. This falls to a minimum
of about 0.04 at 350 ° C and
then rises slightly for higher temperatures (Nesbitt, 1993). Up to temperatures
of about 350 ° C the resistivity of seawater,
sw ,
at a temperature T, to a good approximation is given by (e.g. Becker 1985)
There is an ambiguity in the porosity required to explain a given measurement
of resistivity, with a high porosity and ambient seafloor water temperature
giving the same resistivity as a much lower porosity but higher seawater
temperature (Evans etal, 1994). Although conduction in seawater filled cracks
is the predominant mechanism, the presence of alteration products precipitated
in cracks also affects the conductivity (Drury & Hyndman, 1979; Pezard,
1990). Such alteration products in general have a greater resistivity than
seawater, therefore as cracks are in-filled there is an increase in the
resistivity of the crust (Becker, 1985; Pezard, 1990).
The shallow resistivity structure of the AVR is fairly well constrained
around Quail by the short range data collected during the first source tow.
The upper 700 m to 1 km of the structure are constrained primarily
by the 11 Hz data, which are only slightly affected by the 2-dimensionality
of the ridge. Examination of the simple 1-dimensional model is therefore
valid providing the study is limited to the upper part of the crust. The
steep resistivity gradient in the upper 500 m of the structure is the
best constrained feature of the shallow structure, and is consistent with
the steep seismic velocity gradient in the upper few hundred metres of the
crust (Navin etal , this issue) equated with seismic layer 2A.
The upper HS bound can be used to estimate the porosity, taking the resistivity
of the solid phase to be 10,000 , much larger than that of seawater,
so that conduction within the solid is negligible. This parameterisation
gives a good estimate of the conducting volume, but may underestimate the
total porosity if the medium contains isolated pores. The resistivity is
approximately 1 close to the seafloor. Assuming ambient ocean
floor temperatures (around 2°C) the porosity inferred
from this measurement is approximately 50% in the upper 100 m of the
crust. Values of around 40% are inferred from seismic P-wave velocities
(Navin etal, this issue). Purdy (1987) obtained slightly lower seafloor
porosities of around 30% using seismic data from the Mid-Atlantic Ridge
at 23°N. Similar values were obtained by Evans etal
(1994) using CSEM data from the crest of the East Pacific Rise at 13°N.
Submersible surveys reveal significant fracturing and fissuring of the neovolcanic
zone, with fissure widths on the scale of centimetres to metres (Ballard
& Van Andel, 1977). Such fissuring could account for the high porosities
observed in the upper crust.
The results from the Reykjanes Ridge can be compared directly to the CSEM
results of Evans (1994) from 13° N on the East Pacific
Rise. At 1 km below the seafloor the resistivity at the Reykjanes Ridge
is 15, five times less than the resistivity at equivalent depths on the East
Pacific Rise at 13°N. Temperature/porosity trade off
curves are shown in Figure. 9 for these
two bulk resistivities. If the porosity were 0.5% at both ridges, a temperature
difference of approximately 200 could explain the observed difference in
resistivity. Such a temperature difference would be consistent with the
lack of evidence for a large crustal melt body at 13° N
on the East Pacific Rise, which is inferred to be in a state of relative
magmatic quiescence (Evans etal, 1994), compared to the Reykjanes Ridge,
where the RAMESSES experiment has detected a large crustal magma chamber.
It is probable that the difference results from a combination of temperature
and porosity effects. The seismic results (Navin etal, this issue) show
a region of depressed P-wave velocity in the upper crust on the AVR axis,
which is not observed at the East Pacific Rise and could indicate increased
fracture porosity at the Reykjanes Ridge. Results from Deep Sea Drilling
Project hole 409, drilled in crust estimated to be 2.2-3.2 Ma old on the
western flank of the Reykjanes Ridge, indicate that the layer 2 basalts
have a high vesicularity, averaging 27%, caused by the low confining pressure
during eruption at this shallow ridge (Duffield, 1979). Since a connected
porosity of at most 3% in layer 2 is inferred from the CSEM measurements
(Figure. 9), such vesicles must be predominantly
isolated or in-filled so that their contribution to conduction in the basalt
is small.