The canadian atlantic margin: A passive continental margin encompassing an active past

Tectonophysics, 59 (1979) 83-126 @ Elsevier Scientific Publishing Company,

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THE CANADIAN ATLANTIC MARGIN: A PASSIVE CONTINENTAL MARGIN ENCOMPASSING AN ACTIVE PAST

R.T. HAWORTH

and C.E. KEEN

Atlantic Geoscience Centre, Geological Survey of Canada, Bedford Oceanography, Dartmouth, N.S. (Canuda) (Received

Znstitute of

March 9,1979)

ABSTRACT Haworth, R.T. and Keen, C.E., 1979. The Canadian Atlantic margin: tal margin encompassing an active past. In: C.E. Keen (Editor), across Passive Margins. Tectonophysics, 59: 83-126.

a passive continenCrustal Properties

The continental margins of Atlantic Canada described in this paper show the effects of plate tectonic motions since Precambrian time and thus represent an ideal natural laboratory for geophysical studies and comparisons of ancient and modern margins. The Grenville Province shows vestiges of Helikian sedimentation on a preexisting continental block beneath which there may have been southeastward late-Helikian subduction resulting in collision between the Grenville block and the continental block comprised of the older shield provinces to the north. The Grenville block was subsequently split in Hadrynian time along an irregular line so that the southeastern edge of the Grenville exhibited a series of promontories and reentrants similar to those seen at the present Atlantic continental margin of North America. That margin, which had a passive margin history perhaps comparable with that of the present Atlantic margin, was separated by the Iapetus ocean from the Avalon zone whose Precambrian volcanism has been attributed both to that associated with an island arc and with intracratonic rifting. However, the Iapetus ocean appears to have been subducted in early Paleozoic time with a southeastward dip beneath the Avalon zone, leaving exposures of oceanic rocks in place as in Notre Dame Bay, or transported onto Grenville basement as at Bay of Islands. Plate motions proposed for Devonian and Carboniferous time are numerous, but resulted in the welding of the Meguma block to the Avalon zone of New Brunswick and northern Nova Scotia, extensive faulting within Atlantic Canada which can be correlated with contemporaneous European faulting, and extensive terrestrial sedimentation within the fault zones. Graben formation, continental sedimentation and basaltic intrusion in the Triassic represent the tensional prelude to the Jurassic opening of the present Atlantic Ocean. This Jurassic opening produced a rifted margin adjacent to Nova Scotia and a transform margin along the southern Grand Banks. The width of the ocean-continent transition across the transform margin (approx. 50 km) is narrower than for the rifted margin (approx. 100 km). The eastern part of the transform margin is associated with a complex Cretaceous (?) volcanic province of seamounts and basement ridges showing evidence of subsidence. The western portion of the transform margin is non-volcanic, adjacent to which lies the 360 km wide Quiet Magnetic Zone floored by oceanic crust. Development of the margin east of Newfoundland was more complicated with conti-

84 nental fragments separated from the shelf by deep water basins underlain by foundered and atypically thin continental crust. Although thin, the crust appears unmodified, the similarities between the crustal sections of the narrow Flemish Pass and the wide Orphan Basin suggesting that the thinning is not simply due to stretching. The Newfoundland Basin shows evidence for two-stage rifting between the Grand Banks and Iberia with both lateral separation and rotation of Spain, leaving a wide zone of transitional crust in the south. The overall pattern of variations in crustal section for the margin east of Newfoundland is comparable with that of the British margin against which it is located on paleogeographical reconstructions. The major sedimentary unconformities on the shelves (such as the Early Cretaceous unconformity on the Grand Banks) reflect uplift accompanying rifting, Tracing of the sedimentary horizons across the shelf edge is complicated by paleocontinental slopes, which separate miogeocline and eugeocline depositional environments. The subsidence of the rifted margins is primarily due to cooling of the lithosphere and to sediment loading. The subsidence due to cooling has been shown to vary linearly with (time)“‘, similar to the depth-age behaviour of oceanic crust. The consequent thermal history of the sediments is favourable for hydrocarbon generation where other factors do not preclude it. INTRODUCTION

The continental margin of southeastern Canada (Fig. 1) exhibits almost the whole range of characteristics that may possibly exist at a passive continental margin. It is therefore an ideal area in which to define those characteristics by which the situations at other continental margins may be judged. In this paper we will describe the structural framework of this continental margin as we know it today, indicate the extent to which geological and geophysical investigations have proceeded, and indicate where additional work is necessary to cover those deficiencies in knowledge that remain. We agree with Williams and Stevens (1974) that “an understanding of present continental margins is a key to the understanding and delineation of ancient margins”, and that “conversely, an understanding of ancient continental margins, their depositional and structural evolution, and the position and nature of the transition between continental and oceanic crust aids in understanding recent margins”. Williams and Stevens (1974) provided an excellent expose on the geological development of the ancient margin. We hope to complement that by stressing the geophysical attributes of that ancient margin, extending its structures offshore and to depth, and discussing the structures and processes active at the present margin, primarily through the interpretation of geophysical measurements. Fig. 2, which depicts the overall structural setting of the margin with as much detail as feasible, will serve as the focal point of our discussion of margin characteristics. The history of development of the margin can be summarized as follows. In Precambrian time the Iapetus ocean (Harland and Gayer, 1972) separated the land masses whose remnants we now see as the Grenville. and Avalon cratons. In Paleozoic time, that ocean closed, the remnants of that ocean being well exposed on the north central coast of Newfoundland in the central portion of the Appalachian orogen. In Meso-

Fig. 1. Index map of Atlantic Canada showing place names and structural and geophysical features discussed in the text. QZ = outer boundary of the Magnetic Quiet Zone; J = J-magnetic anomaly; SD. = Shelbume dike; A.D. = Avalon dike; C.L.F. = ChanceportLobster Cove fault; ++ = indicates Gravity Edge Anomaly.

m

cn

A

APPALACHIANS

100

ktlometres

200

AVALON GRAND BANKS

ORPHAN

BASIN ORPHAN

KNOLL

:

Ij

s..

.

.

.~’

”

N. ATLANTIC

Fig. 2. Geophysical cross-sections across the margins and location maps (lower right). The crustal layering is based on seismic refraction lines whose positions are shown by numbers on the location map. The numbers are related to the original reference and lint numbers below. Velocities are shown in km/s. The upper portions of the cross-sections are based upon the geological corn~iiati~~n oi

GRENVILLE

..

B

19. 20. 21. 22. 23. 24. 25. 26. 27. 28. 29. 30. 31. 32. 33. 34.

1. 2. 3. 4. 5. 6. 7. a. 9. 10. 11. 12. 13. 14. 15. 16. 17.

Berry and Fuchs (19’73), Grenville Front. Berry and Fuchs (1973), Grenville Province. Hobson and Overton (1973). Dainty et al. (1966), East Point. Dainty et al. (1966), Mainland. Dainty et al. (1966), Cape Forolle. Dainty et al. (1966), St. Anthony. Bankin et al. (1969), Gasp& Sheridan and Drake (1968), Line 151. Dainty et al. (1966), Tracadie. Sheridan and Drake (1968), Line 150. Dainty et al. (1966), Cheticamp. Dainty et al. (1966), Cape Freels. Dainty et al. (1966), Port aux Basques. Press and Beckmann (1954), Line 16. Sheridan and Drake (1968), Line 148 with 4.9/6.6 interface elevated to match horizon determined by slope to 147. Fenwick et al. (1968), Continental line AB (8.0 horizon doesn’t match because crust in Orphan Basin is thinner than North of Dover Fault). Barrett and Keen (1978), Orphan Basin line. Fenwick et al. (1968), Line GH. Fenwick et al. (1968), Line CD. Fenwick et al. (1968), Line EF. Press and Beckmann (1964). Line 12. Press and Beckmann (1954), Line 15. Jackson et al. (1976), Line 6. Jackson et al. (1975), Line 5. Bentley and Worzel(1956), Line 7. Bentley and Worzel(1956), Line 6. Barrett et ai. (1964), Cole Harbour. Dainty et al. (1966), Sable Bland. Keen et al. (1975), Line 4, Keen et al. (1975), Line 2. Keen et al. (1975), Line 1. Keen et al. (1975), Line 3,

Williams (1978) and, in submerged areas, on GSC Map 1400A, on assorted seismic reflection lines and deep exploratory well data referred to in the text. The uppermost cross-section depicts the crustal structure along line AB and crosses ancient as well as modem margins, The two shorter cross-sections (lower left) depict the structure across the transform margin (uppermost, CD on location map) and across the rifted margin (lowermost, EF on location map). The following list gives the relation between numbers on the location map and original references and line numbers for the refraction data,

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zoic time, rifting was initiated, and a new ocean developed whose margins lay within the (Avalon) craton that formed the eastern margin of the PrcCambrian ocean. That ocean is still opening. Its western margin (off southeastern Canada), although passive, contains evidence for the existence of an earlier ocean that had both passive and active phases. The evolution of the margin includes the following, which will be discussed chronologically(a) The Precambrian cratons: The Grrnville and Avalon are wrll exposed and have received detailed geological attention, The location of the edges of the cratons may be identified at depth by geophysical means, the geophysical data also providing the means by which the craton may be located offshore. The basement to the Meguma is nowhere exposed and therefore its relation to the Grenville or Avalon is unknown. (b) Remnants of Paleozoic ocean crust: Well exposed in north central Newfoundland, with allochthonous slices present along much of the Grenville margin in Canada, Paleozoic oceanic rocks have been extensively compared with samples obtained by deep drilling in the present oceans. By the application of geophysical methods, their subsurface extent has been partially determined as a clue to their history of emplacement. (c) Sedimentation and rifting: Both the cratons and the Paleozoic oceanic rocks are partially covered by late Paleozoic sedimentary rocks from which the (pre-rifting?) tectonic history of the region may be deduced. The effects of subsequent early Mesozoic rifting have been preserved in the Bay of Fundy . (d) Transform and rifted margins: The location of the Mesozoic rift within which ocean crust developed east of the Scotian Shelf was offset from that which developed east of the Grand Banks by a transform fault that roughly follows the southern edge of the Grand Banks. The mature rifted margin off Nova Scotia and the mature transform margin south of the Grand Banks have been investigated geophysically to determine their contrasting characteristics. They form the western and northern margins of a classic quiet magnetic zone. (e) Complicated margins north and east of Newfoundland: East of the Grand Banks, the continental fragments of Orphan Knoll and Flemish Cap lie at different distances from the morphological edge of the continental shelf. The nature of the crust landward of the fragments provides clues to the mechanism of Mesozoic rifting. The Newfoundland Basin which hosts the Newfoundland Seamounts, contains a similarly anomalous transition zone in which lies the evidence for the pre-rift configuration of Iberia and North America. (f) Uplift, subsidence and sedimentation: Draping every unit of the margin and drilled at various locations as part of hydrocarbon exploration programs, Mesozoic and Cenozoic sedimentary rocks provide evidence of the tectonic history of the region; in particular its vertical movements provide clues to deeper processes.

