Observed Magnetic Anomalies

The most prominent magnetic anomaly in the Mesozoic crust of the Newfoundland-Iberia rift is the so-called "J Anomaly." This anomaly occurs over crust from just seaward of anomaly M1 to just seaward of M0 in TE1, and it has been attributed to a thickened magnetic layer, to unusually high magnetization, or to both (Rabinowitz et al., 1978). The anomaly amplitude is highest (~1000 nT) over the Southeast Newfoundland Ridge and conjugate parts of the Madeira-Tore Rise at the southern edge of the rift, and it decays to both the north and south (Tucholke and Ludwig, 1982; Tucholke et al., 1989; Miles et al., 1996). Within the Newfoundland-Iberia rift, the anomaly amplitude reaches much lower, relatively constant levels (~200 nT) just north of the southern margin of Galicia Bank. Thus, if the unusual magnetization of the J Anomaly is associated with melt distributed from the plume under the Southeast Newfoundland Ridge, the plume magmatism may have affected the rift as far north as Galicia Bank and Flemish Cap.

Landward of the J Anomaly, the broad negative magnetic anomaly M3 is readily recognized. This anomaly has similar amplitude on both sides of the rift, but the amplitude is half that observed in the Central Atlantic Ocean (Srivastava et al., 2000). Magnetic anomalies older than M3 have very low amplitudes (~100 nT or less) (Miles et al., 1996) compared to their Central Atlantic counterparts. In addition, the amplitudes are higher on the Newfoundland side than on the Iberia side, where anomalies are barely visible in surface-ship data. Although pre-M3 magnetic anomalies are weak, they are parallel to subparallel to the margins and appear to be symmetrical across the rift. Thus they bear some of the characteristics of seafloor-spreading anomalies, which led Srivastava et al. (2000) to identify them as part of the M-series sequence.

Figure F7 shows examples of magnetic profiles from the conjugate margins of the rift near the southern margin of Flemish Cap and Galicia Bank. The conjugate sea-surface profiles in Figure F7A and F7B show general symmetry with one another. Deep-tow profiles along the western part of the Leg 149/173 drilling transect (Fig. F7D, F7E) show higher-amplitude anomalies. Comparison of the sea-surface and deep-tow magnetic anomalies with model calculations shows good correlation for anomalies M1 to ~M8 but no clear correlation for older anomalies. However, the amplitudes of even these observed anomalies can be fit only by using model magnetization values that vary widely (1–6 A/m) (Fig. F7G); this requirement raises questions about whether the anomalies are true seafloor-spreading lineations (Whitmarsh et al., 2001; Russell and Whitmarsh, 2003). For this reason, we reexamine the source of the magnetic anomalies in the next section. There we conclude that the anomalies are caused largely by serpentinization of exhumed mantle rocks and not by magnetization of igneous rocks emplaced during seafloor spreading.

Source of Magnetic Anomalies

The Role of Magmatism

As discussed previously, the composition of peridotites recovered at drill sites in zones TE1 and TE2 indicates that the mantle was strongly to moderately refractory and thus was unlikely to have produced significant volumes of melt as mantle was exhumed. Nonetheless, the drilled peridotites also show significant heterogeneity, and it is possible that relatively fertile mantle is present in the rift but remains unsampled. If we consider decompression melting of normal mantle, gabbroic and basaltic crust ~2 km thick might be produced (Bown and White, 1994), and we thus could expect substantial magmatic products to be present somewhere within the TE zones. For several reasons, however, it seems highly unlikely that these igneous rocks are present at shallow levels in the basement. First, it is clear from drilling results that there is a dearth of volcanism and magmatic intrusions in the uppermost basement, and this makes it doubtful that there are large underlying magmatic bodies at shallow levels. Second, seismic refraction results, summarized earlier, show little evidence for the presence of igneous rocks in the upper basement except in limited (younger) parts of TE2. Finally, persistent shallow magmatic intrusions should produce a "layer" with relatively uniform magnetization, counter to observations that magnetization intensity appears to be highly variable (Fig. F7).

