Magnetotelluric measurements over the Alpha Ridge

Physics of the Earth and Planetary Interiors, 45 (1987) 101—118 Elsevier Science Publishers B.V., Amsterdam — Printed in The Netherlands

Magnetotelluric measurements over the Alpha Ridge

101

*

E.R. Niblett, R.D. Kurtz and C. Michaud Geophysics Division, Geological Survey of Canada, Energy, Mines and Resources, Ottawa~KIA 0Y3 (Canada) (Received Februamy 10, 1986; revision accepted July 24, 1986)

Niblett, E.R., Kurtz, R.D. and Michaud, C., 1987. Magnetotelluric measurements over the Alpha Ridge. Phys. Earth Planet. Inter., 45: 101—118. Magnetic and telluric field variations were recorded on the Arctic Ocean sea ice over the Alpha Ridge in April and May, 1983. The magnetotelluric measurements were made near the base camp of Operation CESAR (Canadian Expedition to Study the Alpha Ridge) to investigate the deep structure and electrical properties of the lithosphere and upper mantle beneath the ridge. Thirty days of usable data were obtained within the period range 120—20000 s. During the experiment the base station drifted across a deep (2000 m) median valley or graben lying parallel to the ridge axis and then over a steep rise on its northern flank. One data set was obtained over the graben, and the remainder of the measurements were made as the station travelled back and forth over the rise. The directional properties of both the geomagnetic induction arrows and the magnetotelluric impedance tensors were consistent with changes in water depth along the drift path. One-dimensional inversions of E-polarization MT data revealed a resistive lithosphere about 70—85 km thick with a conductive (10 fim) upper mantle (asthenosphere?) below. Sea-water depths estimated in the inversion process agreed remarkably well with the known bathymetry, indicating that the ocean floor contained less than 100 m of higbly conducting sediments along most of the drift path. Two-dimensional models required simulation of the topography of the entire section of the Alpha Ridge between the Canada and Makarov Basins in order that a satisfactory fit be obtained to the H-polarization data. Two-dimensional forward modelling indicates that the anisotropic character of the impedance measurements can entirely be attributed to the bathymetry of the ridge. The data provided no indication of structural inhomogeneity within the oceanic crust and lithosphere, consistent with. results from companion seismic refraction and gravity surveys. The thickness of the lower conducting region was not resolved because reliable long period data (T> 20000 s) were not available for analysis. Impedance estimates at these periods were poorly determined and subject to rapidly rising tensor skew. They were also subject to bias errors caused by water movements, in particular the M 2 tide.

1. Introduction Early in 1983 the Canadian Department of Energy, Mines and Resources (EMR) coordinated a multidisciplinary expedition to investigate the nature and origin of the Alpha Ridge beneath the Arctic Ocean. This ridge is a broad (— 500 km) elevated feature separating the Canada and Makarov Basins and extending northwest from the Canadian continental shelf off Ellesmere Island. Bathymetry suggests the feature is probably *

Geological Survey of Canada contribution no. 25886 and CESAR publication no. 21. -

0031-9201/87/$03.50

© 1987 Elsevier Science Publishers BY.

,

continuous with the Mendeleyev Ridge which extends into the ocean from the East Siberian Shelf. The Polar Continental Shelf Project of EMR established a base camp on the sea ice from which a variety of geological and geophysical observations were undertaken to investigate the structure, composition and tectonic history of the region. Some preliminary results are given by Jackson et a!. (1985). The base camp was located over the central part of the Alpha Ridge at approximately 850 45’N and 110°W,over 650 km northwest of Alert on Ellesmere Island (Fig. 1). Scientific observations extended from 31 March to 23 May, 1983. The drift path, shown in Fig. 2, began over a broad east—west trending depression on the sea

102

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Fig. 2. Drift track of the ice station. The shaded portions indicate intervals for which band-pass or low-pass data were selected for analysis. Bathymetry is in metres.

103

floor which Mudie and Jackson (1985) named the Alpha Ridge Graben. Movement continued in a northeasterly direction over the northern rim of the graben, an abrupt rise of about 800 m. Towards the end of the experiment the camp moved rapidly to the south, away from the high ground, and back across the valley floor. To study the deep structure of this region

ing temperature. The excellent insulation provided by dry snow enabled a low power heater to maintam a temperature above the freezing point inside the instrument box. Air temperatures frequently fell below 40°C. The fluxgate magnetometer has been described by Trigg et al. (1971) and the telluric amplifiers by Trigg (1972). All five components (Er, E~,H5, H,,, H2)

seismic refraction and magnetotellurics (MT) experiments were conducted. The ability of MT to resolve resistivity structure within the crust and lithosphere is much reduced by a layer of sea water, the depth of which varied between 1200 m and 2000 m during the course of the experiment, However, there is one advantage in working on the surface of an ocean. The sea water represents the closest approximation to be found in nature of a truly homogeneous upper layer and effectively eliminates the distortion of measured impedances by near-surface changes in structure or composition. It is therefore of interest to see what MT data can reveal of underlying Earth structure under ihese conditions.

were sampled once per minute and the data stored on cassette tape. Magnetic and telluric outputs were low-pass filtered with the —3 dB points at approximately 120 s. Band-pass filtered telluric outputs were recorded separately with —3 dB points at 120 s and 30000 s (8.3 h) to ensure that telluric data could be obtained which were free from contamination by tidal effects, voltage offsets caused by electrode polarization and other unwanted long period drifts. Using 12 bit A-D converters, we obtained a resolution of 1 nT for the H5 and H,, components and 0.5 nT for H2. Telluric signals were recorded at a resolution of 0.02 mY km’. Thirty consecutive days of good band-pass data were obtained by the end of the experiment. Unfortunately one of the low-pass telluric channels was unstable during the early part of the observation period when the weather was extremely cold. However, this difficulty was eventually overcome and 18 days of good quality low-pass data were finally recorded. Plots of these data revealed high visual correlation between telluric and magnetic horizontal components and very strong attenuation of short-period magnetic H2 fluctuations as would be expected over a substantial thickness of conducting sea water.