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THE PRECA~BR~N

CRATONS

The Grenville Province has an average width of approximately 400 km, extending from the Grenville Front, its junction with the older provinces of the Canadian Shield in the north, to the St. Lawrence, beneath which it has most of its contact with the younger Appalachian structural province to the south. Deposition in the Grenville orogen began 1600 m.y. ago, followed by extensive intrusion between 1400 and 1100 m.y. (Wynne-Edwards, 1972). Subsequent erosion of much of the supracrustal geosynclinal sequences has left much of its basement complex exposed. The basement contains structures with ages between 1700 and 2500 m.y., so that at least part of the orogen was built upon a preexisting continen~l plate. Although there is considerable variety between the geological models proposed for development of the Grenville Province, most models presume that plate tectonic mechanisms were active in the Precambrian (Baer, 1976). Burke and Dewey (1973) suggested that the Grenville Province has a thick crust as a result of continental collision of the Tibetan type, the suture having existed south of the present southern edge of the Grenville Province in the present relative location of the Appalachians. Wynne-Edwards (1972) saw no evidence for a suture line joining two separate continental plates within or at the boundary of the Grenville. However, Brown et al. (1975) believe there is evidence for oceanic rocks in southeastern Ontario, potentially in agreement with the paleomagnetic data that indicate movement of the southern part of the Grenville Province with respect to the remainder of the shield between 1125 and 1000 m.y. (Irving et al., 1974). Reconciliation between these interpretations is likely to proceed only with the benefit of investigations using seismological and gravity data which have been the primary geophysical adjunct to the geological work. Berry and Fuchs (1973) determined an average crustal thickness of 39 km in the Grenville Province compared with 34 km in the Superior Province to the north. There is local thickening of the crust to 45 km close to the Grenville Front which everyone seems to agree is not the precise location of a suture zone, although it may be an artifact of suturing (Baer, 1974). The average Grenville crustal velocity of 6.6 km/s overlies an upper mantle velocity of 8.1 km/s, and Mereu and Jobidon (1971) suspect that the velocity of the crustal material just above the Moho is at least 7 km/s. Preliminary analysis by Rankin et al. (1969) of arrivals in Gasp6 from the same explosive sources used by Mereu and Jobidon (1971), Berry and Fuchs (1973) gave an upper mantle velocity of 8.75 km/s at a depth of 52 km within the contact zone between the Grenville and Appalachian provinces. This is in good agreement with the velocity of 8.5 km/s at a depth of 46 km found by Ewing et al. (1966) at Tracadie, although their more ex-tensive data suggest such upper mantle velocities are more typical of the Appalachian Province than of the Grenville-Appalachian boundary. Gravity analysis has concentrated either on such specific structures as the anorthosites (Kearey, 1978) within the Grenville Province, or the Gren-

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ville Front itself (Thomas and Tanner, 1975), but the best overview of much of the Grenville is presented by Thomas (1974). Thomas and Tanner (1975) have analyzed the 40-50 mgal gravity low lying parallel to, and approximately 50 km south of, the Grenville Front and interpret that subduction occurred from northwest to southeast between the Superior and Grenville provinces. Their model is based on all the available gravity and seismic data, and although it may appear to contradict some geological interpretations which do not consider the Grenville Front to be a suture zone, the two can be more easily reconciled if consideration is given to the extent of the main structural units in three dimensions. For example, at an active continental margin, oceanic lithosphere is subducted beneath continental lithosphere with a zone of several hundred kilometres separating a “pure” continental from a “pure” oceanic lithospheric section (Isacks and Molnar, 1971). On completion of continental suturing, at any one erosional level there will be a surface line defining contact between the two continental blocks. However, the deeper the erosional level, the closer will that line be to the edge of the continental section beneath which subduction took place. Geophysically we see the entire width and thickness of the transition zone whereas geologically we tend to see only the surface contact. For the northeastern edge of the Grenville, the obvious gravity anomaly representing the transition (suture zone) between the two different continental sections therefore lies southeast of the Grenville Front which is defined as the surface geological contact, Seismic data (Dainty et al., 1966) show that the crustal thickness isgreater in the Appalachians than in the Grenville Province. In northern Newfoundland, for example, the crustal thickness of the Grenville Province is approximately 30 km, increasing to approximately 45 km within the central Appalachian zone of New Brunswick and Newfoundland (Fig. 2). This change in crustal thickness coincides with a general increase in gravity across the Grenville-Appalachian boundary (Weaver, 1967) that can be traced as a general feature of the western edge of the Appalachians within the United States (Diment et al., 1972). There is some difficulty tracing this gravity anomaly from Gasp4 to Newfoundland in order to locate the edge of the Grenville within the Gulf of St. Lawrence because of the superposition of the gravitational effects of overlying Carboniferous sedimentary basins reaching their greatest depth near the Magdalen Islands (Watts, 1972; Haworth, 1978). The magnetic expression at the junction between the Grenville and Appalachian provinces in the eastern U.S. (Zietz et al., 1966) is not as evident in the Canadian Appalachians where ophiolites, either autochthonous or allochthonous adjacent to the junction give rise to localized magnetic highs that have an overwhelming effect on the magnetic field (Haworth and MacIntyre, 1977). However, the combination of the gravity and magnetic signatures of the boundary is sufficient to identify that a major salient of the Grenville underlies much of the Gulf of St. Lawrence (Haworth, 1978), and that the edge of the Grenville then lies immediately east of northern Newfoundland and off southern Labrador (Haworth et al., 197613; Fig. 1).

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The Grenville salient within the Gulf of St. Lawrence is the northernmost expression of a series of offsets that lay along the southeastern margin of the Grenville “continent”. Rankin (1976) interpreted these offsets as having originated because the Proterozoic continent, of which the Grenville basement was part, broke up (rifted) along a line joining a series of triple junctions, so that the southern edge of the Grenville was the western margin of the Iapetus ocean. Each of the angularities that cut into the Grenville was therefore interpreted as being the seaward termination of a “failed arm” aulacogen that can be traced into the Grenville. Burke and Dewey (1973) have recognized several such aulacogens at the edge of this and other margins. Thomas (1977) interpreted the offsets as being the transform segments of the western margin of Iapetus in the same way that transform segments such as the southern margin of the Grand Banks create offsets at the present margin of the Atlantic. The transform margin interpretation of Thomas (1977) does not necessarily imply a triple junctionaulacogen origin, while Rankin’s (1976) interpretation is a further development of Burke and Dewey’s ideas that require such an origin. Common to these interpretations is the recognition that features similar to those on the present Atlantic margin should (unless erosion has removed them) be preserved at the southeastern margin of the Grenville Province if plate tectonic processes as we know them today were active in the Proterozoic. Rankin (1976) and Thomas (1977) imply that such processes were indeed active then, having found Precambrian and Paleozoic analogs of present-day continental margin structures. At each of the indentations at the edge of the Grenville, Rankin (1976) concluded that the increased thickness of Precambrian-Cambrian elastics lay within an aulacogen trending into the Grenville. The Ottawa-Bonnechere graben is the best example of such a trough, but geophysical anomalies trending into the Grenville at the other indentations are interpreted by Rankin to define the location of mafic igneous rocks signifying the failed rift activity of the other aulacogens. The Monteregian intrusions are an effect of reactivation of the Ottawa-Bonnechere graben in Cretaceous time. Burke and Dewey (1973) have proposed that aulacogens are a common feature along all presentday margins, and that they are responsible for initial continental break-up. Thomas (1977) primarily examined the geology on the seaward side of the Proterozoic continental margin and recognized: Proterozoic microcontinents such as the Sauratown Mountains of North Carolina, equivalent to the modern Flemish Cap; Late Precambrian-Early Cambrian elastic-volcanic sequences such as the Ocoee-Mt. Rogers groups representing opening rift sequences, equivalent to those of the modem Gulf of Maine-Bay of Fundy region; Early Cambrian-Silurian shallow-marine carbonate bank facies such as those within the northern Gulf of St. Lawrence, equivalent to for example the Jurassic Iroquois Formation on the modem Scotian Shelf; and postCambrian elastic wedge sequences such as the Blount Group of Tennessee prograding from the carbonate bank, equivalent perhaps to the Verill Canyon

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Formation involved in the constructional-destructional phases of modern margin development as discussed later. The apparent existence of Precambrian and Paleozoic analogs of modern continental margin structures suggests that an iterative approach may be possible towards a determination of the processes responsible for their development. For the present margin we have extensive seismic control on structure, but few geological samples, limited to relatively shallow depth. In the case of most hydrocarbon wells, the samples are generally only well cuttings obtained from anomalous structures. The ancient analogs which Rankin (1976) and Thomas (1977) have found can provide complementary information because they are generally on dry land where exposed segments can be examined in detail, the older analogs having been eroded to depths from which almost no sample information is available from the present margin. The detailed geological information from these analogs can then be- used to hypothesize greater detail within the present margin. Such hypotheses can be tested, for example by deep drilling at selected locations, and, if correct, will both substantiate the reality of the analogs and permit better hypothesis of the processes active at present margins. The Avalon zone formed the eastern margin of the Iapetus during the Late Proterozoic and early Paleozoic eras. Whereas the Grenville is well exposed over a large area in Canada, the Avalon is only well exposed in the small area of eastern Newfoundland, with more limited unquestionable exposure in Nova Scotia, New Brunswick and Massachusetts. However, Williams (1978) has recognized equivalents over a wider area in these regions and extended identification of the Avalon zone to the southeastern United States. The AvaIon zone is characterized by a variety of Late Precambrian volcanic and sedimentary rocks, whose origin (tectonic setting during deposition) is still not known, Hughes and Bruckner (1971) considered it to have originated as part of a volcanic island complex, with the implication therefore that the eastern margin of Iapetus was the active one adjacent to which the oceanic crust was subducted. Alternatively the volcanic highs and sedimentary lows have been interpreted structurally as a series of horsts and grabens that formed as the result of rifting preceding the opening of Iapetus (Papezik, 1970; King, 1977). This dispute has to be settled, since inability to resolve the dispute implies an inability to recognize the tectonic state of ancient margins on the basis of its preserved fragments. Because of the limited extent of Avalon outcrop, the ability to recognize geophysically the pattern of Avalon structural units offshore, and thereby deduce their structural history, may be invaluable. The volcanic highs are associated with gravity and magnetic anomalies and seismic velocities that are all significantly higher than those associated with the sedimentary lows (Haworth and MacIntyre, 1975). On the continental shelf east and south of Newfoundland the sedimentary-volcanic alternation continues to be seen on the basis of seismic, gravity and magnetic data and indicates that the Avalon zone underlies much of the Canadian margin between the Charlie fracture zone and the

