The Newfoundland–Iberia conjugate rifted margins

In this chapter 10.1 Introduction

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10.2 Geologic setting

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10.3 Continental basement rocks 345 10.4 Rift phases

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Late Triassic–Early Jurassic extension 347 Epeirogenic subsidence phase 350 Late Jurassic to Early Cretaceous extension 350 Transition to seafloor spreading 352 Late-stage magmatism 356

10.5 Plate kinematic reconstructions

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10.6 Geological/geophysical constraints on the Early Cretaceous transition from rifting to seafloor spreading

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Rift segmentation 359 Southern segment 359 Central segment 361 Boundary between the central and northern segments Northern segment 364 Significance of transition zone structure 366

10.7 Extensional models and melt supply

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Relevant numerical models 368 Conceptual models for asymmetry of the transition zones

10.8 End of rifting and post-rift sedimentary history 10.9 Conclusions Acknowledgments References

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10 The Newfoundland–Iberia conjugate rifted margins Brian E. Tucholke,* Robert B. Whitmarsh { *Department of Geology and Geophysics, Woods Hole Oceanographic Institution, Woods Hole, Massachusetts, USA { School of Ocean and Earth Science, National Oceanography Centre, Southampton, United Kingdom

10.1 Introduction How continental rifting and breakup evolve into seafloor spreading is a fundamental, yet poorly understood, part of the plate-tectonic cycle. Processes of continental separation involving crustal thinning, magmatism, faulting, possible uplift, and eventual thermal subsidence profoundly affect the architecture of continental margins and leave important records of their operation. On volcanic margins, this architecture is commonly obscured by thick igneous sequences, and on nonvolcanic margins it is often deeply buried beneath sedimentary sequences and/ or deformed evaporites. The opposing margins of the Grand Banks and Iberia, however, formed in a non-volcanic rift where such deposits minimally conceal deeper structures. For this reason, the Newfoundland–Iberia rift has been the subject of extensive geophysical study and scientific drilling. Although investigations have concentrated on the Iberia margin, geophysical surveys and scientific drilling conducted in 2000–2003 seaward of the Grand Banks also provide important information that help us assess the evolution of the whole rift (we use the term ‘rift’ throughout this chapter to indicate the broad extensional zone that lay between North America and Iberia as the conjugate Newfoundland and Iberia margins developed). In this chapter, we summarise the basement structure and the history of extension in the rift from initial rifting in Triassic time to the onset of seafloor spreading in the Early Cretaceous. At the end of the chapter, we also briefly summarise the principal features of the sedimentary record in the rift.

10.2 Geologic setting As discussed here, the Newfoundland–Iberia rift (Fig. 10.1) was bounded at its southern end by the Newfoundland–Gibraltar Fracture Zone along the southwestern edge of the Grand Banks. This fracture zone comprised a transform

Phanerozoic Passive Margins, Cratonic Basins and Global Tectonic Maps DOI:10.1016/B978-0-444-56357-6.00009-3 343 Copyright © 2012 by Elsevier B.V. All rights of reproduction in any form reserved.

Phanerozoic Passive Margins, Cratonic Basins and Global Tectonic Maps

344 Figure 10.1 Bathymetry of the North Atlantic encompassing the Newfoundland and Iberia margins, from IOC, IHO and BODC (2003). Contour interval is 1000 m. Precambrian and Paleozoic terranes shown on the Newfoundland margin are adapted from Keen et al. (1990) and on the Iberia margin from Pinheiro et al. (1996). AB, Alentejo Basin; ABR, Azores–Biscay Rise; AM, Armorican Massif (off map); CA, Collector magnetic anomaly; CIZ, Central Iberian Zone; GB, Galicia Bank; GIB, Gibraltar; LB, Lusitanian Basin; MTR, Madeira-Tore Rise; OMZ, Ossa Morena Zone; SPZ, South Portuguese Zone; TS, Tore Seamount.

Phanerozoic Passive Margins, Cratonic Basins and Global Tectonic Maps

margin between Africa and North America-Eurasia during the early opening of the central Atlantic from about Bajocian time (180 Ma) to the Early Cretaceous. (The time scales of Channell et al. (1995) and Gradstein et al. (1995) are used throughout this chapter). On the Newfoundland side, the fracture zone was subsequently covered by volcanics that formed the Southeast Newfoundland Ridge (SENR) in the middle Cretaceous (Tucholke and Ludwig, 1982). On the Iberia side, the fracture zone broadly corresponds today to the Azores–Gibraltar plate boundary; this boundary was significantly modified by plate motion between Iberia and Africa after the Newfoundland–Iberia rift opened. The northern end of the rift is at the eastern tip of Flemish Cap on the Newfoundland margin and at the northwestern edge of Galicia Bank on the Iberia margin (Fig. 10.2). The northern margin of Galicia Bank was modified by compression and underthrusting in Paleocene–Eocene time (Boillot and Malod, 1988). The northeast margin of Flemish Cap is conjugate to Goban Spur (Fig. 10.1); this conjugate pair of margins separated during Early Cretaceous rifting after the initial opening of the Newfoundland–Iberia rift (Sibuet et al., 2004).