2. Experimental procedure North—south and east—west (geographic) electrode pairs were set out on the ice. Low noise silver—silver chloride electrodes (Filloux, 1973) were suspended in the sea water, about 2 m below the lower ice surface, through 10 cm diameter holes. Ice thickness in this location varied between 3 and 5.5 m. The electrode separation was 500 m for each pair, and a fifth electrode was installed at the centre of the array. The magnetometer, telluric amplifiers and two digital cassette recorders were packaged in an insulated aluminium box and powered by air depolarized primary cells which delivered about 5 W for the duration of the experiment. The instrument box was set directly on the sea ice, sealed with an aluminium cover, and buried beneath several feet of snow to keep ternperature variations to a minimum. The three-component fluxgate sensing head was mounted separately on the ice at a distance of about 20 m. The head was also covered with snow to shield it from the wind and to ensure a nearly constant operat-

—

3. Analysis of band-pass data The sea ice supporting the CESAR base camp drifted erratically as shown in Fig. 2 but underwent no appreciable rotation. Its position and azimuth were monitored to a high degree of accuracy several times each day by means of satellite navigation techniques. Four sets of band-passed magnetic and telluric data were selected for analysis (Fig. 2), each set being of 5—8 days duration and corresponding to times when the drift was

104

and quadrature (Q) induction arrows with magnitudes defined by

either relatively slow or confined to a small area. For each of these four intervals both the geomagnetic transfer functions and the MT impedances were computed by the procedure described by Kurtz and Garland (1976). Band averaged autoand cross-power estimates were obtained at various periods for each 48-h record section. Ensern-

1~2,MQ= (A2Q + B~)”2 (A~+ Bfl and azimuths tan O~=B 1/A1, tan 0~= BQ/AQ M1

ble averages were determined from all data in each of the four intervals and used to calculate the impedance tensors and the geomagnetic transfer functions. The period-dependent impedance tensor elements are derived (Sims et a!., 1971) from the linear relationship between telluric and horizontal magnetic components given by

The in-phase arrows so defined will normally point away from regions of high conductivity where induced currents tend to concentrate. In this analysis the directions of all induction arrows have been reversed so that the in-phase arrows will point toward regions of high current density. In-phase and quadrature transfer functions for one of the four data sets are shown plotted against period in Fig. 3. The quadrature amplitudes are only marginally different from zero. The in-phase amplitudes, though also very small (<0.1), appear to be significant over the period range from 102 to iO~s. The azimuth of the in-phase transfer function is most stable and well-determined at periods close to 400 s. Similar transfer function character-

E~= Z55H5 + Z5,,H~+e~ E,, = Z,,5H5 + Z,,,,H,, + e~ The geomagnetic transfer functions are calculated by a method similar to that proposed by Everett and Hyndman (1967) and Schmucker (1970), with assumed linear dependence between the horizontal and vertical magnetic field components of the form

istics are shown by the remaining three data sets selected for analysis. Induction arrows derived from the in-phase functions for a period of 6 mm are plotted in Fig. 4. Each arrow points directly towards deep water and away from topographic

H2 = A H~+ B H1, + e~ .

The complex and period dependent transfer functions, A and B, can be displayed as in-phase (I)

DAYS 99-103 u180~

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Fig. 3. Arimuths and amplitudes of in-phase and quadrature transfer functions (TF) plotted against period for band-pass data corresponding to days 99—103.

105

—

DAYS 99 - 103 1

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BPDATA LOCATION 1

BPDATA LOCATION 3

DAYS 99-103

DAYS 115-122

::~: Fig. 4. In-phase induction arrows for the 6 mm period at locations along the drift track where band-pass data were analysed. Short bars represent ±1 standard deviation in azimuth.

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highs on the sea floor. Thus the in-phase transfer functions, although small in amplitude, were sufficiently well-determined to reflect accurately the influence of local variations in the bathymetry. They provide no evidence of major conductivity contrasts in the crust near the CESAR camp. Figure 5 shows apparent resistivity and phase •

corresponding to major and minor axes of the impedance tensor for two of the four band-pass data sets. The data set for days 99—103 corresponds to a location over the graben floor where the water depth is about 2000 m and ocean bottom topography is relatively smooth but slopes downward to the west with a gradient of about 50 m km’. Both apparent resistivities and phases appear to be nearly isotropic over the frequency range of 102_104 s and the skew is very low, factors which suggest that a one-dimensional model might provide a valid representation of the data. This near-isotropic effect will be discussed in greater detail in a subsequent section. The three remaining data sets were collected at times when the CESAR camp was drifting across an abrupt rise or sea mount on the northern side of the valley, and significant anisotropy is evident in each case at periods greater than 1000 s. This is illustrated in the plots shown for days 115—122 (Fig. 5). The skew of the impedance tensor is small and well-behaved in all four band-pass data .

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Fig. 5. Apparent resistivity, phase, tensor skew and predicted coherencies (E E~)atError two of therepresent locations+1 where bandpass data were5 and analysed. bars standard deviation. —

sets. The data quality is, in general, very good. Predicted coherencies for E5 and E~(Fig. 5) lie between 0.95 and unity over the whole period range from 200 to 10000 s, and exceed 0.98 in the central portion of this band in every case. The azimuths of the principal axes of the impedance tensor (not shown) are stable and well-determined and vary little over the two decades of available data. Figure 6 shows the direction of the major axis of the impedance tensor corresponding to a period of 17 mm for the four band-pass filtered data sets. In each case the direction is well-determined and lies transverse to the sea-floor topography and roughly parallel to the direction of the in-phase induction arrow at the same location. Both the magnetic and MT response functions are therefore clearly influenced by changes in water depth that occurred along the drift path.