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Collector Anomaly, including the deep water area of the Orphan Basin (Haworth, 1977; Haworth and Lefort, 1979). Geological support for suggesting that the Avalon zone extends as far east as Flemish Cap comes from samples of granite from the cap, for one sample of which a latest Precambrian age was obtained (Pelletier, 1971). Correlation of the Avalon zone with Precambrian structures in Iberia has been carried out (Lefort and Haworth, 1979; this volume) and indicates continuity of the Avalon craton over distances greater than the width of the Grenville zone. Whereas an insight into Avalon structure has been possible using reconnaissance data, a more detailed correlation between geological and geophysical indicators is necessary to reach the level of detail available for the Grenville margin. We must, for example, determine how good is the separation of seismic velocities between the volcanic and sedimentary elements of the Avalon zone. Also, a deep seismic refraction profile within, but close and parallel to the western edge of, the Avalon zone would provide comparison with the passive Grenville margin and perhaps help to indicate whether the Precambrian plate tectonic setting at the Avalon margin was active or passive. The only refraction results at present show that the crustal thickness is of the order of 30 km, similar to that seen at the eastern edge of the Grenville (Dainty et al., 1966). The southern half of Nova Scotia, separated from the Avalon zone by the major transcurrent Cobequid--Chedabucto fault (Webb, 1968; Eisbacher, 1969), exposes early Paleozoic sedimentary rocks (the Meguma Group) that are atypical of eastern Canada. Schenk (1971) interprets the sedimentary rocks as having been deposited in a lower continental rise and abyssal fan setting, the continental source of sediments lying to the south with respect to the Meguma as seen today, that source having been subsequently displaced by continental drift. Schenk interprets that source to have been North Africa, with Nova Scotia having lain adjacent to Morocco in the early Paleozoic. Structural correlation between the two areas supports this view (Lefort and Haworth, in press). There seems little doubt that the Meguma zone underlies most of the Scotian Shelf. Magnetic anomalies associated with narrow mineral-rich zones within the Meguma sedimentary rocks can be traced offshore, particularly off southern Nova Scotia (Bower, 1962) and northeastern Nova Scotia. Parallel magnetic zones can also be seen on the Scotian Shelf, but as the thickness of younger sediments increases, inference of the extent of the Meguma can only be based on more subtle changes. McGrath et al. (1973) infer from second derivatives of the magnetic field that Meguma sedimentary rocks do not extend entirely to the present continental margin. The Devonian granites that pervasively intrude the Meguma have a sufficiently low density that readily permits their spatial mapping from gravity data (O’Reilly, 1975). Their gravity lows extend across the Scotian Shelf both to the south (Watts, 1974) and to the east of Nova Scotia (Stephens et al., 1971), such that they form a zone that parallels the curvature of the magnetic mineral zones of

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the Meguma, and the fault contact between the Meguma and the Avalon zone (Haworth and MacIntyre, 1975). That zone of low gravity appears to extend to the edge of the continental shelf. Its interpretation as a zone of granites has been related to subduction at the boundary between the Avalon and Meguma zones as discussed later. The contact between the Meguma and the Avalon zone is marked by the magnetic Collector Anomaly (Haworth, 1975), which cuts across the Tail of the Banks, where, south of the Collector Anomaly there is a gravity low. The Tail of the Banks may therefore have a Meguma basement. Seismic refraction studies on the Atlantic coast (Barrett et al., 1964), within the Meguma, show a relatively simple crust with total thickness of 33-36 km similar to that within the Avalon zone given the inaccuracies likely in the existing Avalon data. The Meguma sedimentary rocks have a total thickness of 9 km (Barrett et al., 1964; Poole, 1970). The major crustal layer (with VP = 6.1 km/s) of the Meguma zone lies at a minimum depth of a few kilometres, but little evidence exists regarding the nature of this basement. On Schenk’s model, the basement could be either oceanic or continental. If continental, but deep water, we might expect to see features characteristic of the Orphan Basin, or other areas adjacent to present continental margins that have undergone large subsidence. Since the crustal structure of such areas as Orphan Basin is presently enigmatic, having a possible ancient analog may prove important, for example, in deciphering the thermal history of such an area. The magnetotelluric measurements that indicate a highly conductive structure beneath the eastern side of the Appalachian system, which also extends beneath the Meguma on the Scotian Shelf (Cochrane and Hyndman, 1974), may be the result of the quasi-continental structure of such an area. Alternatively, if the Meguma overlies what was oceanic crust, a present-day analog of it might be the Laurentian Fan which overlies oceanic crust off the Scotian Shelf. In summary, the Precambrian and lower Paleozoic stable areas forming much of the continental edge of the presentday Canadian continental margin are the remnants of ancient continental margins: Grenville passive continental margin. active continental margin or intra-cratonic rifted province or Avalon both. deep water sedimentation on subsided continental or oceanic Meguma basement. Although these elements may be considered stable parts of a modem passive margin they do preserve evidence of an active past, from which we may deduce details that increase our understanding of present margin tectonics. REMNANTSOFPALEOZOICOCEANICCRUST Modem plate tectonic concepts have led to the interpretation of the sequence of lower Paleozoic vokanic and sedimentary rocks within New

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Brunswick and Newfoundland’s central mobile belt (Williams, 1964) as the remnants of the ocean (Iapetus) that separated the Grenville and Avalon zones. Many syntheses have discussed the geology of this central belt in those terms (e.g. Dewey, 1969; Bird and Dewey, 1970; Church and Stevens, 1971; Strong et al., 1974; Williams, 1975) and this review will not attempt to duplicate these, concentrating primarily on recent geophysical results that pertain to its development. Allochthonous slices of ophiolitic rocks found on or adjacent to the western margin of the Appalachian system have been interpreted as thrust from the Paleozoic ocean during its closure (e.g. Williams, 1975; St. Julien and Hubert, 1975), the age of emplacement of the allochthon in western Newfoundland being Early Ordovician (Stevens, 1970). This dates the climax of the erogenic events associated with closure of the Paleozoic ocean as seen on the western margin of that ocean. Suturing between the eastern and westem continental margins occurred in the southwestern portion of Newfoundland (Brown, 1976) at a salient of the Grenville craton. This contact between the Grenville and Avalon cratons, combined with the oblique approach between the two cratons (Haworth, 1975), provided a protected environment north of the suture (Notre Dame Bay) in which the Paleozoic oceanic rocks were preserved. Geophysical evidence has delineated and extended seaward and to depth the early Paleozoic structures of Notre Dame Bay that are the vestiges of that Paleozoic ocean (Haworth et al., 1978). The volcanic sequences that outcrop on the southern side of Notre Dame Bay can be traced northeastward across the continental shelf by their magnetic signature (Haworth et al., 1976a). The ultramafic rocks that underlie similar units in the vicinity of Betts Cove (DeGrace et al., 1976) may be shown from gravity and magnetic data to be near surface along the seaward extension of the Chanceport-Lobster Cove fault north of Fogo Island. Notre Dame Bay and its northeastward prolongation is therefore a synclinal structure in which the entire range of oceanic crustal rocks has been preserved. The early Paleozoic rocks on the southern margin of Notre Dame Bay (Strong, 1973; Strong and Payne, 1973; Kean and Strong, 1975) appear to form a sequence that is representative of a typical active island arc complex as inferred by Mitchell and Reading (1971), but the position of the arc is ambiguous on the basis of the geological section at any single location (for example, compare the southward subduction invoked by Strong (1973) with the northward subduction invoked by Kean and Strong (1975) for the same rock units). The bouldery mudstone of the Dunnage melange, southeast of Notre Dame Bay, has been interpreted as melange deposition in a trench above a northwest dipping Paleozoic subduction zone (Kay, 1972), a view supported by Williams and Hibbard (1976), but denied by Jacobi and Schweickert (1976) on the grounds that such melange deposits are not uniquely associated with island arc complexes. Models for emplacement of the ophiolitic allochthon have invoked subduction both to the northwest (e.g., Williams, 1975) and to the southeast

(e.g., Strong, 1977). Sections of the ophiolitic suite that are complete, and possibly rooted, have been preserved. The best of these, on Burlington fenin sula, Newfoundland are the Baie Verte and Betts Cove ophiolite complexes. High magnetic anomalies that can be traced adjacent to the edge of the Grenville craton in a more or less continuous fashion all the way to southem Labrador (Haworth et al., 1976b) are associated with those ultramafic sections. The magnetic properties of the ophiolitic units have been determined directly from those units outcropping around Notre Dame Bay (Deutsch et al., in prep.) and also inferred from the observed magnetic anomalies over the known structures of Bay of Islands, Hare Bay, Baie Verte and Betts Cove. These properties may then be used in analysis of the magnetic anomalies observed offshore to locate and determine the structure of the root sources for the allochthonous slices (Srivastava et al,, 1977). Whereas the Bay of Islands complex appears to have been transported at least 200 km from its origin in Notre Dame Bay (Williams, 1975) the root of the Hare Bay complex appears to be only approximately 20 km from the shore (Srivastava et al., 1977; Haworth et al., 1978). Samples drilled from within the zone of high magnetic anomalies associated with this inferred root zone have confirmed the presence of rocks correlative with those in the allochthon (Haworth et al,, 1976a). This substantiates the inferred presence of an ophiolitic root zone stretching from Notre Dame Bay to southern Labrador (Haworth et al., 1976b), this overall zone probably being the source of the allochthon of both western and northern Newfoundland. Seismic refraction experiments to define the structure of this ophiolitic root zone offshore should provide evidence for the origin and mode of emplacement of the onshore units. Despite the apparent inability to uniquely infer from the geological evidence provided by the ophiolites the direction of subduction that destroyed the Paleozoic ocean, both geochemical (Strong et al., 1974) and geophysical (Haworth et al., 1978) data favour subduction to the southeast beneath the Avalon zone. The primary geochemical evidence is that the potassium content of plutons in eastern Newfoundland increases eastward, implying by comparison with the increase in potassium content of igneous rocks with landward distance from an island arc (Dickinson, 1968), that Paleozoic subduction in Newfoundland had an eastward component. Mineralization also shows a zonation in Newfoundland which is analogous with that across and adjacent to Pacific island arcs. Subsequent to the Strong et al. (1974) analysis of plutonic geochemistry, the age and origin of the plutons has been questioned (David Strong, personal communication, and Strong and Dickson, 1978) thereby casting doubt on the inferred direction of subduction. However, southeastward subduction is also supported by the geophysieal evidence. Beneath the central and eastern portions of Newfoundland, a crustal layer with a seismic velocity of 7-7.5 km/set was found by Sheridan and Drake (1968). That layer follows a regular pattern dipping to the southeast with respect to the pre-Carboniferous surface. It lies closest to the surface

