Temperature anomalies under the Northeast Atlantic rifted volcanic margins

EPSL ELSEVIER

Earth and Planetary

Science Letters 146 ( 1997) 195-211

Temperature anomalies under the Northeast Atlantic rifted volcanic margins Peter D. Clift Department

of Geology and Geophysics,

*

Woods Hole Oceanographic Institution, Woods Hole. MA 02543, USA

Received 21 November 1995; revised 3 October 1996: accepted 21 October 1996

Abstract Subsidence analysis of ODP/DSDP drill sites located on oceanic crust on the Southeast Greenland, Edoras Bank, and Vorin Plateau margins, as well as on the Iceland-Faeroe Ridge, shows that the subsidence of these areas does not follow the P-age relationship of normal oceanic crust. By correcting for the effect of thickened oceanic crust in raising the level to which subsidence will occur and analyzing the rate of thermal subsidence, it is possible to provide maximum temperature estimates for the underlying asthenosphere through time by identifying periods of anomalous depth to basement. Isostatic models predict crustal thicknesses of 27 km under the Iceland-Faeroe Ridge, around 20 km at Edoras Bank and Southeast Greenland, and 16- 17 km at the Voring Plateau. Asthenospheric temperatures at the time of continental break-up range from 50°C to 100°C above normal mantle, which are insufficient to account for the crustal thicknesses if melting is purely a passive adiabatic process. Asthenospheric upwelling must thus have been more rapid than spreading following break-up. At

Edoras Bank the thermal anomaly dissipated within 5 Myr of rifting, similar to that inferred from the eastern US margin, where no plume is considered to have affected the rifting process. The need to invoke thermal input from the Iceland Plume in generating the thickened crust at Edoras Bank, and possibly elsewhere in the Northeast Atlantic, is called into question. However, a 14-20 Myr anomaly, peaking at 12 Myr post-rift, in Southeast Greenland suggests that, although the plume did provide heat to this margin, its strongest influence post-dated break-up. Keywords: Northeast

Atlantic;

Paleocene;

continental

margin; ocean basins; subsidence

1. Introduction Current classification of passive rifted margins recognizes two distinct groups: volcanic and nonvolcanic. Yet the mechanisms behind the formation of one rather than another type of margin remain relatively poorly understood. Volcanic margins differ from their non-volcanic counterparts in having sharp

* Fax: + 1 508 457 2187. E-mail: [email protected] 0012-821X/97/$12.00 Copyright PII SOOl2-821X(96)00228-2

continent-ocean transitions and thick tholeiitic basalt sequences erupted at the time of break-up. The lavas, which typically dip toward the ocean basin [l], are characteristic features of these regions, although the driving force behind their generation remains obscure. In particular, the relationship between volcanic margin formation and the presence of upwelling asthenospheric plumes is a key issue in the understanding of these systems. The Northeast Atlantic Ocean represents the classic example of rifted volcanic margins generated in the presence of a

0 1997 Elsevier Science B.V. All rights reserved.

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mantle plume [2] and is thus a good place to test the link between asthenospheric temperature anomalies and excess volcanism at the time of break-up. In simplified form, the excess volcanism associated with continental break-up has been attributed to greater degrees of melting in the rift axis caused by anomalously high temperatures in the underlying asthenosphere, associated with a mantle plume [3]. In the Northeast Atlantic, White and McKenzie [3] invoked the influence of a large diameter proto-Icelandic Plume to explain all the Paleocene volcanism from Disko Island in western Greenland to western Scotland (Fig. l), which they related to simple, adiabatic upwelling and melting of hotter than normal asthenosphere. Most recently, White et al. [4] used a rare earth element inversion technique to estimate a mantle potential temperature of around 1500°C (i.e., a +22O”C anomaly) for Krafla in the Iceland Neovolcanic zone, demonstrating the feasibility of a simple adiabatic melting model for the generation of the thickened oceanic crust of the Faeroe-IcelandGreenland Ridge [51.

1, I. Objectives and methodology

In this study the record of vertical motions since rifting is used to draw conclusions about the variations in temperature under the rifted margins through time, and test the model of adiabatic upwelling as a petrogenetic mechanism for the seaward-dipping lavas of the Northeast Atlantic. The vertical motions in any one place are recorded in the sedimentary column overlying the volcanic basement and are manifest as shallowing or deepening of sedimentary facies, changes in the benthic foraminifer assemblages, as well as unconformities and paleosols. Special interest is focused on the time of rifting and the immediate post-rift period in an attempt to quantify the thermal state of the asthenosphere under the Northeast Atlantic at that time. The sites chosen all lie on oceanic crust close to the continent-ocean transition (COT). Oceanic crust is chosen for study in preference to sites located on extended continental crust because the subsidence history is simpler to model and there is no risk of complications related to

60’

50”N 6

Fig. 1. Present day bathymetric map of the Northeast Atlantic, showing the location of the ODP/DSDP as existing deep seismic cross sections across the COT. Bathymetric contours are in 500 m intervals.