106

DAYS 99 - 103 1 107- 113 2 115-122 3

124

One-dimensional interpretations of the apparent resistivity and phase data have been de-

DIRECTION OF MAXIMUM IMPEDANCE period-li mm

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and using Fischer the inversion and LeQuang method (1981). ofband-pass Fischer Bothetmajor al. (1981) and rived the been four sets minor for axiseach data of have treated and data for each case the resolution of the estimated depths and

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decomposition resistivities has(SVD) been (Edwards examined et byal.,singular 1980; Chave value et al., 1981; Jones, 1982a; Ilkisic and Jones, 1984). The resistivity of the sea water was fixed at 0.3 ~Im at the outset; all other parameters were derived in the inversion process. The results are

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summarized in Table I. The following points are worth noting regarding these results. (1) The short period phase data are not very

Fig. 6. Direction of the major axis of the impedance tensor at the 17 mm period at the four locations where band-pass data were analysed. Error bars are ±1 standard deviation in azimuth.

TABLE I Inversions of MT data Loca-

Days

Axis

tion Band-pass data 1 99—103

107—113

3

115—122

4

124—129

Water

Pi (Fixed)

d

1

SVD

P2

SVD

d2

SVD

inverted

depth (m)

(clm)

(m)

(d1)

(~2m)

(P2)

(km)

(d2)

p~ SVD (~2m) (ps)

0.3 0.3 0.3 0.3 0.3 0.3 0.3 0.3 0.3 0.3 0.3 0.3 0.3 0.3 0.3 0.3

1735 1726 1959 1921 1427 1477 1545 1537 1338 1372 1498 1513 1369 1379 1510 1553

1.00 1.02 1.00 1.05 1.02 1.02 1.02 1.02 1.05 1.05 1.02 1.02 1.02 1.02 1.05 1.00

5,350 3,245 2,829 4,366 689 20,400 688 663 792 188,000 448 1,230 1,673 5,582 241 3,683

10.0 10.0 10.0 10.0 9.8 10.0 10.0 10.0 9.6 10.0 9.8 10.0 10.0 10.0 7.6 10.0

106 108 101 87 110 137 84 77 161 165 75 75 143 147 85 86

1.12 1.26 1.10 1.62 1.26 1.22 1.05 1.09 1.17 1.54 1.05 1.09 1.15 1.25 1.15 1.08

5 7 25 37 45 41 2 7 25 42 4 7 34 39 9 9

1.86 3.02 1.29 2.14 1.41 1.50 1.55 1.36 1.70 3.50 1.23 1.32 1.62 1.95 1.41 1.30

0.3 0.3 0.3 0.3

1353 1352 1479 1451

1.02 1.02 1.00 1.02

1,110 2,015 1,909 2,415

9.3 9.8 10.0 10.0

150 151 68 55

1.32 1.20 1.05 1.17

90 63 9 11

1.35 1.38 1.15 1.32

-

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Data

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Pa + 4)

Low-pass data 110—113 118—125

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Data sets were inverted at each location as follows: major axis apparent resistivity (Pa); major axis apparent resistivity and phase combined (Pa +4)); minor axis apparent resistivity; and minor axis apparent resistivity and phase combined. d1, Pi represent the depth and resistivity of the sea water. d2, P2 represent the depth and resistivity of the second layer. P3 represents the resistivity of the third layer. SVD stands for singular value decomposition and indicates the limits within which a parameter is resolved by the process of inversion. For example, if d2 =106 km and SVD (d2) =1.12, then log d2 log 106±log 1.12. For SYD ~10 the parameter is essentially unresolved.

107

reliable. At all four locations the phase values at periods below 600 s are too low to provide a reasonable fit to the known water depths. The low values could be explained by a 10% change in the filter RC constant at the low operating temperatures which prevailed on the sea ice. The phase data also show a rather high degree of scatter at the long period end of the range. (2) Apparent resistivities are generally stable and well-determined at all periods. Cross-power coherencies between E5 and H,,, E,, and H5 and predicted coherencies for E5 and E,, are close to unity (— 0.98 on the average), confirming the high quality of the apparent resistivity data. Small changes in filter characteristics would have very little effect on short period apparent resistivity values. (3) At locations 2, 3 and 4, where the data are anisotropic, the minor axis apparent resistivities and phases are a better match to the near-isotropic data of location 1 than the major axis results. At each of these stations the minor axes of the impedance tensor lies roughly east—west and parallel to the bathymetric trend (Fig. 6), the direction which we would identify with E-polarization in a two-dimensional model representation. The minor axis data are therefore the least affected by topographic irregularities on the ocean floor, (4) The estimated water depths are in good agreement with the known bathymetry, particularly for the minor axis data. There is no evidence for appreciable thicknesses (>100 m) of highly conducting sediments or unconsolidated material on the sea floor in the vicinity of the CESAR drift path. (5) SYD shows that, next to the water depth, the thickness of the second layer is the best resolved parameter. The resistivity of the lower conductmg layer is also fairly well resolved, while that of the second layer is essentially unresolved. This is to be expected when a resistive layer lies between two conductive ones (Wait, 1962; Cavaliere and Jones, 1984). The SYD analysis also shows in every case that the parameters d2 and p3 are better resolved when apparent resistivity only is inverted, rather than apparent resistivity and phase combined. This result supports the conclusion that

the short period phase estimates are unreliable for all band-pass data sets. (6) The best parameter estimates are those obtamed using the E-polarization apparent resistivity data at each of the four locations. Averaging these, we obtain the results: 2nd

layer

p2

—

1700 ~m

d2

22

88 km

layer p3 5 ~2m Note that E-polarization corresponds to the major axis of the impedance tensor at location 1 (over the graben), and to the minor axis at locations 2, 3 and 4. The value of p2 is determined only to the extent that it must be substantially greater than either p1 or p3. 3rd