An alternate possibility is that magnetic anomalies are caused by magnetic bodies located deep in the basement (~5–8 km), as has been proposed for the southern Iberia Abyssal Plain (Russell and Whitmarsh, 2003). To test this idea, Sibuet et al. (2007) conducted a combined inversion for the structural index and the source location using Euler deconvolution, which only uses derivatives of magnetic anomalies (Hsu, 2002). The primary result of this analysis is that the main magnetic sources are at depths <3 km within the basement and structural indexes are ~2, which means that the magnetic sources approximate cylinders that are probably horizontal and parallel to the magnetic lineations. Considering the above arguments that there probably is not a significant, shallow igneous source layer generating the magnetic anomalies, we must to look for other ways to account for the magnetization. As discussed below, magnetization of serpentinites in the upper basement provides a likely explanation.

The Role of Serpentinization

The contribution of peridotites to magnetic anomalies becomes significant when the peridotites are strongly serpentinized (>75%) (Oufi et al., 2002). This degree of serpentinization may be widespread in the upper 1–2 km of exhumed mantle in the zones of transitional extension, where observed low seismic velocities would correlate with >75% serpentinization (e.g., Christensen, 2004; Chian et al., 1999; Lau et al., 2006b).

Direct evidence that serpentinites can contribute significantly to the magnetic field has been presented by Zhao et al. (Zhao, 1996, 2001; Zhao et al., this volume, 2006), who studied heating and cooling curves for serpentinized peridotite samples from both the Newfoundland and Iberia margins. In general, they found that yellow and brown peridotites have Curie temperatures near 420C; this indicates the presence of maghemite and appears to account for low natural remanent magnetization (~1 A/m). Curie temperatures of gray and green peridotite samples are ~570C, indicating presence of magnetite and accounting for high remanent magnetization (up to ~9 A/m). These values of remanent magnetization are comparable to or higher than those of oceanic basalts. All samples show inclinations close to those predicted for the paleolatitude of the drill sites, indicating that natural remanent magnetization (NRM) intensities probably represent primary magnetization.

Oxygen isotope profiles of serpentinized peridotites at Sites 1068 and 1070 show evidence for two phases of serpentinization (Skelton and Valley, 2000), and these may explain the differences between the gray-green and yellow-brown serpentinites. The first phase (>175C) involved pervasive infiltration of water and bulk serpentinization that is interpreted to have occurred before the mantle was exhumed at the seafloor. The second phase (<50–150C) occurred at or close to the seafloor. Consequently, strong magnetization was first acquired at depth during an initial phase of serpentinization when gray-green serpentinites recorded the polarity of the ambient magnetic field. The second phase affected only the uppermost basement at the seafloor, giving rise to maghemitized yellow-brown serpentinites. Paleomagnetic results show that the yellow-brown "altered" serpentinites acquired their final magnetization later than the gray-green "fresh" serpentinites. Zhao (2001) suggested that the magnetization in the yellow-brown serpentinites zone was an overprint imposed during the Cretaceous Long Normal Superchron.

Serpentinites from Holes 897D, 899B, and 1070A on the Iberian margin yielded mean NRM intensities of 0.35, 1.8, and 1.6 A/m and mean magnetic susceptibility values of 2.1 x 10–2, 2.9 x 10–2, and 2.9 x 10–2 SI units, respectively (Zhao, 2001), with variations in magnetic susceptibility mimicking those of the NRM intensity. The higher magnetization intensity in Hole 899B is consistent with the presence of the large magnetic anomaly recorded on the SAR-93 deep-tow profile (Fig. F7). Furthermore, several samples from Hole 899B retained nearly half their NRM intensity after 400C demagnetization, suggesting that remanent magnetization in magnetites can contribute significantly to the strong 800-nT magnetic anomaly (Whitmarsh et al., 1996a). At all three sites a 40- to 75-m-thick yellow-brown "altered" zone with predominantly normal polarity and probably with maghemite overlies a gray-green "fresh" zone where magnetite is likely present (Zhao, 2001).