drill sites used in this study, as well

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early rifting events. Although mechanical coupling across the COT is a possible source of error, it is a reasonable approximation where the lithosphere is weak or the applied load is wide and evenly distributed. Seismic reflection studies [6] show that the sediment load at these margins was approximately evenly distributed over a wide area during the early post-rift period. In addition, Karner and Watts [7] showed that the rifted continental crust of the Coral Sea and Lord Howe Rise still only has an elastic thickness of 5 km, although rifting was complete at 60 Ma. Although the rifted margins of the North American continent do show much greater elastic thicknesses, approximately 15 km, these are much older features that any of the margins in the Northeast Atlantic. Replacement of normal asthenosphere by less dense material causes most of the temporary plumerelated uplift, or dynamic support. By choosing those sites on the oldest crust, depth anomalies related to temperature anomalies under the newly initiated spreading axis may be recorded. Temperature estimates derived from the vertical motions at each studied volcanic margin transect are compared with those inferred from the thickness of the oceanic crust measured by direct geophysical surveys and through isostatic calculations. I aim to demonstrate whether there were temperature anomalies under the volcanic margins at the time of rifting and, if so, whether the excess asthenospheric temperatures alone are capable of generating the volcanism seen. Sites 918 (east Greenland), 343, 643 (Voring Plateau) and 336 (Iceland-Faeroe Ridge) were chosen for analysis (Fig. 1). Site 343 on the Northern Voring Plateau was used because it is located away from the Jan Mayen Fracture Zone that may have affected the thermal evolution at Site 643. Site 343 provides a second, and possibly more representative, picture of the thermal evolution of the VGring Plateau. Site 336, located on the Iceland-Faeroe Ridge was also included because it has a relatively continuous sedimentary record. The basement at this site was dated at 44 k 3 Ma [8] using K/Ar methods; however, due to the unreliability of this approach a basement age of 52 Ma was used, taken from the continuation of magnetic anomaly 24 on to the Iceland-Faeroe Ridge from the adjacent oceanic crust to the south [9].

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In addition, Site 553 on the Edoras Bank (southern Hatton Bank) is considered. This site is located close to the COT, but on the basis of Pb isotopic studies of the igneous basement it appears to overlie strongly rifted continental crust [lo]. Nonetheless, several lines of evidence suggest that the degree of extension is so high here that the thermal subsidence of the site is effectively of oceanic character. The MORB-like major element and trace element chemistry of the basalts [lo] shows that, in this place, the amount of continental contamination of the melts is minor and that melting of the upwelling asthenosphere is occurring as though the overlying lithospheric lid was no greater than in an oceanic setting (i.e., that extension was extremely high). Sr and Nd isotopes of the same lavas do not show a noticeable departure from normal Atlantic MORB [l 11. Furthermore, attempts to fit continental thermal subsidence curves to the reconstructed data show major discrepancies. Barton and White [ 121 suggested a p of 3 at Site 553 on the basis of refraction seismic experiments. Comparison of the rifting model of McKenzie [ 131 with p = 3 and the reconstructed tectonic subsidence curves implies a brief period of temporary, relative uplift at the time of break-up, presumably explicable in terms of hotter asthenospheric temperatures under the margin at this time. However, the model does not provide a good fit between 5 1 and 45 Ma, when the basement is temporarily up to 400 m deeper than would be expected, a feature that is difficult to explain in terms of asthenospheric temperatures. In contrast, the backstripped subsidence curve for Site 553 fits the Parsons and Sclater [14] oceanic subsidence curve well for the bulk of the Cenozoic, except for the first 7 Myr after break-up, when hotter than normal temperatures are predicted

[151. 1.2. Models for volcanic margin formation Theories on the origin of rifted volcanic margins divide between those invoking the influence of a mantle plume [3] and those that involve different mechanisms to generate the excess melting observed. Mutter et al. [16] invoked the influence of enhanced convective overturn of the asthenosphere underlying the Voring margin as a mechanism to drive greater degrees of melting without having asthenospheric

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temperatures elevated above normal mantle (i.e., about 1280°C). The sharp COT in this setting is considered to accentuate the induced convection, so explaining the lack of voluminous volcanism in non-volcanic margins where the transition from continent to ocean is more gradual. While the existence of induced convection is recognized as a cause of rift flank uplift in some young ocean basins [17], it is unknown to what extent it may have a role in causing the excess volcanism of volcanic passive margins. Other non-plume theories do invoke increased mantle temperatures prior to rifting. Anderson [18] suggested that the insulating effect of continental lithosphere can act to store heat under supercontinents, resulting in excess melting and volcanism due to adiabatic upwelling when these masses eventually rift to form new ocean basins. In a similar way Kelemen and Holbrook [19] suggested that the volcanism seen during the initial stages of oceanic spreading along the eastern US margin is at least partially caused by the presence of anomalously high temperatures under the continent prior to rifting, although no mechanism was advanced to explain the presence of elevated temperatures in the first place. The closest plume to this province at the time of rifting, the Great Meteor hotspot, was considerably removed from the rift [20] and cannot have made a thermal contribution to the rift volcanism. Higher than normal asthenospheric temperatures under the eastern US margin would have lowered the viscosity of the asthenosphere and allowed rapid upwelling and melting under the rift axis, producing a thicker than normal oceanic crust. The low degree partial melts inferred from the high seismic velocities seen in the lower crust off the shore of the eastern US indicate that passive adiabatic upwelling alone cannot account for the igneous thicknesses observed. The 80- 100 km width of the thickened crust in this region suggests that the thermal anomaly and rapid upwelling dissipated within 5-8 Myr of the initiation of spreading, given the average spreading rate during the Jurassic in the central Atlantic [ 191. Extraction of greater degrees of melt from the asthenosphere than would be predicted for passive, adiabatic upwelling was also proposed by Ribe et al. [21], whose modeling of the modem Iceland Plume showed that a broad, and relatively cool plume (= 90°C above normal - i.e., 1280°C) could best ex-

Science Letters

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plain the bathymetric and gravity data associated with the Iceland hotspot. In this case enhanced convection would be required to produce the 30 km thick crust seismically imaged under Iceland [5]. Using the 220°C temperature anomaly of White et al. [4] produced a geologically unrealistic melt thickness of 68 km using the Ribe et al. [21] model. Geochemical data from the earliest lavas on the East Greenland margin (i.e., the basement at ODP Site 918) and Edoras Bank (Sites 552 and 553; Fitton et al. [22]) now suggest that simple adiabatic melting cannot have been the principle petrogenetic mechanism during initial continental break-up. Fitton et al. [22] demonstrated that the basalts of the main dipping reflector sequences at these localities have high SC (50-65 ppm), implying melting at very shallow levels within the spine1 lherzolite mantle field. Modeling by Langmuir et al. [23] showed that, for the large melt volumes found at spreading ridges, a high SC content would require very shallow, large-degree melting, consistent with active upwelling below the ridge crest.