4. Analysis of low-pass MT data The low-pass MT data were expected to provide information at periods longer than 10 000 s to aid in the interpretation of electrical properties within the crust and upper mantle beneath the Alpha Ridge. The best 12 days of low-pass data (days 110—113, 118—125) were selected for analysis. These days corresponded to times when the CESAR camp was moving across the sea mount and drift velocities were relatively low (Fig. 2). The averaged MT response over these 12 days is shown in Fig. 7. The apparent resistivity and phase data ‘are considered to be satisfactory within the period range 120—20000 s (5 hours). Predicted coherencies for both E5 and E,, are about 0.95 over this range and above 0.98 for more than a decade in the central portion of the band (Fig. 7, Table II). Longer period responses are not welldetermined and are characterized by rapidly rising skew, suggesting the onset of large bias errors (Pedersen and Svennekjaer, 1984). The M2 tide may be partially responsible for this. Thus for the present analysis periods longer than 20000 s were not considered acceptable. When we compare the low-pass apparent resistivity curves with those derived from the near-isotropic band-pass data (days 99—103), it is clear that after allowance is made for the difference in ocean depths between the two locations, the minor ~-

108

L P DATA

data alone are inverted rather than apparent resistivity and phase combined. For this case the sig-

DAYS

mficant results are: 2m(~layer p2—2000~2m d2~68km 3rdlayer p3~9~2m The response of this model is shown superimposed on the minor axis data (both apparent resistivity and phase) in Fig. 7. The one-dimensional inversions of all available E-polarization :aata sets therefore indicate a lithosphere beneath the Alpha Ridge which is a little less than 100 km thick underlain by a conductive layer (asthenosphere?)

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of unresolved thickness and a resistivity of close to 1 id60

5. Two dimensional interpretations

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One-dimensional interpretations are unsatisfactory for a variety of reasons, the most important of which are outlined below. (1) All the MT data, both band-pass and low-

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pass, are responsive in water’ depths (Figs. 4 and 6). This to is variations the case even for the near

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Fig. 7. Apparent resistivity, phase, tensor skew and predicted coherencies for the 12 days of low-pass data available for analysis. The solid curves show the response of the model (see Table I) obtained by inversion of the minor axis (E-polarization) apparent resistivity data.

apparent resistivity curve of the low-pass data most closely matches the apparent resistivities derived from the near-isotropic band-pass data. Again, this indicates that for a one-dimensional interpretation, inversion of the minor axis (Epolarization) apparent resistivity and phase data should give the most reliable results. Table I shows the parameters derived with the Fischer et al. (1981) inversion method and a pre-determined value (0.3 ~lm) for the resistivity of the sea water. Again the SVD analysis indicates that d 2 and p3 are better resolved when the apparent resistivity

isotropic band-pass data where both the impedance tensor and the geomagnetic transfer functions could be related to ocean bottom topography. Therefore the assumption of uniform water .

.

depth is an over-simphfication. (2) The sea-floor topography is very irregular where the low-pass data were acquired and a satisfactory solution would require a reasonable fit to both the E-polarization (minor axis) and Hpolarization (major axis) data. In view of these limitations it is necessary to investigate two- and three-dimensional models to see what measure of agreement these provide with the one-dimensional results. Figure 8 shows the 12 days of averaged low-pass MT data superimposed on the apparent resistivity and phase responses derived for the two-dimensional model shown. A resistive layer of 1000 ~2m (lithosphere) extends to a depth of 87 km, while the underlying conductor (asthenosphere) is assigned a value of 10 ~lm. The model responses were calculated from a transmission surface ana-

logue, originally proposed by Swift (1967), for an observing station centered over the top of the structure. The data, on the other hand, were derived from observations collected while the station drifted over both flanks of the sea mount as well as its top. The fairly large error bars evident in the phase results may be caused by this unavoidable drift. Nevertheless the model provides a reasonable fit to the apparent resistivities and the Epolarization phase. The H-polarization model response lies slightly above the corresponding apparent resistivity curve at periods between 120 and 3000 s, but this effect disappears if the observation point is shifted from the top of the sea mount to one of its flanks. Both E-polarization and Hpolarization phase responses coincide at long periods for this model, an effect which is not evident in the phase data. H-polarization phases appear to be biased toward lower values at periods above 2000 s. If the effect is not caused by spurious data, then a more complex model is required to satisfy the H-polarization phase measurements.

Fig. 8. Low-pass MT data superimposed for the two-dimensional model indicated.

on response

curves

ALPHA RIDGE MODEL

0

I

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60km

I

SEA /CE

CANADA BASIN

d cab 4 441

MAKAROV BASIN

2OOOm

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Fig. 9. Two-dimensional forward model for a north-south section through the CESAR ice station of the entire Alpha Ridge between the Canada and Makarov Basins. Model responses were computed for stations located at positions a, b, c and d.

110 102

102

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111

To seek a better explanation for the H-polarization data, we extended the two-dimensional model to include: (a) the broad topographic outline of the north—south section of the entire~Alpha Ridge between the Canada and Makarov Basins; and (b) the approximate bathymetry of the steepsided east—west trending Alpha Ridge Graben that bifurcates the ridge crest (Mudie and Jackson, 1985). This extended model is illustrated in Fig. 9. Model responses were computed for a station located over the top of the rise (Fig. lOa), over its northern flank (Fig. lOb) and over its southern flank adjacent to the median valley or graben (Fig. lOc). We see at once that the E-polarization (minor axis) and H-polarization (major axis) data are well matched by the response curves, both apparent resistivity and phase, for either flank of the rise. For a station located over the top of the rise the phase data and the minor axis apparent resistivity data are a good fit to the model, but the H-polarization apparent resistivity response is substantially higher than the major axis Pa data. The model response thus provides a better fit for a station located over either flank of the rise, %vith the north flank location probably providing the best fit overall. This result is consistent with the drift pattern, since the base station lay directly over the top of the rise for only 2 of the 12 days in which low-pass data were available. The adopted two-dimensional forward model (Fig. 9) is therefore in excellent agreement with the low-pass MT data (and also with band-pass data sets 2, 3 and 4). The long-period separation of the E- and Hpolarization phase data has been explained by modelling the broad structure of the whole Alpha Ridge rather than just the local bathymetry. The near-isotropic band-pass data set (days 99—103) was obtained when the base station was drifting over the central part of the median valley, Figure 11 shows the two-dimensional model response for a station located in the centre of the trough (Fig. 9, position d), superimposed on the major and minor-axis band-pass data corresponding to these days. Again the match is good for both apparent resistivity and phase, except that the phase data at periods of 300 s and less are too low as mentioned earlier. Note that both the ap-

