Eq 3-1:

(conductive heat flux, *q*) = - (thermal conductivity, *k*) x (temperature gradient, *dT/dz*)

The negative sign reflects that heat is
transported in the direction of lower temperature. For a historical
discussion of Fourier's Law and its influence on the description of
other physical systems, see **(44)**. In dimensional
terms:

recent literature (SI) | older literature | |
---|---|---|

flux | W m | cal cms |

thermal conductivity | W mK | cal cms°C |

gradient | K m | °C m |

The pratical unit from older literature is the HFU (heat flow unit). 1 HFU (heat flow unit) = 41.9 mW m = 1 µcal cms The global average conductive heat flow is about 2 HFU.

The thermal conductivity (or equivalently the specific thermal conductivity, (also called the thermal diffusivity) which includes the density and heat capacity of a material) varies considerably, over about four orders of magnitude in various solid materials:

k [W m°C] | =k/(C) [10 ^{-8} m^{2}s^{-1}] | |
---|---|---|

thermal conductivity | specific thermal conductivity or thermal diffusivity | |

air | .025 | 1938 |

water | 0.6 | 14 |

ice | 2.1 | 104 |

aluminum | 237 | 9975 |

copper | 390 | 11161 |

fiber glass | 0.04 | 44 |

surficial ocean sediment | 0.9 | 20 |

basaltic ocean crust | 2.9 | 77 |

upper mantle | 3.1 | 80 |

quartz | 3 | 141 |

saturated sand | 2.7 | 102 |

Using observations of the conductive heat flow we can estimate the temperature increase with depth into the oceanic lithosphere. At 100 My crustal age the heat flow is ~1.1 HFU = 46 mW m, so:

Eq 3-2:*q=-k(dT/dz)*

Eq 3-3:46 x 10 W m = 3.1 W m°C *(dT/dz)*

Eq 3-4:*(dT/dz)* = 15 x 10 °C m

Thus, at 100 km depth, *T* ~1500 °C. (This is an overestimate, as even
after 100 My, the lithosphere is continuing to cool and so the profile
is not strictly linear.)

This approach is reversed in order to make heat flow measurements: the temperature gradient is measured with a probe consisting of a number of thermistors which is inserted into sediment; the thermal conductivity is measured (by measuring the response after a thermal perturbation) or estimated from other data; the product yields the heat flow.

To make use of Fourier's Law to understand the rate of change of the
temperature of a material, we need to establish a thermal balance for
a small volume element. Again in
one-dimension, we balance the fluxes of heat through the two surfaces and the
internal production of heat
per unit volume, *H*, to determine the temperature change in an interval of time
dt:

Figure 3-1

Conservation of heat requires that:

Eq 3-5:

Rearranging equation 3-5 and taking to limits:

Eq 3-6:

Well away from the ridge axis where the time rate of change becomes small, it
becomes appropriate to consider the
steady-state characteristics of equation 3-6. The goal
is to examine whether the internal production
of heat, *H* has a significant influence on the temperature distribution.

At steady state:

Eq 3-7:

Equation 3-7 is solved by integrating twice to find:

Eq 3-8:

The two constants of integration are found by specifying the temperature at the sediment surface:

Eq 3-9:

and the surface heat flow

Eq 3-10:

Combining equations 3-8, 3-9 and 3-10, we find:

Eq 3-11:

If *H*=0, the temperature increases linearly with depth. For positive
internal production of heat, there is a negative deviation from this
linear trend which grows as z squared. (The deviation is negative
because of the boundary conditions we have chosen; as *H* increases more
of the surface heat flux is derived from internal production of heat and
less from the flux of heat from below).
In the case of the lithosphere, is this deviation important? Because
the elements U, Th and K are partitioned into the liquid as the mantle
melts, considering the radiogenic production of heat
in the basaltic crustal layer represents a worst case (i.e., the mantle
must have a lower content of U, Th, and K). This production is
7.8x10 W m. At 6 km depth, for surface heat flow of 45
mW m and with this value of *H* the temperature deviation is 0.97°C from the value
without internal production, 93°C.
We conclude that we can generally neglect internal production of heat.

Thus for application to conductive heat loss
from the lithosphere we neglect *H* and equation 3-6
becomes:

Eq 3-12:

Eq 3-13:

Imagine that at the axis of spreading that material is held at the temperature of melting, in other words an appropriate initial condition is that at all depths:

Eq 3-14:

(While we will not develop it explicitly, 'temperature' here means 'potential temperature'). For now also imagine that the Earth is infinitely thick, i.e., the geometry is that of an infinite half space and so at great depth:

Eq 3-15:

Finally the upper surface will be will be held at the temperature of the deep ocean:

Eq 3-16:

Omitting the details of solution, we find that:

Eq 3-17:

**(Note the k inside the root operator should be kappa.)** We can use equation 3-17 to calculate the temperature field of an evolving
lithosphere.

The error function is defined by the integral:

and exhibits the behavior:

To illustrate the temperature field, we adopt T=1400 °C, the approximate temperature of the liquidus. For depth we will work in units of km, and for time in units of My so that the appropriate thermal diffusivity is 25 (km)/(My). We will build a grid for 0 to 100 km depth and 0 to 100 My of crustal age in steps of 5 km and 5 My respectively. The solution at t=0 has to be taken directly from equation 3-2 because of the singularity at t=0 in equation 3-5. The calculations are done with the MATLAB script, plate.m:

kappa=25; Tmelt=1400; z=0:5:100; t=0:5:100; for I=1:1:21 T(I,1)=Tmelt; for J=2:1:21 T(I,J)=Tmelt*erf(z(I)/(2*sqrt(kappa*t(J)))); end end cs=contour(t,-z,T); clabel(cs)(New to Matlab? Here is a link to an annotated version of the script.) This script displays a contour plot of the temperature field:

If we think of the base of the lithosphere as corresponding to a particular isotherm, say the 1200 °C isotherm which is approximately the temperature of the solidus, we can trace the growing thickness of the lithosphere as the crust ages. However, because of the semi-infinite geometry adopted, it grows to an unreasonable large thickness, as compared to the constraints provided by seismic studies of the oceanic lithosphere. The temperature gradient (and the conductive heat flux) decreases, most rapidly at first and then more slowly.

While there are some problems evident in the model: specifically

- the
temperature of the seafloor at the ridge axis is not T
_{m}and - the thickness of the lithosphere does not grow without limit

By taking the derivative of equation 3-17 to obtain an expression for
*dT/dz* evaluated at *z*=0 and substituting into equation 3-2
we find that (details):

Eq 3-18:

The heat flow predicted from the boundary layer model decreases with an inverse proportionality to the square root of age.

For now (we will later refine these parameters) we can estimate the relationship using a value of T of 1400 °C and the thermal conductivity data given in Table I:

Eq 3-19:

480 mW/m is ~12 HFU.

Conductive heat flow versus age, from **(43)**

HS is our infinite half space model, PSM is the original Parsons Sclater
model **(8)**, GDH1 is the more recent Stein and Stein (**(43)**) fit to the global data
set. (HS and PSM can only be distinguished on very old lithosphere.)

Oceanography 540 Pages Pages Maintained by Russ McDuff (mcduff@ocean.washington.edu) Copyright (©) 1994-2002 Russell E. McDuff and G. Ross Heath; Copyright Notice Content Last Modified 10/14/2002 | Page Last Built 10/14/2002 |