2. The causes of anomalous vertical motions The subsidence history of normal thickness oceanic crust has been well defined and modeled as resulting from cooling and thickening of the oceanic lithosphere away from the spreading axis [14]. The depth of newly generated crust is typically 2500 m and subsides thereafter as a function of t/age, although it shows a decline in the rate of subsidence after 70 Myr [24]. Departures from this model can be caused by a number of mechanisms. Flexure may be important in causing uplift or subsidence relative to the expected depth for crust of any given age. Early after its generation the strength of oceanic lithosphere is low 171, but it rapidly becomes stronger and able to bear stresses. Larsen and Marcussen E251 suggested that flexural deformation due to ridge push along the Reykjanes Ridge may be the driving mechanism for uplift of the East Greenland coast during the Oligocene, well after the initial break-up. The fact that some ridge push is going on has been shown by downhole logging experiments on the East Greenland shelf [6]. which recorded borehole break-

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Earth and Planetary

outs in accord with compression perpendicular to the margin. Nonetheless, gravity studies of the Northeast Atlantic suggest that, even today, when the flexural rigidity is at a maximum since rifting, this region is basically in local isostatic equilibrium and that the amount of uplift resulting from ridge push is small

D61. Transient uplift of the basement relative to the depth predicted by Stein and Stein [24] at any one time can be caused by the upwelling of asthenosphere under a plate. Indeed, regional domal uplift is a characteristic feature of the presence of mantle plumes [27]. The change from vertical upwelling to horizontal flow along the base of the lithosphere induces an upward force on the plate, causing a dome-shaped uplift around the center of upwelling. The width of this uplift will depend on the rate of upwelling, the width of the upwelling column and the flexural rigidity of the overlying plate. Some sort of coupling is also required to transfer stresses from asthenosphere to lithosphere. Exactly how effective this force is at generating uplift is questionable, since it has been demonstrated that there is a thin low viscosity layer at the base of the lithosphere, which would reduce the effect to negligible values [28]. Temperature variations in the asthenosphere have been considered the most important method of driving uplift in the ocean basins and rift zones. especially during the syn-rift period when flexural rigidity is a minimum [3,7]. The replacement of normal asthenosphere by hotter than normal material above the level of isostatic compensation will cause uplift, due to the decrease in density of the mantle. Further uplift may be caused by a heating and thinning of the overlying lithosphere, effectively thermally rejuvenating the plate. Evidence from the subsidence of the Hawaiian Seamount Chain [29] suggests that this process is active under the fast-moving Pacific plate. However, geophysical studies of hotspots positioned over slow-moving plates [30] indicate that this is much less important in these settings. This prediction was born out by recent subsidence analysis of the Northeast Atlantic margins [15]. Extraction of melt from the asthenosphere may cause uplift due to a decrease in density of the remaining refractory asthenosphere. Recent modeling of the Hawaiian Swell suggests that much of this feature may be due to the extraction of melt in the upwelling column of as-

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thenosphere under the present hotspot [3 l]. While having a smaller effect than the simple thermal expansion, this mechanism is thought capable of causing uplift of up to 500 m [3] above a spreading axis. However, at sites located far from a spreading axis or hotspot the mixing of melt-depleted mantle with undepleted material reduces the ability of this mechanism to cause major uplift.

3. Thermal modeling In this study time periods in which the depth to basement at each drill site is shallower than expected from the Stein and Stein [24] model are identified and used to derive estimates of asthenospheric temperature. Comparison with the Stein and Stein [24] curve must be treated cautiously because oceanic lithosphere cooling above a hotter than normal asthenosphere will tend to subside more rapidly than normal. In order to infer paleo-asthenospheric temperatures it is first necessary to estimate the temperatures under the drill sites today, since this represents the background against which anomalies will be detected. The calculated depth anomalies at each site are measured relative to the Stein and Stein subsidence curve normalized to the present corrected depth to basement. If this depth is elevated above what might be expected for crust above normal-temperature mantle then this difference needs to be accounted for. 3.1. Modern temperature

anomalies

The thermal model of Ito and Lin [32] for mantle plumes is used in this study, in which a value for potential asthenospheric temperatures at different distances from the plume center was calculated using data from mid-ocean ridges (Fig. 2). Ito and Lin [32] note that the present day thermal anomalies in Iceland (which they estimate as 150°C) lie above the averaged curve for all the hotspots considered. This suggests that Iceland may be a hotter plume than other examples. Therefore, for this specific case the Ito and Lin 1321 temperature-distance curve was recalculated to fit the Icelandic data more closely. A maximum 150°C temperature anomaly was chosen rather than the 90°C anomaly of Ribe et al. [21]