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Fig. 11. MT response of the model indicated in Fig. 9 corresponding to a station at position d superimposed on the band-pass data from location 1 on the drift path.

parent resistivity and phase responses cross over at periods of about 5000 and 1000 s, respectively. The apparent resistivity for the E-polarization mode (parallel to the graben) is slightly larger for periods between 200 and 5000 s for both the model response and the data. The isotropic appearance of the data is therefore seen to be a consequence of the station location over the valley floor, and does not reflect a truly homogeneous and one-dimensional structure. To test the validity of the two-dimensional results, we constructed a three-dimensional model which more accurately represents the dimensions and geometry of the rise or sea mount that the base camp crossed over three times during collection of the low-pass data (Niblett and Jones, 1985). The model responses were computed at periods of 1000 and 10000 s using the method of Park et al. (1983) and found to deviate little from those of the two-dimensional model depicted in Fig. 8. The validity of the two-dimensional model for this particular structure is therefore confirmed. -

112

6. Discussion Two naturally-occurring processes limit the application of simple induction theory in a polar oceanic environment. The first of these concerns the scale-length at the Earth’s surface of source fields generated in the ionosphere. The second concerns the production of electric fields within the ocean itself as a consequence of dynamo interaction between sea water velocities and the ambient geomagnetic field. The impact of source fields of finite size on MT data and interpretation has been investigated by Price (1962) and Wait (1962). For a stratified three-layer Earth, Wait concluded that the plane wave or uniform field assumption is not justified unless the condition —1/2

L>> (~oaaw) (1) is fulfilled. Here L is the spatial wavelength of the source field, ~ is the permeability of free space and ; is the apparent conductivity at a frequency w. At a period of 20000 s, the longest used in the analysis of CESAR data, we have; 0.1 S m 1, w = ir x i0~~s~ and ~t0 = 4ir x iO—~H m’. Wait’s condition therefore reduces to L>> 160 km. Accordingly, source fields smaller in horizontalextent than about 1600 km would fail to justify the uniform held assumption and would require that a more complex pretation of the data. procedure be used in interThe spatial extent of the horizontal components of magnetic field variations in the Queen Elizabeth Islands has been investigated quantitatively by DeLaurier et al. (1974). This region lies well within the polar cap regime and far to the north- of the highly localized current systems and electromagnetic fields associated with the auroral zone. For .a period of 18000 s DeLaurier et al. (1974) determined both H~and H~component coherencies at various recording stations with respect to Mould Bay observatory on Prince Patrick Island. For both components all measured coherencies were above 0.8 for station separations of 400 km and less, and above 0.6 for separations of 1000 kin and less. Thus at the 0.8 level the source field scale length is about 800 km, and at the 0.6 level about 2000 km. Higher measured -

coherencies would be expected over the Arctic Ocean owing to the absence of induction effects caused by island coastlines and irregularities in the local geology. It seems reasonable, therefore, to conclude the uniform field assumption is justified at periods up to 20000 s on the basis of Wait’s criterion. At longer periods this may no longer be so, although DeLaurier et al. (1974) observed that source field scale-lengths tend to increase with period. Dmitriev and Berdichevsky (1979) pointed out that the condition determining the applicability of the uniform field assumption in one-dimensional interpretation depended on the linearity of the horizontal field changes rather than their rate of change. This relaxes the stringency of condition (1) to some degree and increases the justification for neglecting source-field restrictions in the analysis of the CESAR data. Data obtained in a conventional MT experiment on sea ice are prone to contamination by electromagnetic fields generated within the ocean itself which differ in spatial dimensions and spectral content from the normal far-field external sources. Ocean-generated fields are far from negligible, particularly in polar regions where the vertical magnetic component is large. In a horizontal direction the ~ x B electric field for a velocity of 0.01 m s’ and ambient magnetic of 2 isan equivalent to 0.6 mYfield hn1. 0.6 x exceeds iO~Wb This themE-field resolution of the CESAR telluric recording system by a factor of 30. Even very small fluid velocities can therefore have an impact on the measured field variations. Furthermore, at a period of 12 h, the tellunc field amplitude induced in a one-dimensional Earth- of apparent resistivity Pa = 10 ~lm (Fig. 5), and corresponding to a typical magnetic field amplitude of 10 nT, is 0.3 mV km’. It is therefore clear that M 2 tidal velocities of only 1 cm s’ can produce larger electric fields than magnetic variations originating in the ionosphere. To avoid spurious effects generated by ocean currents, internal waves and .tides, Chave and Filloux (1984) separated MT data acquired on the Pacific Ocean floor (Filloux, 1980) into parts which were coherent and incoherent with the honzontal magnetic field at Tuscon, Arizona, 1700 km distant but at the same geomag-

113

netic latitude. The correlated part of the sea-floor data they attributed to ionospheric sources; the uncorrelated part to oceanic sources. Such a procedure was not attempted for the CESAR data. The closest observatory is at Alert, some 650 km distant on the northern tip of Ellesmere Island. This station is highly anomalous (Niblett et al., 1974) and is likely to be influenced by tides because of its coastal location. Mould Bay and Resolute Bay are also coastal stations, and College, Alaska is nearly 2500 km distant and lies within the auroral zone where magnetic variations differ greatly from those within the polar cap. For the CESAR experiment it is therefore important that the data used for analysis are, insofar as we can determine, free from bias errors which could be caused by superposition of electric and magnetic fields generated by tidal velocities or wave motions, noisy data, changes in water depth along the drift path and temperature variations at the recording site. The analytical and experimental procedures described previously were designed to minimize bias effects as far as possible in a single station experiment. Aagaard (1981) describes water velocity measurements made at LOREX in deep water (25 and 200 m above the bottom) over the Lomonosov Ridge and at the base of its northern flank. Mean velocities were found to be 2—3 cm s~’in the former location and 0.4 cm s~ in the latter. However, more energetic episodes with peak speeds of 12 cm s ‘over the ridge and 4 cm s ~ in the abyss were occasionally observed. Semi-diurnal oscillations corresponding to the M2 tide were -