In yellow-brown and gray-green serpentinized peridotites at Site 1277 on the Newfoundland margin, measured Curie temperatures of 420C and 550–580C indicate the presence of maghemite and magnetite, respectively. High NRM intensities (up to 9 A/m) in the gray-green serpentinized peridotites may help to explain the amplitude of observed sea-surface magnetic anomalies (Zhao et al., this volume).

It is important to note that magnetic properties of extensively serpentinized peridotites can vary significantly depending on whether the peridotites have been maghemitized (either at the seafloor or in a permeable fault zone), whether the conditions of serpentinization allowed for the crystallization of iron-bearing silicates together with serpentine and magnetite, and whether conditions favored formation of a serpentine meshwork with thin veinlike concentrations or small and/or highly elongated magnetite grains (Oufi et al., 2002). Consequently, there is no reason to expect laterally uniform magnetization in serpentinized basement. From this perspective, the highly variable magnetization values used in the magnetic model of Figure F7 are not unreasonable.

Based on the above discussion, we suggest that serpentinization may explain the magnetization responsible for the magnetic anomalies in the TE zones, exclusive of the interval around anomalies M1–M0 where plume magmatism from the Southeast Newfoundland Ridge may have had an effect. In this model, fluids penetrated deeply along margin-parallel faults and fractures caused by plate bending of the exhuming mantle, which allowed penetrative serpentinization to propagate as deep as 2–3 km. Pseudo-single domain magnetite grains in gray-green serpentinites oriented along the magnetic field below the Curie temperature (~570C), recording the ambient magnetic field as the mantle cooled at crustal levels (Sibuet et al., 2007). Strong spatial variability in serpentinization is likely to have occurred, and this would have created vertically and horizontally heterogeneous crustal magnetization, consistent with forward modeling where the mean magnetization varies from 0.75 to 6 A/m (Fig. F7). As the serpentinized mantle was exhumed at the seafloor, the upper several tens of meters were altered by further reaction with seawater (Fisher-Trop reaction); this increased the degree of serpentinization, and magnetite was replaced by maghemite.

According to this process, seafloor age indicated by a magnetic anomaly would be the same as the age at which the serpentinized peridotite became exposed at the seafloor (i.e., the same relative timing as for anomalies generated from magnetization of igneous bodies), or it would somewhat predate the time of exposure (taking into account the time required for the magnetized serpentinite to be exhumed from some depth). At Site 1070, the ~128-Ma magnetic anomaly age is the same as the U/Pb age of 127 the mantle peridotite (Table T1) (Beard et al., 2002). This is also true at Site 1277, where the magnetic anomaly age (~127 Ma) is the same as a 128 [N1]). The coincidence of magnetic and intrusion ages suggests that the intrusions were emplaced when the serpentinized mantle first acquired its magnetization. However, the error bars on the intrusion ages allow for the possibility that the magnetic ages are somewhat older (up to ~3–4 m.y.), as noted above. Zhao et al. (this volume) determined that the magnetic inclinations in the serpentinites at Site 1277 imply ~35 of counterclockwise tectonic rotation (viewed facing toward N010E) after the magnetization was acquired. Considering all of the above, the following evolution of the basement ridge where Site 1277 was drilled seems likely:

  1. Mantle peridotite was first serpentinized and thereby acquired its magnetization at subseafloor depths of up to ~2–3 km; this was accomplished as fluids penetrated along the master normal fault (and perhaps subsidiary faults) that uplifted the ridge.
  2. Subsequent exhumation uplifted the ridge more than a kilometer above the surrounding seafloor. If the exhumation was along a normal fault on the east side of the ridge (as required to account for counterclockwise rotation) the basement morphology (Fig. F3) suggests minimum fault heave of ~5 km; if we add exhumation from subseafloor depths of ~2–3 km, the total horizontal extension probably was at least 7–8 km.
  3. Minor magmatism, possibly enhanced by decompression melting, occurred during the exhumation.
  4. Upon exposure at the seafloor, a thin "skin" of the exhumed mantle was further altered and maghemitized by low-temperature serpentinization.