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because this will tend to maximize the paleo-mantle temperatures and provide the strongest test of the adiabatic melting model. The 22O’C anomaly of White et al. [4] is rejected as being incompatible with the observed bathymeny and melt thicknesses, as noted by Ribe et al. [21]. Although White et al. [4] estimate modem asthenospheric temperatures under the Reykjanes Ridge, these cannot be used for the margin drill sites, even if the 220” central anomaly was correct, because the estimates refer to the asthenosphere under the ridge, which might be expected to channel hot plume asthenosphere away from Iceland, because of the buoyancy of the plume material [27]. This channelling would not affect the areas far off-axis considered in this work. All the drill sites on the volcanic margins in question are presently located far from the plume center (about 900 km), but within the range of the thermal anomaly, which is seen to continue almost 1000 km south of Iceland on the present Reykjanes Ridge [33]. While the modified temperature-distance curve may not be completely accurate, the flat gradient to the curve at distances greater than 400 km will mean that errors in the modem temperature estimates will be small. All plume models tend to agree that the radius of influence of the Iceland Plume does not exceed 1000 km. A modem temperature anomaly of around 56°C appears appropriate for the asthenosphere under the rifted volcanic margins along the drilled transects. The temperature under Site 336 on the Iceland-Faeroe Ridge is estimated at around 25°C above normal mantle. 3.2. Modeling oceanic subsidence In order to gauge the magnitude of the thermal anomalies in the asthenosphere underlying the vol-

from [32].

canic margin at and since break-up the reconstructed subsidence curves were compared with the subsidence pattern for normal oceanic crust of Stein and Stein [24], and also with two model curves reflecting the subsidence of oceanic crust under which a plume head of anomalous temperature was emplaced at the ridge at the time of spreading and then allowed to cool diffusively through time. Subsidence immediately following spreading is sensitive to the mode of heat loss. To minimize the number of parameters and dynamic effects involved in the calculations a purely diffusive end-member model was employed in which lithospheric cooling starts with the injection of an anomalously hot mantle layer at the time of spreading. The one-dimensional heat equation is solved: aT( t,z) at

=K

a2T( t,z) a22

where K is thermal diffusivity (25.6 km/Myr2); z is depth below the seafloor (in km); (Y is the coefficient of thermal expansion (3.1 X lo-‘/“C); and T is temperature (“C). Temperature is set at 0°C at the surface and increases to 1450°C at the base of a 100 km thick lithosphere in the final steady-state condition. This is similar to the 92 km thick plate modeled by Stein and Stein [24] and results in a subsidence curve effectively identical to theirs in the case where no thermally anomalous layer is present. The model introduces a layer with anomalous temperature and a predetermined thickness (100 or 200 km) beneath the 800°C isotherm at the ridge and then allows the temperatures to diffuse to the final steady-state conductive geotherm in the 100 km thick plate. Fig. 3 shows the model parameters for the case of a 200 km thick, +2OO”C anomalous layer. At the time of spreading the hot mantle material extends up to the

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Fig. 3. Thermal model used in the subsidence analysis, showing the emplacement of 100 km and 200 km thick layers of anomalous asthenosphere under a spreading ridge.

seafloor, gradually decaying by conduction to a 100 km thick lithosphere underlain by normal asthenosphere. The same model successfully predicts the initial depth anomaly for the Hawaiian hotspot, giving a 1.5 km depth anomaly for a 100 km thick plume head with a temperature anomaly of 250°C. The reconstructed curves are compared with a model curve generated for a 100 km thick hot layer, similar to that predicted for the modem, steady-state Hawaiian plume [34], and for the Cape Verde hotspot [35]. A 200 km thick anomalous layer is also modeled, as this may be more appropriate if the temperature anomaly is not plume-related (e.g., building up because of the insulating effects of the continental plate [ 1811, or if the plume is initiating at the time of break-up, when the hot plume head might be expected to be especially deep [36]. 3.3. Reconstructed

subsidence

histories

I use the calculated subsidence histories of Clift et al. [15] for ODP/DSDP Sites 553, 343, 643, 918, and 336 (Fig. 1). The subsidence histories were derived from examination of the sedimentary and microfaunal evidence published in the DSDP and ODP Initial Reports [6,37-391, and the published range of paleo-water depths was used where this was estimated by each scientific party. The paleo-bathymetric scheme of the shipboard party [40] was used for converting foraminifer assemblages into numerical water depths. At Edoras Bank only a minimum water depth was specified by the shipboard party, with the maximum left undefined. In this situation a maximum value was chosen so as not to cause any discrepancy with the predicted depth to basement at that time, assuming a normal Stein and Stein [24]

thermal subsidence. These points are marked on the subsidence diagrams. Subsidence was assumed to follow the form of normally cooling oceanic lithosphere, unless there was evidence to the contrary. By doing this Clift et al. [15] avoided creating apparent anomalies in the subsidence history that are not otherwise supported by good paleo-water depth data. In some deep-water sediments the range of possible water depths is high, and in these cases a detailed subsidence reconstruction is not possible. A full discussion of the paleo-water depth information for the sites discussed here is included in Clift et al. 1151. 3.4. Uncertainties

in the subsidence

analysis

A basement subsidence history for each site was derived using the backstripping subsidence method of Sclater and Christie [41]. This approach uses lithology and age information taken from the cored material to derive a depth to basement, and removes the loading effects of sediment and water so that the tectonically driven subsidence of the basement can be isolated. In each case two possible tectonic subsidence curves are plotted, which represent minimum and maximum estimates of the water depth. The true loading-corrected subsidence pattern of the basement must lie between these two estimates. This method assumes an empirical porosity-depth curve that is based on lithology rather than any actual data derived from the borehole by logging or measurement of recovered material. When making the unloading correction the density of the mantle is assumed to be 3330 kg/m3 1421. Cooling of the lithosphere with time after spreading results in increased density of the mantle, which is not corrected for in the backstripping process. However, this change is very small over the temperature range considered here due to the low thermal expansion coefficient 1431, resulting in uncertainties of < 1% for cooling of > 300°C. This not significant for the thicknesses of sediment considered in this study. When dealing with oceanic crust, and thus significant water depths, estimates of paleo-water depth are crucial to a meaningful result and represent the single largest uncertainty. Variations in the degree of sediment compaction are of a magnitude smaller than any possible errors in the water depth, especially in the Northeast Atlantic where the sedimen-