—

clearly seen in the energy spectra with maximum amplitudes of 2—2.5 cm s’. At the CESAR location the sittiation was quite different (K. Aagaard, personal communication, 1985). Current flow meters deployed at depths of 100, 400 and 1500 m near the base camp revealed average velocities well below 2 cm s’, occasional episodes with flows up to 5—6 cm s~, and much smaller tidal velocities than at LOREX. Contamination of MT data by tidal oscillations in the ocean was therefore substantially reduced during the Alpha Ridge experiment, but would probably have been a significant factor had the data analysis extended to periods of 12 h or longer. Band-

width calculations indicate that for a centre frequency of 5.5 h, the longest period used for interpretation, energy corresponding to an 8 h (tidal) period is reduced by a factor of five, while that corresponding to a 12-h period is reduced by 600. The fields generated by the lunar semi-diurnal tide were therefore essentially eliminated from both the band-pass and low-pass impedance estimates at periods of 20000 s and less. Surface gravity waves associated with winds and ocean swell are greatly attenuated by ice cover. Typically the amplitude of motion of a thick ice flow varies from a few hundredths to a few tenths of a millimetre over periods ranging from a few seconds to several hours (Hunkins, 1962; LeShack and Flaubrich, 1964). From gravity measurements on the Arctic Ocean sea ice Weber (personal communication, 1986) has estimated that the velocities associated with this motion are very small (— 0.02 cm s’). Magnetic fields induced by surface waves are therefore not likely to be an important source of error in MT measurements made from a sea-ice platform. However, internal gravity waves owe their existence to stable stratification in the oceans (Munk, 1981) and are probably not much affected by ice cover. Chave (1984) has obtained theoretical spectra of electromagnetic fields generated by ambient internal waves and found that spectral levels are very high in the interior of the ocean. Boundary effects reduce the spectra by orders of magnitude at the sea surface, but even here internal wave-induced magrietic fields can be within a factor of 10 of the fields derived from ionospheric sources in the period range from about 30 mm to 1 day. Chave and Fffloux(l984) found that internal waves were a significant source of induction in their sea-floor experiments in the deep ocean off the California coast. For the CESAR data internal gravity waves are a possible source of magnetic field contamination over much of the period range available for interpretation. The impedance tensor elements for the low pass data have therefore been examined in detail for evidence of bias. To implement this procedure, we computed the measured impedance estimates Z~ and in four different ways as outlined by Sims et al. (1971)~The appropriate -

114

formulae, expressed in terms of coherencies rather than auto-power and cross-power density spectra are

coh H~E,,coh E~H,, coh H~H~coh EKE,, coh H~ E,, coh H~H~ coh H,,E,, —

X

—

IEX1

(3)

4~

lxyIH coh H~E~coh E~E~ coh H~E,, —

~ =

X

(1)

~JJ~ixY

I H,, I

2xy Z

=

h

—

X

X

IE I ~7~’~2xy

Y

ii l~~-’

(4)

E*)?/l 2 X / ~v ‘H H*~i/ Equations 1 and 2 for Z~ are biased upward by random noise on the E signal but are not biased by random noise on H (Sims et al., 1971). Equations 3 and 4 are biased down by random noise on H, but not affected by random noise on E. In this study impedances were calculated using

coh H~E~ chEH oh H,,H~ 2_i coh H,,E~

IX

=

IE I coh E H —coh H H coh E H Z4~,,= X ~ coh H H 12

(2)

Y

I E, I

fill

IE )‘ I —s---— “(E.“X

IH

TABLE II Stability of impedance estimates Period (s)

tDi~~ ~~2xy

~~3xy

~~4xy

~1~xy

S((~~) ~lyx

~yx

19475 15470 12288 9761 7753 6159 4892 3886 3087 2452 1948 1547 1229 976 775 616 489 389 309 245 195 155 123

1.008 1.048 0.979 1.009 1.019 1.017 0.990 0.998 0.999 0.998 0.992 0.986 1.033 1.032 1.009 1.007 1.003 0.999 1.006 1.003 1.013 1.013 1.011 1.012 1.008 1.007 1.007 1.007 1.014 1.014 1.013 1.013 1.010 1.009 1.013 1.012 1.013 1.012 1.018 1.017 1.025 1.024 1.042 1.041 1.056 1.056

0.891 0.904 0.941 0.901 0.945 0.933 0.961 0.970 0.978 0.982 0.984 0.994 0.990 0.992 0.997 0.995 0.994 0.993 0.993 0.991 0.985 0.971 0.952

0.944 0.936 0.942 0.913 0.947 0.932 0.961 0.969 0.975 0.980 0.985 0.995 0.989 0.992 0.997 0.996 0.992 0.992 0.993 0.990 0.985 0.970 0.952

0.973 0.957 0.980 0.950 0.972 0.961 0.997 0.989 0.989 0.993 0.999 1.003 0.998 1.000 1.006 1.004 1.001 1.002 1.003 1.004 1.005 1.006 1.004

0.796 0.857 0.855 0.832 0.898 0.889 0.866 0.925 0.952 0.954 0.945 0.967 0.965 0.970 0.967 0.966 0.968 0.961 0.962 0.948 0.924 0.868 0.813

0.976 0.741 0.944 0.834 0.989 0.935 0.938 0.876 0.954 0.903 0.988 0.923 0.986 0.960 0.990 0.959 0.978 0.942 0.998 0.968 0.980 0.957 0.994 0.980 0.997 0.985 0.974 0.962 0.980 0.969 0.989 0.980 0.989 0.978 0.994 0.982 0.998 0.980 1.005 0.976 1.014 0.967 1.034 0.954 1.055 0.945