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adjacent to the southeastern edge of the Grenville orogen, beneath which it does not extend. At that edge of the Grenville orogen, are found the ophiolite suites that are the remnants of the Paleozoic ocean (Williams, 1975), and within which are gabbros generally having compressional velocities of 6.87.4 km/set, and ultramafics with compressional velocities of the order of 8.4 km/set (Salisbury and Christensen, 1978). Such velocities are comparable with those within the major layers of oceanic crust (Raitt, 1963; Christensen and Salisbury, 1975) and are not found extensively elsewhere within Newfoundland. They can be traced down dip to the southeast, beneath the Avalon peninsula, indicating that the ocean within which the ophiolites originated was subducted to the southeast beneath Avalon (Haworth et al., 1978). The interpretation of an island arc suite existing on the southern margin of Notre Dame Bay (Strong, 1977) and the Gander ultramafic zone as a zone of mantle diapim above a subduction zone (Stevens et al., 1974) are compatible with this interpretation that the continental margin of the Avalon zone was an active one with subduction taking place to the southeast. Cochrane and Wright (1977) have shown from geomagnetic data that a zone of high conductivity exists beneath eastern Newfoundland. They correlate this with the existence of a fossil subduction zone by analogy with similar conductivity anomalies over similarly inferred subduction zones in Scotland (Hutton et al., 1977). Bailey et al. (1978) have summarized the geomagnetic data for eastern North America, but although similar correlations can be inferred in various localities along the length of the Appalachians, such conductivity anomalies are in apparent continuity across major Appalachian boundaries leaving the impression that their regional significance may not be directly related to the existence of a subduction zone. Stewart (1978) has shown that PP arrivals in Europe from central American earthquakes are missing for a reflection zone adjacent to the western margin of the Avalon zone. He suggests that this “shadow zone” could be due to asymmetry in the teleseismic reflection path which might be caused by the residual effect of a dipping subduction zone. The results do not directly indicate which way dipping occurs, but since the shadow (and hence the dipping of the lithosphere) lies at the eastern edge of the central “mobile” zone, southeast of the island arc sequences, by analogy with the mapped cross-section of presentday subduction zones where the major lithospheric dip lies landward of the trench axis (Isacks and Molnar, 1971), the lithospheric dip must have an eastward component. The geophysical and geochemical data are therefore unanimously indicative of a southeastward dipping Paleozoic subduction zone in Newfoundland. Seismic velocities of the order of 7.2 km/set are not, however, confined to the layer observed within the upper 15 km of the crust. Long seismic refraction lines within and adjacent to the Appalachian system (Dainty et al., 1966) revealed a layer with velocity of 7.3-7.5 km/set whose upper surface lies at 20-30 km depth, extending to Moho (VP approximately 8.5 km/set)

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at a depth of 45 km. The analysis of Dainty et al. (1966) assumed the shallow 7+ km/set layer observed by Sheridan and Drake (1968) to be part of an upper crustal layer having an average velocity of 5.9 km/set. These similar velocities at different depths have to be resolved but probably are the result of the shallow, thin 7+ km/s velocity layer not being resolved by the longer refraction lines which were shot perpendicular to the strike of the structures. The deeper 7+ km/set layer is being recognized more and more in continental areas, lying in general between 30 and 40 km depth within the Eurasian platform where its composition is unknown but is assumed to have a high ultramafic content (Pavlenkova, 1979, this volume). The relationship between the coexisting shallow and deep ultramafic(?) layers beneath the Appalachian system has still to be resolved. Following early Paleozoic closure of Iapetus, the determination of subsequent plate tectonic activity during the Paleozoic in Atlantic Canada is problematical. Granitic intrusion was widespread in the Atlantic provinces in Devonian time. The plutons of Devonian age in Newfoundland have a patchy surface expression scattered over a wide area (Williams, 1967; Strong et al., 1974; Bell and Blenkinsop, 1975) but the large negative gravity anomaly that underlies the area of their greatest concentration in eastern central Newfoundland (Haworth, in press) provides an indication that ostensibly individual plutons may have a common source at depth. Strong and Dickson (1978) argue that the correspondence between the composition of the plutons and their host rock in various tectonic settings implies that the granites in northeast Newfoundland were formed by partial melting. In the southern (Meguma zone) part of Nova Scotia the granites onshore form a more cohesive group with the gravity data providing a southward (Watts, 1974) and eastward (Stephens et al., 1971) extension into a curved zone parallel to the southeastern unit of the Avalon zone in New Brunswick and Nova Scotia (Haworth, 1975). Since some granites have been recognized as a product of partial melting above subduction zones (Brown, 1973), and the Cobequid--Chedabucto fault separating the Meguma and Avalon zones appears to have undergone major transcurrent motion during Devonian time (Webb, 1968), major plate tectonic activity within Nova Scotia, including Devonian subduction, was proposed (Haworth, 1975). Poole (1976) suggested that this subduction was the northwestward subduction of an ocean that lay between southern Nova Scotia and northern Nova Scotia-southeastem New Brunswick. However, this scenario cannot account for the location of the granites, if they were produced above a subduction zone, located as they are in southern Nova Scotia beneath which Poole (1976) does not have subduction taking place. Production of the granites through partial melting related to subduction processes can only have occurred if those processes were also active beneath southern Nova Scotia. The plate motions deduced by Irving (1977) from paleomagnetic data imply that an ocean existed between Laura&a and Gondwana (of which southern Nova Scotia was probably part) in Middle Devonian time. Irving

99

(1977) infers collision of southeastern Canada with northwestern South America in Late Carboniferous time, followed by collision with northwestem Africa in Late Triassic time. Apparent commonality between Late Carboniferous faulting on the continental margin of North America between Massachusetts and Newfoundland with that off western Europe (Lefort and Haworth, 1978) and northwest Africa rather than northwestern South America (Lefort and Haworth, in press) is not compatible with these paleomagnetic results. Also, neither of these interpretations seems compatible with the fauna1 evidence from cores obtained northeast of Newfoundland (Haworth et al., 1976a) indicating that a deep, wide seaway (ocean?) separated eastern Canada from Ireland in Early Carboniferous time (Jansa et al., 1978). It is apparent from the discrepancy between these interpretations that much work remains to be done in developing a compromise between the geological and geophysical interpretations that fits all the facts. SEDIMENTATION

AND RIFTING

Whatever plate movements were effective in Devonian time, they were probably also responsible for the extensive faulting and subsequent Carboniferous deposition that affected the entire Atlantic region. The faults lie generally within a narrow northeast trending zone from Bay of Fundy to White Bay, Newfoundland (Webb, 1968). The faults may have been created during the final stages of lower Paleozoic collision between obliquely approaching continents and subsequently reactivated in Carboniferous time (Haworth, 1975). In basins bordered by these faults, Carboniferous deposition was widespread. In the central Gulf of St. Lawrence, and southeastern Prince Edward Island, the thickness of Carboniferous sedimentary rocks exceeds 9 km, the width of the depositional zone decreasing rapidly but maintaining considerable depth (“6 km) as it approaches New Brunswick and Newfoundland (Hobson and Overton, 1973; Howie and Barss, 1975). Northeast of Newfoundland the basin widens again with Carboniferous deposition over much of the continental shelf (Haworth et al., 1976a, b). Evidence of a change in the deformation of the Carboniferous sediments is seen northeast of Newfoundland. Whereas the faulting in the Atlantic provinces is primarily aligned northeastward, and the older Carboniferous sediments are confined by these faults, the younger Carboniferous sediments north of 52”N latitude have fold axes (although the folding is gentle) that veer northwestward, implying a different constraint on their deformation. That deformation marks the completion of the erogenic events of the Appalachian orogen, to be succeeded in Triassic time by the tensional prelude to opening of the present Atlantic Ocean. Throughout eastern North America, the Late Triassic was generally a period of normal faulting, graben formation and continental sedimentary deposition, accompanied by the extrusion and intrusion of basaltic lavas and

sills (Ballard, 1974). Exposures of Triassic rocks in southeastern Canada are limited in extent to the borders of Chedabucto Bay and the Bay of Fundy beneath which the Triassic strata are interpreted to extend (King and Maclean, 1976). The Triassic sedimentary rocks attain a thickness of over 1 km in Nova Scotia, and are all non-marine (Poole, 1970). They thicken to over 2 km on the western side of the Bay of Fundy, the downdropped side of the half graben (King and Maclean, 1976). Within the Gulf of Maine, sedimentary rocks lying in fault-bounded basins have been recognized form seismic reflection profiles (Ballard and Uchupi, 1972; King and Maclean, 1976). These sedimentary rocks are inferred to be of Triassic age because of their structural continuity with basins in the Bay of Fundy and the western Scot,ian Shelf, and because Late Triassic faults provide control over the deposition (Ballard, 1974). The narrow Orpheus Basin east of Chedabucto Bay, linked to the Fundy Basin by faults of the Cobequid-Chedabucto fault system, contains up to 1 km of Triassic non-marine sedimentary rocks. It is only towards the edge of the present continental shelf that the Triassic facies become evaporitic in nature (Jansa and Wade, 1975; Jansa et al., 1977) showing a trend towards more marine conditions in the east. Jansa and Wade (1975) conclude that open marine conditions can only have existed during Triassic time in the area of North Africa and southern Iberia that lay adjacent to southeastern Canada at that time. The North Mountain basalt, which conformably overlies the Triassic sedimentary rocks on the eastern shore of the Bay of Fundy is tholeitic and reaches a thickness of over 400 m (Jansa and Wade, 1975). The basalts can be recognized offshore by their magnetic signature and can be traced from their northeastward trend along the eastern shore of the Bay of Fundy into a more northerly trend across the Gulf of Maine. The basalts are also exposed on Grand Manan Island adjacent to the New Brunswick coast, and on islands in the upper reaches of the Bay of Fundy. Triassic dikes may also be widespread as an indicator of the tensional regime in which the area existed prior to final rifting. A dike roughly paralleling the southeastern coast of Nova Scotia, the Shelbume dike, can be traced onshore for a distance of approximately 100 km, and by its magnetic signature it can be traced offshore for a similar distance both south and east of Nova Scotia. The Shelbume dike has a latest Triassic age, almost contemporary with the Bay of Fundy basalts. The presence of another dike, the Avalon dike, has been inferred from the existence of a linear magnetic anomaly crossing the Avalon peninsula of Newfoundland (Papezik et al., 1975). The Avalon dike’s magnetic anomaly can be traced offshore to a location 90 km east of St. John’s, Newfoundland, and there is an indication that it also extends offshore to the southwest. Hodych (personal communication) reports sampling what he believes to be outcrop of the Avalon dike, the measured paleomagnetic pole for the sample being compatible with a Triassic age. Linear, northeast trending magnetic anomalies are also a pervasive feature of the continental shelf northeast of Newfoundland. Diabase has been

101

drilled from a structure correlated with one of those linear anomalies, but its age has not been determined. The general pattern throughout the Bay of Fundy and Gulf of Maine of graben development, sedimentation and igneous activity is consistent with the classical interpretation (for the more extensive Triassic in the eastern United States) of Triassic uplift representing rebound of the thickened Appalachian crust finally released from the compressive erogenic forces that produced and deformed it (Rodgers, 1970). The faulting and sedimentation was interpreted to be a crustal effect which would be primarily constrained by the preexisting Appalachian structures. The igneous activity and subsequent dike injection was presumed to be associated with deeper processes and therefore more likely to be related to the mechanism of continental separation. Ballard (1974) concluded that whereas the Appalachian structures were finally controlled by a right lateral shear couple along a northeastward trending fault system, relaxation of that stress by a left lateral shear couple would produce a tensional regime controlling graben development. Late Triassic-Early Jurassic continental separation, which occurred east of Nova Scotia, then produced renewed faulting and differential movement of the Fundy graben, causing folding of the strata within them. TRANSFORM