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tary cover is relatively modest. No attempt has been made to correct for fluctuations in eustatic sea level, as current predictions of rates and magnitudes of subsidence remain controversial [44] and produce unlikely. saw tooth-like subsidence curves when taken into account. The long-term sea variations of Haq et al. [44] show sea level in the Early Tertiary being about 150 m higher than today. A correction for this difference results in a steepening of the reconstructed curves and thus an increase in the apparent depth and temperature anomalies shortly after break-up. Since the initial depth anomaly is calculated at 300-500 m [ 151, a 1.50 m sea-level correction would represent a significant change, although this would be less important if sea-level variations based on oxygen isotope work are used because these show fluctuations of 30-50 m above modem levels in the Early Tertiary [45]. Such variations would result in total uncertainties of around 10%. Comparison of porosity taken from shipboard physical property measurements shows a generally good agreement between the model of &later and Christie 1411, especially for sandy lithologies. At Site 643 physical property measurements suggest higher than expected porosities for muddy sediment at > 500 mbsf, near the base of the sediment column [38]. However, even in this case the measured porosities of around 60% are not greatly different from the modeled value of 55%. Errors in the age data are typically small, and the duration of nannofossil and foraminiferal zones are short (l-2 Myr) in the Eocene. At Site 918 the ages at the base of the column are derived from both foraminifer and nannofossil sources, which show good agreement [6]. The ages at Site 553 are determined solely by nannofossils at the base of the section. and by a mixed assemblage of foraminifers and nannofossils in the late Eocene section. No discrepancies are noted between the two schemes. At Site 336 on the IcelandFaeroes Ridge the oldest sediments are dated entirely

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based on planktonic foraminifers [39]. Dating at Site 643 differs from the other sites in that the oldest rocks are dated by palynology and through benthic foraminifers [38]. The latter provide poor age control but the former allow even more precise dating than using other methods [46]. At none of the sites do different dating schemes suggest significantly different ages of sedimentation. 3.5. Correcting

,for thick oceanic crust

Anomalous subsidence behavior is identified by comparison of the reconstructed subsidence curves with the modeled curves. A direct comparison between the Northeast Atlantic drill sites and regular oceanic crust will not reveal temperature anomalies because the thickened crust in this area would be expected to subside to a shallower than normal depth when in isostatic equilibrium, even without the presence of anomalously hot asthenosphere. Nonetheless, for a given asthenospheric temperature, it might be expected that the depth to basement would increase at the modeled rate because of the cooling and thickening of the underlying lithosphere. which is not governed by crustal thickness. In the case of Sites 553, 918, 343, and 643 the modem asthenospheric temperature is close to normal (5°C anomaly) meaning that the basement might be expected to have a subsidence history close in form to the Stein and Stein [24] curve, vertically shifted to match the modem corrected depth to basement. assuming that the underlying asthenosphere had remained at this temperature since break-up. Fig. 4 shows the reconstructed subsidence curves compared with the model curves for both 100 km and 200 km thick plume heads. It can be seen that at all margin drill sites both the minimum and maximum water depth curves lie above the normal subsidence curve immediately after spreading, but below the + 100°C temperature anomaly curve, except at Site 553 modeled with a 100 km thick hot layer, where

Fig. 4. Calculated subsidence curves for Sites 553, 918, 343, 643 and 336. Superimposed theoretical curves are from 1241 and with the . . addition of a 100°C and 200°C anomalous layer. 100 km thick in (A)-(E), and 200 km thick in (F)-(J). The solid line with dots represents the tectonic subsidence of the basement with sediment and minimum water depth correction: the solid line with circles represents the maximum water depth correction. White squares in the maximum water depth for Site 553 show values assigned to show no depth anomaly [ 151.Vertical dotted lines represent epoch boundaries.

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the asthenospheric temperature could rise as high as + 140°C above normal, and at Site 343 where both 100 km and 200 km thick plume models give higher temperatures (160-280°C and 80- 160°C respectively). Although the number of dated points which depart from the model curve are not numerous in each case, the tightly defined nature of their water depth estimates means that a high degree of confidence can be placed in their interpretation. The small number of actual dots merely reflects the modest number of biostratigraphic tie points available, between which the section is uniformly a shallow water marine sediment. At Site 553 the reconstructed curve appears to follow an elevated temperature subsidence path for 7 Myr before showing a rapid deepening and subsequently no further evidence for a departure from the normal curve. Such an evolution suggests that the hot asthenosphere underlying the margin at the time of break-up is not allowed to disperse in the diffusive fashion modeled, but is more actively removed and replaced by cooler material, presumably by convective processes. This stepped shape to the subsidence histories is quite distinct from a normal oceanic crustal model. A similar history is noted at the Varing Plateau (Site 643) although in this case the wide span of paleo-water depths after the end of volcanism makes it impossible to estimate the duration of anomalous temperatures. Site 918 is the exception to this simple pattern of thermal evolution because, after an initial period of subsidence at an elevated temperature, this site shows increasing temperatures, peaking at around 42 Ma, a pattern that Clift et al. [15] interpreted as reflecting the closest approach of the Iceland Plume to the site, during its migration across the Greenland coast at this time [47]. This post-rift temperature maximum is not seen at Site 553, which is approximately conjugate to Site 918, and it is an observation that is difficult to reconcile with a ridge-centered plume [3]. As mentioned above, Site 336 presently lies over asthenosphere estimated at 25°C above normal and, consequently, the modeled curves are matched with the present corrected depth to basement at the + 25°C anomaly level. The reconstructed curves appear to show approximately constant asthenospheric temperatures since the start of the sedimentary record. Unfortunately, the presence of a large hiatus at the base of the section here means that no information

on the temperature under the ridge during eruption of the basement can be extracted.

