and begins to move vertically (upward on the left and downward on the right) toward its original “isopycnal.”
Fig. 2.

(a) A three-dimensional perspective of the projected nonorthogonal coordinate system that is commonly used in layered ocean models. (b) The projected nonorthogonal gradient of a property in the neutral tangent plane, in the limit as the distances tend to zero, is equal to the tracer difference between points a and b in the above figure, divided by the exactly horizontal distance or .
Fig. 3.

(a) Sketch of the horizontal components of the projected nonorthogonal epineutral tracer gradient and the exact epineutral tracer gradient . The horizontal components of these gradients are and , respectively. The components of these horizontal fluxes in the direction are equal [see Eq. (19)], while the components in the direction are in the ratio [see Eq. (18)]. The vector points in the direction and is equal to [see Eq. (20)]. Panel (a) has been drawn with . (b) Illustrates the special case when both and point due north (or due south) and these gradients are in the ratio .
Fig. 4.

Sketch showing how, in the absence of epineutral diffusion, the so-called dianeutral diffusion of the Redi (1982) diffusion tensor diffuses a small initial blob of tracer (the black dot) only in the dianeutral direction, leading to a line of tracer in the direction normal to a neutral tangent plane. In the small-slope approximation to the Redi diffusion tensor, the so-called dianeutral diffusion diffuses tracer only in the vertical direction, leading to a vertical line of tracer. However, in reality, small-scale mixing processes actually diffuse properties isotropically so that a small initial blob of tracer will spread spherically in the absence of epineutral diffusion.
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