AND RIFTED

MARGINS

The rifted margin southeast of Nova Scotia and the large transform margin along the southern Grand Banks were created during the early Mesozoic stages of formation of the present Atlantic Ocean when Africa separated from North America. The Nova Scotian margin is part of the extensive margin stretching from Florida to the Laurentian Channel. The ocean crust south and east of these margins is as old as Jurassic and Cretaceous in age and various oceanic volcanic provinces can be recognized whose tectonic significance may be relevant to the development of the margin, particularly the transform margin. The deep structural characteristics of the rifted and transform margins have been defined mainly from geophysical studies. We briefly review the results obtained so that the more recent studies along the more complex margins to the north can be reviewed in their proper perspective. Two cross-sections, one across the transform margin and one across the rifted margin are shown in Fig. 2. These were compiled from a variety of sources including deep exploratory well data, seismic refraction, reflection, gravity and magnetic data. Some of the main conclusions regarding the rifted and transform margins are summarized below: (1) To the south and east of the margins lies the Quiet Magnetic Zone, exhibiting magnetic anomalies no larger than 50-100 nT. This zone, some 350 km wide, appears to be underlain by oceanic crust (Pitman and Talwani, 1972; Barrett and Keen, 1976), although some studies have suggested that it is occupied by foundered continental or “transitional” crust (Drake et al., 1968; Gradstein et al., 1977). This zone is magnetically quiet because of the

absence of magnetic reversals during the time that oceanic crust was formed there (Larson and Hilde, 1975; Barrett and Keen, 1976). The characteristics of the basement rocks seen on seismic reflection profiles are similar to those of oceanic basement observed elsewhere and this basement reflector can be traced westward through the Quiet Zone to the foot of the continental slope where it is obscured by large thicknesses of sediment. (2) The Quiet Zone is bounded to the east by a series of lineated magnetic anomalies of Cretaceous age, the Keathley sequence (Larson and Hilde, 1975; Barrett and Keen, 1976). The strike of the anomalies and of the eastern edge of the Quiet Zone is roughly perpendicular to the southern transform margin of the Grand Banks (Fig. 1). The most prominent feature of the Keathley sequence, the J-anomaly, is associated with a ridge of oceanic basement which intersects the southern Grand Banks margin near its eastern termination (Sullivan and Keen, 1978; Rabinowitz et al., in press). (3) Running sub-parallel to the rifted margin are linear, gravity and magnetic anomalies (Fig. 1; the East Coast Magnetic Anomaly, Taylor et al., 1968). Many hypotheses have been presented concerning their origin (Keen, 1969; Emery et al., 1970; Rabinowitz, 1974; Keen et al., 1975; Sobczak, 1975). Most theories imply that the east coast magnetic anomaly is associated with the ocean-continent boundary, while the gravity anomaly may be caused by a combination of the change in crustal thickness at the margin, and the presence of the thick sedimentary wedge which is not entirely isostatically compensated. Parallel to the transform margin, linear magnetic anomalies are not observed and the amplitude of the gravity anomaly is about one-half that found off Nova Scotia (Jackson et al., 1975). (4) Seismic refraction measurements on the rifted margin have shown that the crust is about 32-36 km thick on the east coast of Nova Scotia and about the same thickness under Sable Island, near the shelf edge (Barrett et al., 1964, Dainty et al., 1966). However, the crustal layer with velocity 6.16.3 km/s thins between the mainland and the shelf break, from 30 to 20 km, and the total crustal thickness beneath Sable Island is maintained at 35 km by the additional thickness of over 10 km of Cenozoic-Mesozoic sediments. On the seaward side of the margin, about 100 km from the shelf break, oceanic crust is observed, with a total depth to the Mdiscontinuity of 15 km (Fig. 2). The boundary between oceanic and continental crust cannot be resolved by seismic measurements alone to better than within this 100 km wide zone centered on the continental slope. Refraction measurements along the transform margin indicate a relatively sharp transition zone between oceanic and continental crust, less than 50 km wide, and in the eastern part the oceanic crust seaward of the transition zone is defined by the presence of the Fogo Seamounts province (Hall et al., 1977) which lies within 70 km of the shelf break (Jackson et al., 1975). Little thinning of the continental crust is apparent (Fig. 2) although more deep refraction measurements on the shelf are necessary to confirm this. Several questions of fundamental importance can be posed regarding these

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transform and rifted margins. The position of the oce~~ontinent boundary appears to be better defined along the transform margin, as might be predicted. The position and width of the transition zone along the rifted margin are more controversial and depend upon the evidence used. If the velocities of the main crustal layer (continental = 6.1-6.3 km/s, oceanic = 6.7-7.5 km/s) are used to define the ocean-continent boundary, it would lie near the shelf break, south of Sable Island. If the east coast magnetic anomaly is used to define the boundary, it lies near the foot of the slope. In this case the continental slope region, about 50 km wide, represents transitional crust characterized by high seismic velocities (7.4 km/s) in the main crustal layer and the magnetic anomaly is caused by the edge effect between magnetic oceanic rocks and relatively non-ma~etic, tr~sitional and continents rocks (Keen et al., 1975). The latter interpretation is more consistent with plate tectonic reconstructions (Rabinowitz, 1974; Klitgord and Schouten, in prep.) which use the magnetic anomaly as a marker of the boundary and coincides with changes in the sedimentary strata near the foot of the slope (Jansa and Wade, 1975; King et al., 1975). No basement ridge such as that suggested by Drake et al., (1959), Emery et al., (1970), and Sheridan (1974) lies beneath the shelf break nor is there any evidence for a substantial carbonate reef there (King et al., 1975). However, the distinction between post-rifting sedimentary rocks of mid-Jurassic age and basement (Meguma) rocks, both of which exhibit velocities in excess of 5 km/s is difficult seismically. Some of the layering identified as “basement” in Fig. 2 may be in fact part of the overlying sedimentary strata. The transform margin appears to have experienced very little thinning of the continental crust relative to the rifted margin. The two most commonly suggested means of thinning the margins, subaerial erosion, and necking of the crust under tension (Bott, 1971; Artemjev and Artyushkov, 1971; Sleep, 1971), would not necessarily lead to similar thinning beneath rifted and transform segments. If sub-aerial erosion during uplift was largely responsible for the observed crustal thinning, the rifted margin may have experienced a longer period of uplift and therefore more thinning than the transform margin. The amount of uplift and erosion along the transform margin will be variable; similar to that along the rifted margin near the junction between the two and less along the transform away from the rift zone before sea floor spreading began. During sea floor spreading, the ridge crest, associated with high temperatures and uplift, migrated along the transform margin producing a transient uplift of the margin which may have been responsible for further erosion. Necking of the crust might be expected in association with tensional forces at rifted margins and may play a relatively small role along transform margins (Scrutton, 1976). Near the southern transform margin of the Grand Banks lies a complex volcanic province. This includes the Fogo Seamounts, probably Cretaceous in age, which form part of the oceanic crust adjacent to the margin, the Newfoundland Ridge, a major physiographical lineament which extends south-

1 o-1

east from the Tail of the Banks region and appears to be underlain by oceanic basement highs, and the Spur, or J-anomaly Ridge, another basement ridge trending perpendicular to the Newfoundland Ridge and associated with the J-magnetic anomaly (Sullivan and Keen, 1978; Rabinowitz et al., in press). The elevated basement topography and seamounts near the transform margin suggests excessive volcanism, which may be either a common property of transform margins, or the result of passage of a “hot-spot” through the area. Gradstein et al., (1977) and Grant (1977) have suggested an alternative explanation for this volcanic region: that the Spur Ridge and at least part of the Newfoundland Ridge is underlain by subsided continental crust. This is not supported by the studies of Sullivan and Keen (1978) and we favour the interpretation of the region as oceanic. The arguments supporting continental material beneath these areas are discussed elsewhere (Grant, 1977, 1979, this volume). DSDP Site 384 is located on the Spur Ridge. The discovery of shallow water limestones of mid-Cretaceous age just above basaltic rocks which are now at a depth of 4100 m suggests that the Spur Ridge, and perhaps much of the volcanic province mentioned above, has subsided about 4100 m since mid-Cretaceous time (Tucholke et al., 1975). As Gradstein et al. (1977) show, this implies a total subsidence comparable with that measured on the southern Grand Banks in the same time interval. The total subsidence is greater than that associated with normal oceanic crust (Sclater and Francheteau, 1970). If the area is, as we suggest, oceanic, these results could have two explanations. First, if the oceanic and adjacent continental lithosphere is coupled across the transform margin, the effects of thermal contraction, and other deep-seated causes of subsidence, and sediment loading would act upon both oceanic and continental sides to produce similar subsidence. Second, if a hot spot were located along the transform margin during the Cretaceous (Burke et al., 1973), an anomalously large amount of subsidence might be expected, similar perhaps to the subsidence associated with mid-plate hot spots like the Hawaiian Chain (Detrick and Crough, 1978). The presence of a hot spot is supported by a study of teleseismic PP-P travel time residuals (Stewart and Keen, 1978), which show a large anomaly beneath the Fogo Seamounts, where upper mantle velocities must be low to explain the observations. Stewart and Keen (1978) have suggested this is due to a zone of residual high temperatures and partial melting in the upper mantle which was produced by a Cretaceous “hot spot” or partially molten mantle plug from which volcanic material is now represented in the Fogo Seamounts was injected. The junction between the rifted and transform segments is not well defined, partly because of the large quantities of post-rifting sediments on the Laurentian Fan which obscure basement structures. Volcanic seamounts are not found along the transform margin west of 53”W and large normal faults affecting basement. are observed there. This region was the location of the 1929 Grand Banks earthquake (Heezen and Drake, 1964; Emery et al.,

105

1970). The change in structural style at 53”W is not well understood and further work is required in the area. It is possible that this boundary reflects a change in the nature of the motion between Africa and North America. Perhaps the faulted western region indicates the area where extension between the plates was primarily by stretching of the continental crust, while the seamount region marks the start of true sea floor spreading between them. This would place the ocean-continent boundary along the rifted margin further seaward than discussed previously, within the eastern part of the Quiet Zone and some 200 km east of the east coast magnetic anomaly. It is generally anticipated that transform margins develop along old lines of weakness in the continent (Francheteau and Le Pichon, 1972; King et al., 1975). It is therefore perhaps surprising that this margin does not follow the Avalon-Meguma fault boundary discussed earlier. Instead the margin appears to lie within the Meguma Platform. King et al. (1975) suggest that the continental prolongation of the line of weakness lies along the AvalonMeguma boundary in Nova Scotia (the Cobequid-Chedabucto fault system) but this is not consistent with the divergence between this boundary and the transform margin on the Grand Banks. The western extension of the transform more closely follows an apparent structural discontinuity within the Avalon in the vicinity of the Laurentian Channel (Haworth and Lefort, 1979).