4. Oceanic crustal thicknesses The relationship between oceanic crustal thickness (i.e., melt thickness) and asthenospheric temperature has been highlighted by McKenzie and Bickle [48]. Using the revised relationship for high asthenospheric temperatures calculated by White and McKenzie [3] it is possible to estimate the temperature of the asthenosphere at the time of break-up from the crustal thickness, assuming that melting was a passive adiabatic process. By comparing the temperature estimates derived from this source with the estimates taken from the subsidence anomalies it will be possible to see if simple decompression melting of anomalously hot asthenosphere is sufficient to generate the igneous thicknesses observed. The thickness of the crust itself can be measured seismically, and appropriate surveys have been made across the margins on the Edoras Bank, Voring Plateau, and East Greenland (transects A, B, and C, Fig. 1). Barton and White (121 estimate a thickness of around 15- 17 km for the oldest oceanic crust at the Edoras Bank, close to the COT. Mutter and Zehnder [49] estimate a maximum thickness of 24 km across the COT in the central Voring Plateau, although a value of around 20 km may be more appropriate for Site 343, which is on the oldest true oceanic crust, slightly west of this maximum. In addition, Site 643 might be expected to have thinner crust because of its proximity to the Jan Mayen Fracture Zone. A crustal thickness on the order of 16-18 km is predicted here. Larsen and Jakobsdottir [50] provide estimates of around 18 km for the East Greenland margin near ODP Site 918. Recent deep seismic surveys on the Iceland-Faeroe Ridge indicate crustal thicknesses of around 25 km for the crust between the Faeroes and the Iceland Plateau, with the crust approaching 30 km in thickness under present day Iceland [5]. These figures are also broadly in agreement with the deep structure inferred from regional gravity and bathymetric data [511. An alternative method of estimating crustal thickness relies on assigning standard densities for the upper and lower crust and performing an isostatic

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643 343 918 552 553 336 East Iceland

calculated

from present depth to basement

(km)

(km)

3.151 3.350 2.607 2.522 2.689 1.231 - 0.055

5.0 5.0 5.0 5.0 5.0 4.9 3.1

1.8 1.6 2.4 2.5 2.3 3.7 3.8

17.2 16.1 20.2 20.7 19.8 27.3 27.8

Present corrected depth to basement

Present depth to basement without temperature anomaly

(km)

(km)

(km)

85 86 86 86 86 85 54

0.023 0.023 0.023 0.023 0.023 0.094 0.745

3.128 3.327 2.583 2.499 2.665 1.137 - 0.800

at each of the drill sites are shown corrected

calculation to determine what thickness of crust is required to produce the depth to basement observed today. Mutter and Zehnder [49] noted that increases in the oceanic crustal thickness were principally manifest as increases in the thickness of the sheeted dikes and gabbros, rather than the eruptive sequences. In this study the model of Mutter and Zehnder [49] was used and therefore the ratio of lavas and dikes versus gabbro thickness is 3:4. A mean density of 2800 kg/m3 is used for the lavas and dikes, while 3000 kg/m3 is used for the gabbros. A mean density of 3330 kg/m3 is attributed to mantle peridotite at normal temperatures. Since the present depth to basement is not only controlled by anomalously high crustal thickness, but also by modem thermal anomalies in the underlying asthenosphere, this effect must be corrected for before the crustal thickness can be estimated. As estimated above, a temperature anomaly of 5°C is predicted for each of the passive margin drill sites from the model of Ito and Lin [32], while 25°C is the value used for Site 336 on the Iceland-Faeroe Ridge. Once the depth of basement has been recalculated to account for the asthenospheric temperature anomaly, as well as the sediment loading removed by the backstripping calculation (Table l), then this value can be compared with the expected depth to basement for normal crust of this age [24]. The difference between the model and the reconstructed depth to basement can then be used to determine the thickness of the crust at each site, as shown in Table 1. The estimates for Sites 336, 643 and 918 compare well with the

Predicted

(km)

Uplift due to temperature anomaly in asthenosphere

Present depths to basement anomalies.

Depth anomaly

Predicted depth to basement for normal oceanic crust (km)

Lithospheric thickness

for sediment

loading and estimated

asthenospheric

crustal thickness

temperature

estimates from seismic surveys. In the case of Site 343 the crustal thickness measured seismically capproximately 20 km) is rather greater than that derived from the depth to basement (16.1 km). The reason for this discrepancy is not clear, but it may relate to the distance between the site and the precise track of the seismic survey, so that the estimate derived seismically may not be truly applicable to the drill site. In addition, if the densities of the lower crust are higher than normal, due to compositional variations, such as the presence of ultrarnafic cumulates, while retaining velocities more typical of the crust, then the seismic survey will predict a greater crustal thickness than the isostatic calculation. Similarly, the thickness at Site 553 is greater than inferred from seismic data, and this may relate to the presence of relatively light continental material close to the COT. As a test for this method the crust of eastern Iceland was also considered, using an age of 12 Ma and an elevation of 800 m above sea level to represent the land close to the modem hotspot center. An asthenospheric temperature anomaly of 150°C is taken from Ito and Lin [32]. In this case the calculated crustal thickness is 27.8 km, corresponding well with the approximately 30 km thick crust measured by wide angle seismic methods [5].