0.904 0.903 0.989 0.920 0.952 0.994 0.986 0.992 0.983 1.002 0.980 0.992 0.999 0.974 0.980 0.989 0.991 0.996 0,999 1.006 1.015 1.035 1.055

~3y~

~4yx

‘Z~

0.785 0.852 0.864 0.886 0.935 0.962 0.887 0.905 0.905 0.928 0.922 0.959 0.960 0.973 0.959 0.975 0.940 0.961 0.966 0.984 0.958 0.969 0.981 0.987 0.984 0.991 0.962 0.968 0.969 0.974 0.980 0.984 - 0.976 0.984 0.982 0.988 0.979 0.989 0.975 0.990 0.967 0.991 0.953 0.994 0.945 1.000

S(I~~)DF

coh coh (E~E~)(E~~’E~)

0.659 0.845 0.894 0.900 0.900 0.867 0.948 0.936 0.921 0.935 0.955 0.975 0.973 0.976 0.978 0.982 0.974 0.974 0.962 0.941 0.909 0.850 0.802

0.963 0.979 0.964 0.968 0.976 0.976 0.971 0.983 0.990 0.990 0.988 0.992 0.991 0.994 0.993 0.993 0.993 0.991 0.991 0.987 0.981 0.967 0.950

10 4 28 36 32 56 60 94 108 132 180 214 276 348 432 552 684 870 1088 1324 1436 1120 454

0.951 0.981 0.977 0.981 0.979 0.979 0.988 0.986 0.984 0.989 0.990 0.994 0.994 0.995 0.995 0.996 0.995 0.994 0.991 0.986 0.977 0.961 0.948

~2~y etc. are defined by eqs~1—4. O~is the mean value ~nd S(~~)is the stability factor defined by eq. 5. DF depicts the number of degrees of freedom in each frequency band and coh (E~”E~) is the coherency between E~(measured) and E,~” (predicted). ~i~y’

115

eq. 4 on the strength of the usual assumption that the electric field is likely to be more contaminated with noise than the magnetic field. Table II shows a comparison between the coherency terms ~ ~2xy etc. in each of the four equations. For perfectly coherent and unbiased E and H data E =

1 and

=

I H,, I and all four relations are equivalent. Bias caused by cultural noise, electrode polarization and other instrumental drifts, sea-water motions etc. will cause 1 to differ from unity. Table II also gives the mean values 4~ and the stability factor, s, defined by Sims et al. (1971) as

s

=

[z3~,,j [z4~,,} [z,~,,J [z2~,,]

(5)

This number will be unity if the four estimates of are identical, and will decrease as the estimates diverge. Values of cI~, their means and their stability factors are also included in Table II as are the predicted coherencies for E~and E,, and the number of degrees of freedom for each frequency band. Zr,, corresponds to the minor axis of the impedance tensor (E-polanzation) and Z,~to the major axis. The effects of bias are clearly evident in the data, but these are for the most part very small and well within the error levels associated with the impedance estimates. ~1xy’ ~2xy are within 2% of unity over most of the frequency image. tb3xy and ~4xy values are all less than one; they are down by 5—9% at periods exceeding 6000 s, but by less than 3% over most of the remaining bands. For the major axis results, ~ and ~2yx values are very close to unity except those with the longest periods, while ~3yx and ~ are consistently lower by a few percent. Again, bias is indicated by noise in the data, although the effect is large only at the two longest periods, It is not possible to determine the bias contamed in the common term I E~I/I H,, I in eqs. 1—4. However, the clear evidence of bias in the coherency terms (Table II) indicates that internal gravity waves could have contributed in a small way to the measured H fluctuations in the period

range used for interpretation. The CESAR location was removed by hundreds of miles from major sources of cultural noise which so often contaminate E-field data acquired on land. Fortunately the downward bias in the minor axis impedance estimates (Z4~,,)is less than 10%, even at the longest periods. The small number of degrees of freedom associated with the long-period estimates is partially responsible for their lower reliability. If eqs. 1 or 2 had been used to calculate the impedances, the apparent resistivity values for the low-pass data would have been slightly larger at periods longer than 6000 s, but still within the error bars presented in the figures. The main effect would be a decrease of about 15% in the estimated thickness of layer 2 (see Table I), and an increase in the resistivity of layer 3 to about 15 ~l m for the minor axis data. The generally favourable comparison between band-pass data and low-pass data sets, the urnformly high values of predicted coherency, and the well confined low values of the tensor skew over the period range of 100—20000 s indicate that the interpreted data are trustworthy. The depth of 70—85 km to the top of the conducting layer is reasonably consistent with one-dimensional inversions of both low-pass and band-pass minor-axis data and is compatible with the two-dimensional and three-dimensional models we have fitted to the apparent resistivity and phase results. Neither the thickness nor the conductance of the lower layer could be resolved with the available data. Its resistivity appears to lie between 5 and 15 ~2m according to the model calculations, a value which agrees well with results from other investigations (Vanyan et al., 1978; Law and Greenhouse, 1981; Chave et al., 1981; Jones, 1982b; Chave and Filloux, 1984). The existence of an ‘electrical asthenosphere’ or low resistivity layer beneath both continental and oceanic lithospheres has been indicated in many previous investigations. In general the conductive zone is thought to reside at shallower depths beneath the oceans than beneath the continents, and a correlation between the age of the lithosphere and the depth to the conducting layer has been inferred for the north Pacific region by Filloux —