COMPLICATED MARGINS NORTH AND EAST OF NEWFOUNDLAND

North of the rifted and transform margins produced by the AfricanNorth American plate motions, the continental margins are fragmented into segments of different ages and different characteristics. The diversity in character of the margin arises because of the timing and geometry of plate motions between North America, Iberia, and Europe (Fig. 3). During the Early Cretaceous, spreading began between Iberia and the Grand Banks and formed the present rifted margins of the Newfoundland Basin, east of the Grand Banks (Keen et al., 1977). North of the Newfoundland Basin, sea floor spreading between Europe and North America began in the Late Cretaceous, forming the margins east of Flemish Cap and Orphan Knoll and at about the same time the separation of Rockall Plateau from the shelf northeast of Newfoundland occurred (Haworth, 1977; Le Pichon et al., 1977; Srivastava, 1978). The development of these margins is complicated by the formation of the continental fragments, Flemish Cap and Orphan Knoll (Fig. 1). Between these features and the continental shelf proper lie relatively deep basins in which the water depths (1000-2800 m) are intermediate between those of the deep oceans and of the continental shelves. These basins, Flemish Pass and Orphan Basin, are bounded to the north by the Charlie fracture zone, forming a transform segment of the margin north of Orphan Knoll. North

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of this fracture lies a more normal rifted margin along the northeast Newfoundland shelf. The nature of the margins of the Newfoundland Basin has been discussed by Sullivan (1978) who showed that the ocean-continent boundary east of the Grand Banks is marked by a deep sediment filled graben near the foot of the slope, about 100 km wide, which may be occupied by “transitional” or foundered continental crust. This graben is associated with a quiet magnetic zone, seaward of which a large magnetic anomaly trending north-northeast has been identified as the northern extension of the J-anomaly (Keen et al., 1977). This anomaly denotes the oldest identifiable marker in oceanic crust within the Newfoundland Basin. A similar graben marks the ocean-continent boundary south of Flemish Cap. The J-anomaly has not been recognized in that area. The directions of oceanic magnetic lineations within the Basin suggest two directions of motion during the separation of Iberia from the Grand Banks (Sullivan, 1978). First, a phase of east-west motion during the time of the J-anomaly (115 m.y.) produced a rifted margin east of the Grand Banks. During this period, sea floor spreading occurred only in the southern part of the Newfoundland Basin, implying a component of counter-clockwise rotation of Iberia and no motion between the Grand Banks and Galicia Bank. Second, at about 100 m.y. a change in spreading direction took place, from about 105” to 145”. The spreading centre migrated northward initiating spreading between Galicia Bank and the Grand Banks and within the Bay of Biscay. During this second phase, a rifted margin south of Flemish Cap was developed. The two-phase development of the Newfoundland Basin explains how rifted margins characterized by similar structural features developed around its perimeter, even though the strike of the margin changes from N-S to ENE-WSW at about 45”N. A single E-W plate motion of Iberia relative to the Grand Banks would have resulted in a transform margin south of Flemish Cap. The width of the zone of transitional crust appears to be greatest in the southern part of the basin. This alleviates the problem created by the gap between the continents south of Galicia Bank which has often been encountered in pre-rift reconstructions of Iberia and North America (Le Pichon et al., 1977, Lefort and Haworth, 1978). However, a significant gap remains, and further studies in this problem area are required on both sides of the Atlantic. The Newfoundland Seamounts form a roughly linear east-west line across the centre of the Newfoundland Basin. The geochemistry of the basalts and trachytes recovered from them suggest affinities with volcanic rocks forming oceanic island chains and not with mid-ocean ridge basal& One of the trachytes has been dated at 100 m.y. but it is not known whether an age progression occurs along the seamount chain (Sullivan and Keen, 1977). Keen et al. (1977) suggest that the seamounts may reflect reactivation in

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oceanic crust along the prolongation of an old line of weakness in the continent, perhaps along the Avalon-Meguma boundary (Haworth, 1975) and that they represent volcanic activity along a leaky transform, related to the shift in spreading direction at about 100 m.y. (Sullivan, 1978). Volcanic rocks of similar age have been sampled on the Grand Banks just south of the Avalon-Meguma boundary (Gradstein et al., 1977) so that reactivation of this faulted boundary may have affected the continental crust also. The Newfoundland Basin is bounded to the south by the Newfoundland Ridge described previously. In plate tectonic terms it may mark the eastern extension of the Grand Banks transform but differential subsidence across the ridge, with greater subsidence to the south, may have played an equally important role in its development. Shallow water elastic limestones of midCretaceous age have been dredged from the top of one of the Newfoundland Seamounts, now at a depth of 3000 m. South of the Newfoundland Ridge at DSDP site 384 subsidence of about 4100 m has been measured during the same time interval suggesting a differential subsidence of 1100 m. This difference may have been accommodated across the Newfoundland Ridge. North of the Newfoundland Basin, the continental fragment of Flemish Cap from which granodiorites of Precambrian (Avalon) age have been sampled (Pelletier, 1971) is separated from the Grand Banks by the Flemish Pass. The latter is a linear feature, some 1000 m deep and 50 km wide, which opens at its northern end into the wider (250 km) and deeper (2800 m) Orphan Basin which separates the Orphan Knoll continental fragment from the shelf northeast of Newfoundland. Orphan Knoll is a positive bathymetric feature whose summit lies at a depth of 1800 m. It was drilled at DSDP site 111 and shown to be underlain by rocks at least as old as Jurassic, leading to the conclusion that it is a small detached piece of continent (Ruffman and Van Hinte, 1973). The results at site 111 indicate that Orphan Knoll was subjected to a period of uplift and erosion during the Late Jurassic and Early Cretaceous, followed by slow subsidence in the Late Cretaceous and extremely rapid subsidence in the early Tertiary. The Orphan Knoll has apparently been near its present depth since the Eocene and most of the subsidence took place in the Paleocene (Ruffman and Van Hinte, 1973). Analysis of subsidence on the adjacent continental shelf and near Flemish Pass using data in deep exploratory wells shows a similar subsidence history there, with rapid subsidence in the early Tertiary (Keen, 1979; R.A. Folinsbee, personal communication). The times of uplift and subsidence on the shelf and on Orphan Knoll correlate reasonably well with the rifting and spreading episodes in the southern Labrador Sea which started to open in Late Cretaceous time (Srivastava, 1978). These results suggest that the entire area was radically affected by the Labrador Sea opening and that this event may be responsible for both the subsidence and for the present configuration of the continental fragments with respect to the shelf. In some earlier studies, oceanic crust was postulated to occupy the Orphan

Basin (Sheridan, 1974), as if a failed arm of a mid-ocean ridge once occupied the area, but more recent analyses of gravity, magnetic and seismic data have shown that this region is underlain by thin continental crust. Haworth (1977) showed that the magnetic trends, expressing trends within the Avalon Precambrian basement, cross the shelf break and extend into Orphan Basin. He further suggested that the boundary between the Avalon Platform and the Paleozoic rocks to the north has a seaward expression in the Charlie fracture zone, marking the northern edge of the Orphan Basin and of the Avalon Platform. Thus the true ocean-continent boundary lies east of Orphan Knoll and Flemish Cap and a transform margin lies north of Orphan Basin. Few results are available from this transform margin. Large positive free air gravity anomalies lie at the edge of the shelf landward of Orphan Basin and Flemish Pass (Haworth, 1977). These have been interpreted by Folinsbee (in prep.) to be caused by thinned continental crust extending from Orphan Basin inshore beneath the shelf. This thin crust beneath the shelf is isostatically uncompensated, as it has not subsided sufficiently under the large thickness of sediments now near the shelf break. The crust beneath the deep central part of Orphan Basin does appear to be isostatically compensated. The zone of thin, or alternatively exceptionally dense, crust beneath the shelf is unusual but appears to be a property which this area shares with the shelves off central Labrador and Baffin Bay (Keen and Hyndman, 1979; Hinz et al., 1979, this volume). The gravity results are supported by recent crustal refraction measurements in Flemish Pass and Orphan Basin which show them to be underlain by thinned continental crust (Barrett and Keen, 1978). In both areas, normal mantle velocities are measured at about 22 km depth..The main crustal layer is 9 km thick under Orphan Basin and 12 km thick under Flemish Pass, with velocities of 6.3-6.4 km/s. Basement rocks interpreted to be Paleozoic in age, or older, with velocities of 5.4-5.8 km/s overlie the main crustal layer and are 4-6 km thick. The uppermost layers are mainly Late Cretaceous and younger sediments, below which older sediments of inferred Early Cretaceous and Jurassic age are occasionally preserved between basement highs. The Mesozoic and younger sediments are about 5 km thick. The normal mantle and main crustal layer velocities and the apparent absence of a deep crustal layer with a velocity of about 7 km/s such as may characterize the rifted margin off Nova Scotia and the eastern U.S. (Keen et al., 1975; Sheridan et al., 1979) suggest that very little modification of the crust accompanied the observed crustal thinning in these subsided continental regions. This is puzzling, given the history of uplift and subsidence experienced since the Early Cretaceous and the proximity of the area to the early opening of the Labrador Sea. If Orphan Basin and Flemish Pass were produced as part of the development of the continental margins of northeast Newfoundland and now represent the axis of a rift which failed to develop into a mid-ocean ridge, more volcanism and modification of the deep crustal layers producing higher crustal velocities would have been expected. Flemish

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Pass is relatively narrow and lies adjacent to Flemish Cap, a large, shallow fragment. Here less crustal modification associated with the rifting process would be anticipated than in Orphan Basin, which separates the much smaller Orphan Knoll fragment from the shelf by a much larger distance. The lack of significant differences in the crustal structure of the Orphan Basin and Flemish Pass, suggests that crustal thinning is not related in a simple way to horizontal stretching of the crust as McKenzie (1978) has recently suggested. The rifted margin north of the Charlie fracture zone is characterized by a belt of large magnetic anomalies, resembling, but larger than, those along the Nova Scotian margin. Seismic refraction studies have shown a fairly narrow transition from oceanic to continental crust, unlike the Orphan Basin region (Fenwick et al., 1968; Sheridan and Drake, 1968). The continental part of this margin is underlain by rocks of the Paleozoic mobile belt and the accentuated magnetic expression of the ocean-continent boundary is thought to be caused by thick magnetic continental rocks abutting thin magnetic oceanic rocks (Fenwick et al., 1968). Here the upper continental crust consists partly of Paleozoic “oceanic” rocks and also exhibits a deep intermediate layer of 7.4 km/s velocity immediately above the Mdiscontinuity, which lies at a depth of about 40 km (Dainty et al., 1966). The inhomogeneity of the crust and the presence of Paleozoic oceanic layer 3b beneath the continental shelf make seismic velocities poor indicators of the present ocean-continent transition. The results discussed above have been used in recent pre-rift paleogeographic reconstructions of the area (Fig. 3, Le Pichon et al., 1977; Srivastava, 1978; Haworth, in press). These reconstructions place Flemish Cap opposite the Goban Spur and -Orphan Knoll adjacent to Porcupine Bank. Porcupine Bank is an elevated basement feature, partially detached from the adjacent shelf and separated by the deep water indentation of Porcupine Seabight from the Goban Spur. Orphan Knoll is similarly separated by a deep water indentation (the Orphan Basin) from Flemish Cap thus suggesting that Porcupine Seabight may be the eastern equivalent of the Orphan Basin. It similarly appears to be underlain by thinned continental crust (Roberts, 1975; Bailey, 1975). Direct sampling of the rocks of pre-rift age on both sides of the Atlantic has so far given ambiguous results but does confirm the extent of continental rocks. Pautot et al. (1976) and Didier et al. (1977) describe pre-late Hercynian granodiorite dredged from the Goban Spur area in about 4000 m water depth. The granodiorite drilled on Flemish Cap is Precambrian in age and similar to rocks of the Avalon peninsula. On Orphan Knoll limestone pebbles of Devonian age have been collected and may be locally derived (Ruffman and Van Hinte, 1973), and Devonian sediments have also been sampled on Goban Spur (Scrutton, personal communication). “Basement” outcrops on Orphan Knoll and every effort should be made to obtain in situ material from this area. Recent comparisons of continental shelf geology (Bailey, 1975; Haworth,