5. Melting models Using these estimates with the decompression melting model of White and McKenzie ([3]; Fig. 5) it

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A

E

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G30

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p E c! s 10 r

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% 0

0 1200

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1300

1400

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is seen that the inferred asthenospheric temperatures from crustal thickness are in most cases higher than those predicted from the subsidence reconstructions at the time of break-up. At Site 336 there is an inferred hiatus of approximately 8 Myr, between eruption of the initial volcanics and the first transgression, so no good subsidence-derived temperature estimate exists for the time of break-up. However, at all other sites the time gap is small, as a result of rapid transgression of the lavas by the sea. Using the 100 km thick plume model, Site 343 shows asthenospheric temperatures hot enough to have generated the oceanic crust by adiabatic decompression melting alone, and this may also be true at Site 553 if the seismically derived crustal thickness is accepted. However. at Sites 918 and 643 the maximum temperatures are insufficient to account for the melt thickness. This is also true of the 200 km thick plume head, the preferred model for an initiating plume, in which only Site 343 is hot enough for the melt thickness observed. These results indicate that melting at Sites 918 and 643, and probably Site 553, was not purely adiabatic but that melt volumes were probably increased by melting greater volumes of asthenosphere, implying enhanced convection under the earliest spreading ridge.

6. Thermal evolution

The subsidence records at each of the drilled volcanic margins provide for the first time a record of the long-term thermal evolution of this area. Sites 553, 918, 343, and 643 share common features in the predicted evolution of their underlying asthenosphere. It is apparent from the subsidence histories that significant positive thermal anomalies are restricted to the period after rifting, and there is no evidence to indicate the anomalies’ presence even until the Miocene. The drill sites of the Voring Plateau (Sites 343 and 643) both show evidence for elevated asthenospheric temperatures under the rift at the time of break-up, an anomaly that lasts at least 6

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Myr according to the record from Site 343. The measured anomaly of 50-100°C above normal mantle at Site 643 (80-180°C at Site 343) is approximately in accord with the estimate of 50-80°C of Skogseid et al. [52], derived from melt volumes and degrees of crustal extension measured on the continent side of the COT. In addition, Zehnder et al. [53] predicted an asthenospheric temperature of this range based on the Na,O content of lavas from the Vdring Plateau, after correcting for the effects of fractional crystallization. The large unconformity at Site 343 removes the rest of the Paleogene record so that the duration of the anomaly cannot be determined, and the water depth constraints at Site 643 are too poor to provide a useful estimate. A more complete record is shown by Site 553. Here, the temperature anomaly lasts 5 Myr following eruption of the seaward-dipping lavas at this site. The record at Site 918 shows a similar pattern to Site 553 but in this case the anomaly is seen to last longer, at least 14 Myr, and possibly as long as 20 Myr, suggesting an additional heat source at this margin compared to its conjugate. Site 336 lacks the crucial record of the earliest post-spreading history, so that no inference can be drawn regarding melt processes under the earliest Iceland-Faeroe Ridge. Nonetheless, it appears that the temperature of the asthenosphere under Site 336 has remained relatively constant over the past 44 Myr, with the suggestion of a slight long-term rise with time to the present.

7. Melting processes in the rift zone The thermal histories deduced from the subsidence histories of the volcanic rifted margins can be used to assess the possible influence of the Iceland Plume on their development. It is clear that, at all sites analyzed, the temperature of the asthenosphere underlying the rift axis was elevated above normal mantle during the time of continental break-up. This observation is at odds with the model of Mutter et al. [16l, which invoked only induced convective up-

Fig. 5. Diagram showing the thickness of melt generated by adiabatic decompression of asthenosphere temperatures. Estimated thicknesses for oldest oceanic crust at Sites 553. 918, 343, and 643 are shown, temperature anomalies above normal mantle derived from the subsidence patterns.

over

together

a range of potential with the estimated

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welling of normal asthenosphere to explain the large amounts of melting at this time. However, a comparison of the temperatures deduced from the subsidence histories with those predicted from the thickness of the oceanic crust at each site, using the melting model of White and McKenzie [3], indicates a discrepancy. The asthenosphere under the rifted margins at Sites 918 and 643, and probably 553, was about 50-100°C hotter than normal mantle at the time of break-up, but was not hot enough to generate the observed thicknesses by simple adiabatic melting of upwelling asthenosphere. In order to generate the observed crustal thicknesses at the temperatures deduced from the reconstructed subsidence histories it would be necessary to melt greater volumes of mantle than would be involved in passive upwelling of asthenosphere under the rift axis. In this case upwelling has to exceed spreading rate. Exactly what is causing this enhanced convection cannot be deduced from these observations; however, higher asthenospheric temperatures, lowering viscosity, will facilitate more rapid convection of the mantle, while the stresses related to rifting may initiate upwelling into the rift axis. The duration of the temperature anomaly at the Edoras Bank is comparable to the 5-8 Myr duration for the thermal anomaly inferred from the width of thickened crust offshore the eastern US margin [19]. Although not so well constrained; the observations at the Voting Plateau are also compatible with a relatively short-lived thermal anomaly; however, the incompleteness of the stratigraphic record at Site 343 and the great water-depth uncertainties at Site 643 make the exact duration of the anomaly difficult to assess. Given the similar crustal thicknesses and duration of the thermal anomaly at the Edoras Bank (Site 553) and the eastern US margin, it may be questionable how much influence the Iceland Plume had on the generation of the seaward-dipping sequences at the Edoras Bank, if not the Voring Plateau, too. The case for the Bdoras Bank being principally the product of rifting rather than plume-related processes is strengthened by the N-MORB chemistry of the lavas recovered there [lo]. However, trace element and isotopic evidence does suggest some enriched plume influence on petrogenesis at the VBring Plateau and on the East Greenland margin [22]. Site 918 differs from the other drilled margins in

