Filloux (1973, 1987) and Constable et al. (1998) have described instrumentation for ocean-bottom electromagnetic studies. A sche- matic representation of an ocean-bottom MT instrument is shown in Figure 9.1. The instrument sinks to the ocean floor under the weight of a 60 kg concrete anchor that is fastened below its centre.
Electric dipole cables are encased in 5 m lengths of 5 cm diameter polypropylene pipes, which terminate with Ag–AgCl electrodes.
Different coded sequences of timed pings allow commands to be transmitted to the instrument whilst on the ocean floor. One such command code allows the user to instruct the instrument to release the anchor. On receiving the relevant command code, a stainless- steel burn wire that connects the anchor to the system is electrolysed away by means of a voltage supplied by internal batteries, and the instrument rises to the surface aided by glass flotation spheres.
172 The special link to other geosciences
The magnetic variational field at the ocean bottom is attenuated by the conductance of the ocean as described in Section 10.4 (cf. discussion of Equation (3.2) in Section 3.1.1, also). Owing to the skin effect, the attenuation is stronger for shorter periods: the ocean behaves as a low-pass filter. Hence passive ocean-bottom experi- ments have generally been band-limited to periods longer than 1 minute, and crustal structures have rarely been resolved. A recent attempt to resolve partial melt in the crust and shallow mantle below the East Pacific Rise is reported by Key and Constable (2002).
It has been suggested that differences in the resistivity of oceanic and continental mantle might be expected owing to differences in chemical composition (e.g., Jordan, 1978). Conductivity–depth models based on recent ocean-bottom measurements in stable oceanic regions of the NE Pacific Ocean (Lizzaraldeet al., 1995) and Tasman Sea (Heinson and Lilley, 1993) both indicate resistivities in the 5–100m range in the depth interval 100–400 km, whereas models based on long-period electromagnetic data from the Canadian Shield indicate resistivities in the 15–2000m range (e.g., Schultz et al., 1993). However, comparisons between the electrical conductivities of oceanic and continental mantle are inconclusive owing to the small number of regions for which data are available, and to the expected heterogeneity of the mantle.
Comparison of MT data acquired on Mediterranean islands, which are located in the vicinity of oceanic or transitional mantle, with data from stable continental Europe revealed no significant differences in upper-mantle conductivities in the 100–400 km depth range (Simpson, 2002b).
AgCl electrodes
+Exdipole
–Eydipole +Eydipole Bxcoil
Glass flotation spheres (4)
Recording compass
Acoustic release/navigation unit Datalogger
Pressure case By coil
–Ex dipole
AgCl electrodes anchor
Figure 9.1Schematic drawing of ocean-bottom MT instrument. The electric dipole arms are 5 m lengths of 5 cm diameter polypropylene pipes.
Dipole cables run along the insides of the tubing. (Redrawn from Constableet al., 1998.) 9.4 The oceanic mantle: seafloor and island studies 173
Acquisition of ocean-bottom data is logistically difficult and expensive. On the other hand, as far as the inductive attenuation of electromagnetic fields is concerned, at periods long enough to penetrate through the seawater layer, measurements on small islands should be approximately equivalent to measurements made on the ocean bottom, since electromagnetic induction is governed by a diffusion equation, and what results is a volume sounding, not a point sounding. (Notice the similarity between the complex attenuation factor (Equation (10.13)) at the ocean bottom and the complex reduction of theSchmucker–Weidelt transfer func- tion (Equation (2.40)) assuming that an hypothetical instrument swims on the surface of an ocean with conductance .) However, island-based electromagnetic studies are made more complicated by distortion of electromagnetic fields at the interface between the relatively resistive island and highly conductive seawater.
Bathymetric effects can be inductive or galvanic, and distor- tion of magnetic as well as electric fields can occur. For long- period electromagnetic measurements, the electromagnetic skin depthis large compared with the seawater depth, and the seawater layer can, thus, be approximated as a thin sheet of equivalent conductance (Section 6.2). The magnetic distortion can then be formulated by considering the magnetic perturbation tensor, W (Schmucker, 1970; Siemon, 1997), which we describe in more detail in Section 10.1, and which links the magnetic field at a local site to the magnetic field at a reference site. If we identify the reference site with a normal site located outside of the oceanic region that is specified in the thin-sheet model, then Equation (10.1) also describes the distortion at the island site located within the synthe- sised ocean. The corrections to the horizontal magnetic field are thus given by:
Bc¼ W
2þI h i1
Bo , (9:2)
whereBoandBcare the observed and corrected horizontal magnetic fields,W2is the 22 perturbation tensor consisting of the horizontal components in Equation (10.1) andIis the identity matrix.
A similar treatment of the electric fields allows the undistorted impedance tensorZ
cto be calculated from the observed impedance tensorZ
o:
Zc¼ W
2þI h i
ZoW
2Z
(9:3)
withZ being the impedance due to the thin sheet at the location of the island site and W2 the modelled horizontal perturbation
174 The special link to other geosciences
tensor. Since the correction derived is dependent not only on the thin sheet, but also on the underlying resistivity structure, iteration may be necessary if a significant difference is revealed between the resistivities assumed for the mantle during initial correction of the impedance tensor and the conductivity–depth profile modelled subsequently from the corrected impedance tensor. This approach to correct island MT data for bathymetric effects was applied to data acquired on the Mediterranean islands of Mallorca and Montecristo (Simpson, 2002b). A similar approach was used by Nolasco et al.(1998) for removing the bathymetry effects from ocean-bottom data from the Tahiti hotspot. The corrected ocean- bottom data across the Tahiti hotspot are compatible with a modelled mantle plume of limited extent (less than 150 km radius) and slightly higher conductivity than the surrounding mantle (Figure 9.2). Beneath the most active zone, the asthenosphere is laterally heterogeneous, and is more resistive than elsewhere, possibly owing to depletion of volatiles. A possible depression of the olivine–
spinel transition (manifest as a significant decrease in resistivity) from 400 to 450 km is also imaged. The data provide no evidence for strong thermal influence over an extensive area, and therefore do not sup- port a superswell model (unless the electromagnetic signature of the superswell is only present in the lower or mid mantle, in which case electromagnetic data with periods longer than 1 day would be neces- sary to find it). Nor is there evidence for lithospheric thinning owing to plume–lithosphere interaction.
If magnetometers operate simultaneously on the ocean floor and on an island, then a vertical gradient technique (Section 10.4)
Lateral distance (km) –200 –100 0 100 200 700
600 500 400 300 200 100 0
5000 1000 500 100 50 10 5 1
Depth (km)
S10 S11 S15S16S13 S06
ρ (Ω m)
Figure 9.2Best-fit 2-D model obtained from ocean-bottom MT data over the Tahiti hotspot, French Polynesia, following correction for bathymetric and island effects using 3-D thin-sheet modelling. The
electromagnetic strike direction is N1308E, which is in approximate agreement with the spreading direction (N11558E) of the Pacific plate. Station S06 is a reference site located350 km
northwest of the leading edge of the swell. (Redrawn from Nolascoet al., 1998.) 9.4 The oceanic mantle: seafloor and island studies 175
can be applied to infer the vertical conductivity distribution. This technique is particularly useful if telluric data are not available.