N.W. AFRICA

i/i-i

.< ./’

” -

~~__

Lti---

_i

Fig. 3. Fit of the continents’around the Atlantic after Lefort and Haworth (1978), Srivastava (1978), and Lefort and Haworth (in press). The Grenville Front is shown as a broad dashed line. Correlations on major fault zones extending from Newfoundland to the shelf of the British Isles are shown as a solid black line. NB = Newfoundland Basin; G.B. = Galicia Bank; F.C. = Flemish Cap; O.K. = Orphan Knoll; G.S. = Goban Spur;and R = Rockall Plateau.

in press) show that the Charlie fracture zone can be traced both eastward and westward into the continents. Its westward extension into Newfoundland follows the northeastern edge of the Precambrian Avalon Platform (Haworth, 1977). Its possible eastward extension onto the Irish Continental margin may be marked- by an east-west trending magnetic anomaly, and a basement flexure which has then been extrapolated toward the Highland Boundary fault on land (Bailey, 1975). The Irish margin correlation is more tenuous than that for the Canadian margin and in order for the correlation to be valid geologically the Irish landfall must be the European equivalent of the northwestern edge of the Avalon Platform. The proposed continuity between the Great Glen fault and the Cabot fault (Wilson, 1962) is not supported by the offshore extension of the Cabot fault (Haworth, in press).

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The Grenville Front apparently cuts across Rockall Plateau (Roberts, 1975) with rocks of the Grenville Province and perhaps of the Paleozoic mobile belt lying beneath the southern part of the plateau. Rockall is underlain by thinned continental crust but, unlike Orphan Basin, the crust there consists of two main crustal layers with velocities of 6.4 and 7.1 km/s (Scrutton, 1972). Perhaps the presence of the intermediate velocity layer under Rockall Plateau reflects the presence of the Paleozoic mobile belt beneath it, similar to the crustal structure northeast of Newfoundland, north of the Avalon zone. Haworth (in press) has compared seismic velocities on the northeast Newfoundland shelf with those southwest of Rockall Plateau, and suggested that their similarity can be attributed to subsided continental crust seaward of the Rockall Plateau. In conclusion, it is worth noting that whereas the thinned continental crust of Orphan Basin and Flemish Pass lies opposite the sharp rifted margins off Porcupine Bank and the Goban Spur in paleogeographic reconstructions, the sharp ocean-continent boundary north of Orphan Basin lies opposite the seemingly diffuse boundary west of Rockall Plateau. Perhaps this is typical of some conjugate margins. The refinements in paleogeography already made possible by correlations in continental geology indicate the need for further studies of this kind. UPLIFT, SUBSIDENCE

AND SEDIMENTATION

The two main episodes in the development of the margins of Atlantic Canada, rifting accompanied by uplift, and drifting associated with subsidence and sedimentation (Vogt and Ostenso, 1967; Sleep, 1971), are clearly seen in the stratigraphic record obtained from seismic reflection data and from numerous deep exploratory wells drilled on the shelves (Gradstein et al., 1975; Grant, 1975; Jansa and Wade, 1975; King and Young, 1977; Cutt and Laving, 1977; Keen, 1979). Major regional unconformities associated with uplift and erosion are observed, overlain by thick sequences of postrifting sediments. These unconformities may separate the basement rocks from the overlying post-rifting sediments. Along the rifted margin off Nova Scotia, Triassic, Jurassic and younger sediments overlie lower Paleozoic sediments (the Meguma) and Devonian granites. The oldest post-rifting sediments are marginal marine evaporites and elastics of Early Jurassic age. Only in restricted deep basins such as the Orpheus sub-basin are Triassic sediments preserved and these are non-marine (Jansa and Wade, 1975). Sedimentation and subsidence appear to have been generally continuous on the Nova Scotian Shelf from the Jurassic to the present, resulting in the deposition of over 10 km of sediment near the shelf edge. The uplift and subsidence history of the Grand Banks is more complex, perhaps because of the numerous times the area experienced uplift, as first Africa, then Iberia and then Europe separated from it. This appears to have

produced a regional unconformity of Early Cretaceous age over most of the Banks, particularly in the southern and eastern areas, and sediments older than mid-Cretaceous are found only in the northeast-trending grabens where Jurassic and older sediments in excess of 4 km are found. In general, the southern and eastern edges of the Grand Banks have experienced less subsidence since the mid-Cretaceous than the shelf edge off Nova Scotia, allowing only about 2 km of sediments to accumulate there. The northern Grand Banks, adjacent Flemish Pass, and the northeast Newfoundland Shelf, where rifting occurred in the early Late Cretaceous, exhibit unconformities of Early to Late Cretaceous age. These unconformities separate Early Cretaceous and older rocks from latest Cre~ceous-Paleocene and younger sediments (Cutt and Laving, 1977). About 3-4 km of postrifting sediments have accumulated since Late Cretaceous time, aided by the extremely rapid Tertiary subsidence observed for this area and for Orphan Knoll. The ages of the unconformities agree with the times of rifting suggested by studies of sea floor spreading. The duration and intensity of uplift can only be qualitatively assessed from the degree of development of the unconformities. The regional unconformity above basement on the Nova Scotian shelf is the best developed and the Early Cretaeeous unconformity on the Grand Banks is also well developed. The Early-Late Cretaceous unconformities on the northeastern shelves are less pronounced and more complex perhaps because there was no single intensive period of rifting between two massive continental plates. There may be more than one important unconformity, one in the Early and one in the Late Cretaceous. The Late Cretaceous unconformity may be directly related to rifting prior to spreading in the Labrador Sea. The significance of the Early Cretaceous unconformity is not known. The presence of Lower Cretaceous and possibly Jurassic marine sediments on this shelf suggest that an epicontinental seaway existed in the region well before rifting occurred (Cutt and Laving, 1977). Salt diapirs of Triassic and Jurassic age are found beneath the Nova Scotian shelf, in the grabens on the Grand Banks, and in fault-bounded basins on the northeast Newfoundl~d shelf (Jansa et al., 1977). Diapirism is also common along the slope off Nova Scotia but it is not known whether these features are salt or shale. These slope diapirs form the sedimentary ridge complex (Emery et al., 1970; Keen et al., 1975; Jansa and Wade, 1975) which extends as far north as the Laurentian Fan (Parsons, 1975). The southeastern edge of this diapiric province is the limit in extent of the flat-lying sediments over oceanic crust to the east which exhibit the deep oceanic seismic markers, “A”, and ‘$” (King et al., 1975). Diapirs are not common along the eastern part of the transform margin or along the margins further north. King (1975) and King and Young (1977) have discussed the post-rifting sediments on the margins in terms of geosyncline development and proposed that the sed~~nts be described as part of a miogeocline consisting of the sedimentary wedge beneath the shelf, and a eugeocline forming the sedi-

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ments on the continental rise. These two elements of the geosyncline exhibit sediments of similar ages but are separated by a zone of paleoslopes in the region of the present continental slope. This zone can be seen on seismic reflection records as a region of confused reflectors dipping down at 5”20” toward the ocean basin. The paleoslopes are “constructional” when formed by progradation of the sediments over the margin, or “destructional” when formed by slumping and erosion. The type of paleoslope depends largely on the balance between rates of subsidence and rates of sedimentation, with shallow water depths and high sedimentation rates leading to the formation of destructional slopes. King and Young (1977) showed that destructional slopes are most common on the older margin off Nova Scotia, infrequent off the eastern margin of the Grand Banks and absent on the margin northeast of Newfoundland. They attribute this to rapid subsidence during the first 50-100 Ma after subsidence starts, leading to a deep quiet depositional environment and the development of constructional paleoslopes. The younger northern margins would be in this stage of development. Slower subsidence later in the history of the margin allows sediments to upbuild to the point where slumping and erosion occurs. This stage has been operative on the Nova Scotian margin perhaps since the Cretaceous and for a shorter period east of the Grand Banks. The concept of paleoslope development is important and requires further study. If destructional paleoslopes are common along the older mature margins, strata are not continuous across the slope region and stratigraphic correlations cannot be made from the shallow to the deep water strata. This points to the need for high resolution seismic studies on the slope and drilling on the rise to obtain the stratigraphy through the eugeocline. Subsidence of “basement”, here defined as the rocks which pre-date rifting, has been investigated by Keen and Keen (1973), Renwick (1973), Gradstein et al. (1975) and Keen (1979) using the biostratigraphic data from deep exploratory wells. Generally subsidence is most rapid early in margin development and slows as time progresses. The subsidence is mainly due to thermal cooling of the lithosphere and other deep-seated processes and to sediment loading as well as to changes in eustatic sea level and in paleo-water-depth (Van Hinte, 1978). The tectonic subsidence due only to thermal contraction or other deepseated processes, can be estimated from the total observed subsidence by correcting for the sediment load and other effects where applicable (Van Hinte, 1978). Keen (1979) has recently shown that the tectonic subsidence of the Nova Scotian margin is linear when plotted against t1’2 (t = time since subsidence began) for the first 80 m.y. after subsidence began. Thereafter the subsidence may be less rapid than is predicted by the t”’ relationship. These results are similar to the behaviour of the oceanic lithosphere as it cools and thickens away from the mid-ocean ridges (Parsons and Sclater, 1977) and have been used as evidence that the margins subside primarily