showing a much longer-lived thermal anomaly in the post-rift period. Although the subsidence history at Site 918 does confirm a thermal anomaly at the time of break-up, Clift et al. [15] noted that the period of maximum relative uplift, and thus thermal anomaly, occurred at about 42 Ma, some 12 Myr post-rift, which coincides with the time at which the Iceland Plume crossed the East Greenland coast on its relative migration eastward, according to the hotspot track of Lawver and Miller [47]. At this time the plume makes it closest approach to Site 918, resulting in the observed post-break-up temperature maximum. Although it is impossible to quantify the relative contributions of the Iceland Plume and a thermal anomaly related to the rift at the time of break-up, geochemical evidence from recovered basalts at Site 918 [22] suggests generation of the main section of the seaward-dipping lavas at 63”N offshore East Greenland due to remelting of an Icelandic source depleted by earlier melt extraction. In this case it would be expected that at least some of the thermal anomaly is the result of the presence of the Iceland Plume, assuming that the geochemical anomaly is smaller than the thermal anomaly, as it is at the present day [54].

8. Role of the Iceland Plume The 5 Myr long thermal anomaly, the thick oceanic crust, and the MORB-type geochemistry of the basalts in the seaward-dipping sequences at the Edoras Bank all compare strongly with the dipping reflectors of the eastern US margin [19] and thus raise the possibility that the Northeast Atlantic might have rifted and formed seaward-dipping lavas and thickened oceanic crust even if the Iceland Plume had not been located nearby in the late Paleoceneearly Eocene. Anomalous asthenospheric temperatures must have existed below the rift prior to the arrival of the plume, and could represent the insulating effect of the overlying continental lithosphere in storing heat before its dissipation in rifting episodes [ 181. In contrast, the subsidence data from Site 918 show that the Iceland Plume did provide significant thermal input into the region after about 44 Ma. Chemical evidence indicates that some thermal input would have affected those parts of the rift axis

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strong impact that the Iceland Plume has had on the development of the Northeast Atlantic.

Acknowledgements I wish to thank the Woods Hole Oceanographic Institution for support during the course of this study. Peter Kelemen, Steve Holbrook, Garrett Ito, and Trine Dahl-Jensen provided thought-provoking discussion on my subject. Garrett Ito helped with the generation of the oceanic subsidence curves. C. Fox copyedited the manuscript, figures and table. Godfrey Fitton, Chip Lesher, and Bob White provided constructive reviews that greatly improved this manuscript. This is WHO1 contribution number 9276. [CL1

References ItI K. Him, A hypothesis on terrestrial catastrophes: Fig. 6. Reconstruction of the Northeast Atlantic area at 52 Ma, end of anomaly 24, after [7]. Plume center is from [47], with each concentric circle marking 200 km from the plume center. GIR = Greenland-Iceland Ridge; IFR = Iceland-Faeroe Ridge; FSE = Faeroe-Shetland Escarpment; VP = Voting Plateau; ROB = Rockall Bank; S = Shetland Islands; HE = Hatton Bank. Bathymetry is in hundreds of meters.

closest to the plume center at the time of break-up (Fig. 6). The topography of the Greenland-Iceland Ridge, inferred to be a plume track by White and McKenzie [3], does not suggest a major thermal contribution by the plume to the rift axis in the earliest Eocene. This feature is very poorly developed in its oldest section, as shown by bathymetry and gravity data [%I. The ridge only gains its full topographic expression after about anomaly 20 time [9], corresponding to an age of 42-46 Ma. This reinforces the idea that the Iceland Plume has only had a strong effect on the development of the Northeast Atlantic after that time. Since the middle Oligocene the generation of the GreenlandIceland-Faeroe Ridge, the Iceland Plateau, and the modem day shallowing of the Reykjanes Ridge towards Iceland, all provide concrete evidence for the

wedges of very thick oceanward dipping layers beneath passive continental margins, their origin and paleoenvironmental significance, Geol. Jahrb. E22, 3-28, 1981. l21 R.S. White, A hot-spot mode1 for early Tertiary volcanism in the N. Atlantic, in: Early Tertiary Volcanism and the Opening of the NE Atlantic, AC. Morton and L.M. Parson, eds., Geol. Sot. London Spec. Publ. 39, 3-13, 1988. [31 R.S. White and D.P. McKenzie, Magma&m at rift zones: the generation of volcanic continental margins and flood basalts, J. Geophys. Res. 94, 7685-7729, 1989. 141R.S. White, J.W. Bown and J.R. Smallwood, The temperature of the Iceland plume and origin of outward-propagating V-shaped ridges, J. Geol. Sot. London 152, 1039-1045, 1995. El R.S. White, J.H. McBride, P.K.H. Maguire, B. Brandsdottir, W. Menke, T.A. Minshull, K.R. Richardson, J.R. Smallwood, R.K. Staples and the FIRE group, FIRE: Faeroe-Iceland Ridge Experiment, EOS Trans. AGU, in press. I61 H.C. Larsen, A.D. Saunders, P.D. Clift, et al., Proc. ODP, Init. Rep. 152, 1994. [71 G.D. Karner and A.B. Watts, On isostasy at Atlantic-type continental margins, J. Geophys. Res. 87, 2923-2948, 1982. @I G.N. Kharin, G.B. Udintsev, O.A. Bogatikov, J.I. Dmitriev, H. Raschka, H. Kreuzer, M. Mohr, W. Harre and F.-J. Eckhardt, K/Ar age of the basalts of Norwegian-Greenland Sea, Glomar Challenger, Leg 38, Init. Rep. DSDP 38, 755760, 1976. [91 A.G. Nunns, Plate tectonic evolution of the Greenland-Scotland Ridge and surrounding areas, in: Structure and Development of the Greenland-Scotland Ridge, M.H.P. Bott et al., eds., pp. 11-30, Plenum, New York, NY, 1983.

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