If magnetized serpentinites account for the observed magnetic anomalies, it is conceivable that basement topography alone generates the anomalies. We investigated this possibility with the forward model shown in Figure F7I. We assumed a magnetized basement ~3 km thick with a flat bottom and a constant magnetization of 5 A/m, a value close to or somewhat larger than the mean value measured in serpentinized peridotite samples. Model results were computed for the sea surface and sea bottom (Fig. F7H). The results show that the contribution of basement topography is only ~10% of the amplitude of observed magnetic anomalies. A similar result was obtained by Russell and Whitmarsh (2003). We conclude that basement topography alone cannot explain the amplitude of magnetic anomalies.

Implied Extension Rates

If we assume that the magnetic anomaly picks by Srivastava et al. (2000) are correct (but taking note that anomaly picks older than M8 are highly debatable), we can calculate extension/spreading half rates as function of time as shown in Figure F8. The upper part of the figure shows three different timescales used in the calculations: (1) the Kent and Gradstein (1986) timescale (KG86), which was used by Srivastava et al. (2000); (2) the Gradstein et al. (1995) and Channel et al. (1995) timescales (GC95); and (3) the most recent Gradstein et al. (2004) timescale (G04). For comparison, we also show the Fiet et al. (2006) timescale, which is based on ages of glauconites sampled in the Vocontian Basin (southeast France) together with cyclostratigraphy previously obtained in southeast France and basins in central and southern Italy. The large differences between these timescales significantly affect calculated extension rates as shown at the bottom of the figure.

Extension half-rates calculated for Chron M8 (Hauterivian) and older are generally consistent at 6–7 mm/yr (Fig. F8). A similar rate (7.2 mm/yr) for the younger part of this period was derived by Whitmarsh et al. (2001) from geometrical reconstruction of faults (dotted line labeled W in Fig. F8), but they derived a lower rate (>3.5 mm/yr) for the older part. The major inconsistency in extension rate is for the G04 timescale, which shows a large excursion at M11 that probably is not real. Whether this reflects a problem with the timescale or with the anomaly identifications is unknown. In the Barremian after M8 (i.e., in the early part of TE2), rates increase to ~9–13 mm/yr, depending on the timescale used. Then from M0 to C34, rates are ~13 mm/yr or increase slowly to this value. Note, however, that there are no constraints within the long Cretaceous Magnetic Quiet Zone between M0 and C34, and there may be unknown rate variations in that interval.

The 6- to 7-mm/yr rates characterize TE1 and they are ultraslow, comparable to the slowest rates on Earth (i.e., Gakkel Ridge in the Arctic Ocean [Cochran et al., 2003]), and the ~9- to 13-mm/yr rates in TE2 are comparable to those at the slow end of slow-spreading mid-ocean ridges. It is natural to look for lithologic and morphologic analogs at mid-ocean ridges that spread at these rates. For example, on the Southwest Indian Ridge where the spreading half-rate is ~7 mm/yr, ~40% of the crust has a smooth or corrugated surface that may be formed by long-lived normal faults (Cannat et al., 2006) and it is composed of peridotites 60% corresponds to apparently more volcanic and normally faulted seafloor that is more like that of slow-spreading ridges. However, it must be kept in mind that the critical control on morphology and composition of newly emplaced seafloor is not extension rate but melt supply. This is dramatically demonstrated, for example, in the Australia-Antarctic Discordance where the half-spreading rate is 75 mm/yr (i.e., intermediate rate), yet because of very low melt supply the morphology is like that of slow- to ultraslow-spreading ridges (Okino et al., 2004). Furthermore, for some uncertain distance seaward into the TE zones (but apparently seaward past anomaly M1, as discussed earlier), subcontinental mantle lithosphere was being exhumed and its properties may have influenced morphology and mantle melting patterns in much different ways than asthenospheric mantle beneath mid-ocean ridges. Thus, there are good reasons to be cautious about making comparisons solely on the basis of extension/spreading rate.