because they are cooling (Sleep, 1971). If this thermal model is valid, the departure from the linear t”’ relationship after about 80 m.y. suggests that the lithosphere is about 140 km thick under the Nova Scotian margin (Keen, 1979). The crustal thinning of 10 km or more beneath the outer Nova Scotian shelf and slope as observed by seismic refraction measurements and which satisfies gravity measurements across the margins, is necessary for the thick wedge of sediments to have accumulated during margin subsidence. The cause of the observed thinning is not known but is assumed to occur during the rifting stage or early in the subsidence history of the margin perhaps by sub-aerial erosion during uplift or by stretching of the crust under tension, accomplished by normal faulting in the upper crust and plastic flow in the lower crust (Bott, 1971; Artemjev and Artyushkov, 1971; Falvey, 1974). Such processes are especially enigmatic northeast of Newfoundland, where the crust has undergone extensive thinning in Orphan Basin and beneath the outer shelf, but where substantial erosion has not occurred, given the presence of Early Cretaceous and older sediments beneath the shelf. Block faulting in the upper crust and plastic flow in the lower crust under tension could account for the thinning but if this were produced by simple horizontal stretching (McKenzie, 1978), the present continental reconstruction must be drastically revised to allow for this horizontal motion, and furthermore, a simple model of stretching does not explain the similar crustal structures beneath the Orphan Basin and Flemish Pass, discussed earlier. The simple model of McKenzie (1978) does however explain in a general way the anomalously rapid subsidence of this area in the Paleocene but the mode of crustal thinning must have been more complex than in this model. Perhaps, the lithosphere was initially thin or weak in this area due to an earlier tectonic disturbance. This is supported by the presence of Lower Cretaceous basalts on the Labrador Shelf to the north and of Early Cretaceous unconformities suggesting that volcanism and uplift were active at least as early as the Early Cretaceous. If these tectonic events were accompanied perhaps by thinning of the crust in Orphan Basin, it is puzzling that continental break-up finally occurred to the east of Flemish Cap and Orphan Knoll and not between them and the shelf. Geothermal, geochemical and organic maturation studies of the sediments have been carried out with reference to the petroleum potential of the shelf areas (Cassou et al., 1977; Bujak et al., 1977a, b; Robbins and Rhodehamel, 1976). The geothermal gradients measured in deep exploratory wells are about 23°C km-’ in all areas. These are comparable with values measured in the oil-producing Alberta sedimentary basin (Hacquebard, 1977) and much less than the 38°C km-’ gradient found in the Venezuela Basin (KIemme, 1975). Lewis and Hyndman (1976) measured heat flow along two transects across the Nova Scotian margin and found a mean value of 48 mW m-* which is typical of old oceanic crust and of Paleozoic continental platforms (Sclater and Francheteau, 1970). The results exhibited variations along the

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transects of about 20 mW rnT2 and part of this was attributed to the presence of salt diapirs which act to focus the geothermal gradients because of their high conductivities. The geothermal gradients have been used by Robbins and Rhodehamel (1976) to predict that oil and gas should be generated below 3000 m on the Nova Scotian shelf where temperatures exceeding 65°C are found, with optimum depths lying at about 4000 m (93°C). This depth range corresponds to Upper Cretaceous to Jurassic sediments. Their conclusions, however, are based on the assumption that no significant change in the geothermal gradients has taken place in the past. However, if the rifted margins subside mainly because the plate cools, the sediments deposited upon the plate will experience a temperature-time history dependent upon the heat flow through the plate with time, and the rate of subsidence and rapidity with which the sediments are buried under younger strata. Turcotte and Ahem (1977) have quantitatively derived this temperature-time relationship and Keen (1979) applied their methods to obtain the temperature-time history of sediments on the Nova Scotian and Labrador shelves. The results show that temperatures may change rapidly in the first 50 m.y. after deposition and that it is not sufficient to consider only the present temperature distribution in hydrocarbon studies. This analysis of the thermal history showed that Lower Cretaceous and Jurassic sediments on the Nova Scotian Shelf and latest Cretaceous and Lower Paleocene sediments on the Labrador Shelf have experienced temperature conditions favourable to hydrocarbon generation if all other factors were also favourable. However, studies of the degree of organic metamorphism, organic type and organic carbon content of sediments sampled in deep wells indicate that most of the sediments on the shelves between Nova Scotia and Labrador are not promising as sources of oil and gas. In many areas the past thermal history appears to have been favourable but the organic type is not suitable for hydrocarbon generation (Bujak et al., 1977a, b; Cassou et al., 1977). The region northeast of Newfoundland does not conform to these results. Its extremely rapid subsidence in the early Tertiary should, according to the simple cooling plate model, suggest that excessively high temperatures affected the area during the rifting shape. These should have resulted in high geothermal gradients, but there is no evidence that the region is presently hotter than the Labrador Shelf to the north where more normal subsidence rates are observed. Thus the evidence so far available suggests that cooling is not solely responsible for the rapid subsidence of this area but that the other consequences of crustal thinning must also be considered (McKenzie, 1978). More heat flow measurements and accurate geothermal gradient measurements are necessary to confirm this. Further studies of the spatial variations of subsidence, in relation to the past and present geothermal regime should be helpful in determining if the location of deep sedimentary basins along the shelf, such as the Scotian Basin and Georges Basin, which are separated by a basement high, is con-

116

trolled by the temperature distribution in the early rifting history of the margin. Also, deduction of the temperature distribution as a function of time within the plate could be used to define the variations in rheology of the plate with depth and time and thus help to determine whether such basic processes as crustal thinning could have resulted from plastic flow in the early stages of margin formation. EPILOG

The current status of knowledge of the historical development of the continental margin of Atlantic Canada, as discussed within the preceding sections, was summarized in the abstract and will not be reiterated in this section. Instead, a cautionary note will be sounded. The wealth of knowledge available for the Canadian Atlantic margin has, as a result of providing considerable insights into its structure and evolution, posed a large number of questions. These must be answered before the evolution of continental margins can be considered to be understood. Some of the more important questions, particularly as they apply to the Canadian margins are listed below. Can the plate tectonic setting of Precambrian cratons be deduced by geophysical means? What is the history of development of a suture zone, including its thermal history and therefore what evidence of the suture zone will be apparent after l/2 billion years? Can geophysical measurements indicative of modern subduction zones, such as that on the west coast of Canada, be used to identify the location of ancient subduction zones? Would the gravity, seismic and deep electrical properties observed over present convergent margins be expected to persist over the times required for comparisons with their ancient analogs? Is the volcanism in the Avalon zone related to island arc or rift tectonism (or both)? Why is the Avalon zone as wide as it appears to be, if it represents an old island arc complex ? Are the Avalon volcanic and metasedimentary rock velocities separable, enabling these units to be defined offshore? What is the variation in crustal structure across the western boundary of the Avalon zone? What is the nature of basement beneath the Me&ma zone? Is the basement comparable to rocks of the Avalon zone? What is its extent? Were the Meguma sedimentary strata deposited as a fan on oceanic or continental crust? If the boundary between the Avalon and Meguma zones lies beneath the Grand Banks, why did the transform margin now situated south of the Grand Banks form where it did, and not along this major structural boundary?

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What is the relationship between the Paleozoic oceanic rocks exposed Newfoundland and modern oceanic crust?

in

Is it possible to construct a universally valid model of the cross-section of an island arc (e.g. Mitchell and Reading, 1971)? Can the shallow 7+ km/s velocities be tied directly to ophiolites by refraction at the Bay of Islands or Hare Bay? What is the mechanism for emplacement of ophiolites? Can the roots of the ophiolites be recognized by their geophysical character? What is the relationship between the deep and shallow 7+ km/s velocities? Are the lateral limits of the deep 7+ km/s confined between the ancient margins (why is this layer not observed at lines 6 and 7, Fig. 2)? What is the late Paleozoic plate tectonic history? What is the relationship between the granites in Nova Scotia and possible Devonian subduction and how are they related to the granites in Newfoundland? Can the validity of paleomagnetic results which imply Devonian collision between Atlantic Canada and South America be tested by geological or geophysical means? What was the depositional environment of the Dunnage melange and hence what are the implications of this possible trench deposit for Silurian plate tectonics in Newfoundland? How did the transition between Ordovician closure (compressional tectonics) and the initiation of Triassic rifting (tensional tectonics) take place? What do the presence of open marine, Late Carboniferous fauna, with North American affinities imply for paleogeographic reconstructions at that time? What tectonic forces produced the northwest trending folds in Carboniferous sediments off northern Newfoundland in contrast with the northeast trending folds found further south? What provides control over the pattern of continental breakup and what is the mechanism of separation? Why are the basalts of Triassic age so far from the location of the oceancontinent boundary? Are hot spots responsible for breakup at triple points, as Burke and Dewey suggest, or do the continents separate with orthogonal rift and transform segments? Is tensional necking and thinning of the crust largely responsible for the observed crustal thinning at rifted margins thus perhaps explaining why less thinning is observed at transform margins? What is the role of sub-aerial erosion in crustal thinning and how does the amount of erosion on various segments of the transform margin compare with that at the rifted margin? How do we distinguish between the oldest post-rifting sediments and crystalline basement (e.g. Meguma) when V > 5 km/s for both? What can we learn about the mechanism of separation from the Orphan Basin region ? How can similar crustal thicknesses in Flemish Pass and Orphan Basin be compatible with stretching? Why is there little apparent modification of the

crust in Orphan Basin when we might expect volcanism and higher crustal velocities? Why did breakup occur east of Orphan Knoll and Flemish Cap in view of stretching on its west side? Why was there no substantial erosion in Orphan Basin compared with the Scotian margin before subsidence occurred? Can crustal extension explain the very rapid Paleocene subsidence at Orphan Knoll? Can listric faults, indicative of crustal stretching, be defined in Orphan Basin? Where does the ocean--continent

boundary lie?

What causes the gravity and magnetic edge anomalies and do rocks of high density and high magnetization lie at the shelf edge and significantly contribute to these? Do rifted margins have 7+ km/s transitional crust separating continental crust from magnetic oceanic crust as a source of the magnetic edge anomalies? Can deep drilling provide the final proof that oceanic crust lies beneath the Quiet Magnetic Zone in the northwestern Atlantic, the Newfoundland Ridge and the Spur Ridge? How much of the Newfoundland Basin is oceanic and how do we bridge the gap between Iberia and the Grand Banks in paleogeographic reconstructions ? What is the nature of the crust west of Iberia and Galicia Bank? Can further continental structural continuities be recognized by which to constrain and refine paleogeographic reconstructions? How can the history of subsidence and sedimentation at continental margins best be modelled? Under what conditions does the linear t”2 subsidence relationship break down? Can we improve the models for subsiding continental margins by comparing theoretical results with observed geothermal gradients and measures of organic metamorphism of the sediments sampled in deep exploratory wells? Can a realistic model for the development of the sedimentary wedge at a subsiding margin be produced, which includes temporal variations in rheology? What rheological model is most reasonable in modelling subsidence? Can drilling and high resolution seismic data on the slope validate the constructional and destructional paleoslope model? From an examination of this list of questions, two conclusions become evident. First, the depth of knowledge of this margin compared with most other allows a rather complete description of its geological and geophysical characteristics. Despite this, many of the important processes involved in its evolution remain enigmatic or uncertain. Second, examination of the ancient margin in comparison with the modern may yield many important clues to their evolution. However, little more than superficial comparisons between the two have so far been made, particularly with respect to their deep crustal structure. While such comparisons are difficult, studies of this kind may provide the best means of solving some of the problems raised above.

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