116

(1980), Oldenburg (1981) and others. Such a correlation is reasonable in an ocean basin where both the age and thickness of the lithosphere depend on its distance from an active spreading centre~ Oldenburg (1981) indicated that the data of Law and Greenhouse (1981) and Filoux (1977, 1980) are consistent with this hypothesis. However, Oldenberg et a!. (1984) later found that the lithosphere thickness vs. age correlation was not as well developed as they previously thought, owing to large inherent non-uniqueness in interpretation of the data. In much of the Arctic Basin the origin of oceanic crust and lithosphere is not well understood. The Nansen—Gakkel Ridge is an active spreading centre and the adjacent Fran’i and Nansen Basins are presumably underlain by lithosphere that thickens with distance from the ridge axis. However, the Fram Basin is bounded by a massive parallel feature, the aseismic Lomonosov Ridge which may be expected to have a profound effect on deep structure. Camfield’s results from the LOREX experiment (P.A. Canifield, personal communication, 1985) suggest a lithosphere nearly twice as thick below the Lomonosov Ridge (120—150 km) as below the Fram Basin (70—80 km), which lends support to the theory that the former structure is a continental fragment separated from the Barents shelf after spreading began in the Eurasian sector (Sweeney, 1985). On the Amerasian side of the Lomonosov Ridge lies the Makarov Basin, the Alpha—Mendeleyev Ridge and the Canada Basin. Geological, biological and palaeomagnetic data acquired during the CESAR experiment from core specimens indicate a late Cretaceous age for the Alpha Ridge (Jackson et al., 1985). The nature of its origin is uncertain, but is thought to be associated with early stages of the opening of the Canada Basin. Geological and geophysical evidence relating to the age and formation of this basin have been reviewed by Sweeney (1985). A combination of different constraints led to the result that the basin sea-floor was mostly produced during the period 118 to 83 Ma, i.e., mid to late Cretaceous. The Alpha Ridge has not been an active spreading centre in recent time and DeLauner (1978) has argued that it shows too high a relief to have ever been one. Sweeney et al.

(1978) have suggested that it may have been produced by buckling of young oceanic crust and Forsyth et al. (198t~have proposed a hot spot origin on the oasis of similarities in crustal structure with Iceland. Thus there is good reason to believe the Alpha Ridge developed by deformation of the oceanic crust approximately 100 Ma ago. It is uncertain that it was ever an active spreading centre and the nature of its origin is still a matter of speculation. It is therefore difficult to infer a simple relationship between age and lithosphenc thickness as Filloux (1980) and Oldenburg (1981) were able to do in the Pacific. Filloux found 30 Ma lithosphere underlain by highly conducting material at a depth of 85 kin, while for 72 Ma-old lithosphere the corresponding depth was 140 km. We conclude that lithosphere beneath the Alpha Ridge is much thinner than lithosphere in the northeast Pacific of similar age. Russian deep soundings from Arctic Ocean sea ice have been reported by Trofimov and Fonarev (1976), Trofiinov (1979), and other workers. Some of the results of these studies have been reviewed by Fonarev (1982). Trofiinov and Fonarev (1979) obtained depths to the upper-mantle conductor of 165 km near the edge of the Chukchi Shelf, 300—320 km at the Siberian end of the Lomonosoy Ridge, and 325 km on the Amerasian side of the Lomonosov Ridge near the Wrangel Plain. In the Canada Basin off the southwestern flank of the Alpha Ridge, Trofimov (1979) obtained a depth of about 200 km. Magnetovariational soundings reported by Fonarer (1982) at two positions on the Mendeleyev Ridge yielded a similar depth to the upper-mantle conductor. The Soviet MT results therefore indicate a thicker lithosphere beneath the Arctic Ocean than was found in either the CESAR or LOREX experiments. Comparisons between these studies may be misleading, however, because different procedures were used in the collection and analysis of the MT data. For example, in some of the Russian work it is not very clear how the problem of ocean-generated fields has been dealt with. Fonarev (1982) discussed this effect explicitly in his review, however, indicating that the longer period data must have been corrected in some way.

117

Conclusions

Acknowledgments

One-dimensional inversions of E-polarization MT data are compatible with a three-layered structure in which the middle layer is taken to represent the oceanic lithosphere beneath the Alpha Ridge. The thickness of this layer as derived from the inversions is 70—85 km and this is consistent with more realistic two and three-dimensional model representations of the sea-floor bathymetry. A more precise estimate is not possible because measured values of the impedance tensor appear to be biased downward by small amounts. The effect is most pronounced at the longest periods and is probably caused by internal-wave motions and their associated magnetic fields. To provide an adequate match to the observed H-polarization data, we require a two-dimensional model that includes a broad topographic section of the whole Alpha Ridge. The directional properties of both the impedance tensors and the magnetic transfer functions derived at various locations along the drift path confirm that these response functions are sensitive to variations in the depth of the sea water. The resistivity of the lithospheric layer, sandwiched between highly conducting sea water above, and highly conducting upper mantle below, is not well resolved. Singular value decomposition (Edwards et al., 1980; Jones, 1982a) indicates that this layer could be substantially more resistive than the 1000 ~m indicated in the two-dimensional model. Neither was it possible to resolve crustal structure within this region, although the good agreement between derived and measured water depths supports the conclusion that the ocean floor, along the drift path, contains only a thin upper layer of highly conducting sediments and/or unconsolidated matter. However, the thickness of the intermediate layer was quite well resolved as was the resistivity of the lower layer (— 10 ~lm). The seismic refraction experiments described by Forsyth et al. (1986) indicate a laterally homogeneous upper crust beneath the Alpha Ridge. The air gun reflection data (Jackson et al., 1985; Forsyth et al., 1986) confirm that sediment thicknesses varied from zero to 1 km beneath the base station, the accumulation being greater over the graben than over the elevated regions.

We are grateful to R. Groulx for his help in assembling, testing, calibrating and packaging the MT recording system. F. Andersen assisted with its installation on the sea ice under rather severe conditions, and V.S. Allen inspected the equipment and kept it operating during the latter part of the observation period. The experiment could not have been done without the support these individuals provided. We thank A.G. Jones for his keen interest in the project and for the many valuable discussions we had with him during the analysis and interpretation phase. The suggestion to try a three-dimensional forward model was his, and he provided software for both this and the SVD analysis. We also gratefully acknowledge useful discussions with P.A. Cainfield, K. Aagaard and A.G. Green. Transportation and logistical support for CESAR was provided by the Polar Continental Shelf Project and the Canadian Armed Forces.

—

-

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