10.3 Continental basement rocks Basement of the Newfoundland–Iberia rift is an assemblage of Precambrian to Paleozoic rocks that was accreted to the eastern margin of the Laurentian craton during the closing of the Paleozoic Iapetus and Rheic oceans. These rocks formed the Appalachian orogen along the present Atlantic Canadian margin where they are ordered into three zones consisting, from northwest to southeast, of the Dunnage, Gander, and Avalon zones (Fig. 10.1) (Williams and Hatcher, 1982, 1983). The Avalon zone contains thick Upper Precambrian sedimentary and volcanic sequences with overlying Cambrian shales, and it occupies eastern Newfoundland and most of the Grand Banks platform. At the southeastern edge of the Grand Banks, another zone (Meguma) lies south of the Avalon and is separated from it by a prominent magnetic anomaly termed the Collector Anomaly (Figs. 10.1 and 10.2). The Meguma zone consists of thick CambrianOrdovician graywackes and shales overlying Late Proterozoic gneissic basement. These terranes were accreted during the Taconian, Salinic, Acadian, and Alleghenian orogenies (Early to Middle Ordovician, Early Silurian, SilurianDevonian, and Carboniferous-Permian, respectively) (Percival et al., 2004; van Staal et al., 1998; Waldron and van Staal, 2001). Large granitic plutons were emplaced in the existing terranes during the Acadian orogeny. The Gander-Avalon and Avalon-Meguma boundaries are steep ductile shears and brittle faults that imply assembly by transcurrent motion rather than by subduction/obduction (Keen et al., 1990). Structural fabrics within the Avalon and Meguma terranes are well defined by gravity and magnetic anomalies, and in many places they appear to have controlled subsequent development of rift basins within the Grand Banks (Welsink et al., 1989).

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346 Figure 10.2 Reconstruction of the Newfoundland–Iberia rift at chron M0 (North America plate is fixed), based on the reconstruction pole of Srivastava et al. (2000) and adapted from Tucholke et al. (2004). Tectonic and other data are compiled from numerous sources. Plate boundary in the centre of the rift is shown by a heavy black line. The diamond line off Iberia shows the minimum seaward limit of continental crust (poorly constrained in the southern rift); the minimum limit off Newfoundland lies along Flemish hinge and the hinge line at the outer edge of Salar–Bonnition basin. The seaward transition zone (STZ; centre, light gray) extends landward approximately to anomaly M3 and continued to develop following M0 until latest Aptian time. The landward transition zones (LTZ) reach landward from
M3 to the minimum seaward limits of continental crust. Thrust faults along the northern margin of Galicia Bank were formed during postrift, Paleocene–Eocene compression within the Iberia plate (Boillot and Malod, 1988). Bold dotted lines locate seismic reflection profiles illustrated in Fig. 10.4; heavier sections of the LG 12 and SCREECH 2 lines show the locations of profiles in Figs. 10.5 and 10.6, respectively. DSDP/ODP drill sites are shown by black dots and labels. AF, Aveiro fault; A. P., abyssal plain; BI, Berlengas islands; BMT., basement; NF., Newfoundland; NZF, Nazare´ fault; PTF, Porto-Tomar fault; SMT, Seamount; TF, Tagus fault; T.Z., transfer zone; C.A., Collector Anomaly.

Phanerozoic Passive Margins, Cratonic Basins and Global Tectonic Maps

Basement on the Iberia margin consists of rocks deformed during the Variscan orogeny in mid-Devonian to Carboniferous time. Until the opening of the Bay of Biscay, these rocks were continuous with the Armorican Massif of northwestern France (Fig. 10.1). The Variscan Massif in Iberia consists mainly of folded, thrusted, and metamorphosed Precambrian and Paleozoic rocks. These rocks were extensively intruded by large granitoid batholiths during and after the Variscan continent-continent collision that accompanied the closure of the Paleotethys Ocean (Pinheiro et al., 1996). The Variscan Massif is divided into a number of distinct zones with NW-oriented boundaries. Various authors have attempted to trace these zones offshore, and even to correlate them with the Grand Banks margin, by using plutonic and metamorphic rocks from dredge hauls and cores (e.g., Capdevila and Mougenot, 1988) and the distribution of magnetic susceptibility (Silva et al., 2000). The results are tentative and even conflicting. A late Variscan fault system developed in Iberia in response to east–west compressive stress (Ribeiro et al., 1990). The main faults were sinistral, NNE to ENE strike-slip faults with subsidiary conjugate dextral NNW to NW faults. In addition, two major, roughly N–S, strike-slip faults (the Porto-Tomar Fault and a fault separating the Berlengas islands from the mainland, Fig. 10.2) were also reactivated (Ribeiro and Silva, 1997). Basement thus possessed an anisotropic fabric of faults, folds, and thrusts before extension began in the Mesozoic. The faults, when reactivated, strongly controlled basin geometry, facies distributions, sites of salt structures, and locations of extensional and compressional faults (Wilson et al., 1989).

10.4 Rift phases Late Triassic–Early Jurassic extension The Newfoundland and Iberia margins experienced two main phases of extension (Fig. 10.3). The first phase was concentrated in Late Triassic to earliest Jurassic time, coincident with rifting that led to Middle Jurassic (Bajocian) opening of the central Atlantic immediately to the south of the Newfoundland-Iberia rift. Large evaporitic rift basins formed primarily within the Grand Banks but also in Iberia (Fig. 10.2). With the exception of rifts that developed into deep-water basins because of later rifting along the outer Newfoundland and Iberia margins, all the Triassic Basins are within continental crust of normal thickness (
30 km). On the Grand Banks, structural trends of the major basins mostly follow preexisting basement heterogeneities (Silva et al., 2000; Welsink et al., 1989). The Whale Basins, Horseshoe and Jeanne d’Arc Basins, as well as basins to the east (Fig. 10.2) align with magnetic-anomaly trends that largely reflect contrasts in basement magnetisation and thus in pre-existing structure. Numerous transfer faults are roughly orthogonal to the predominantly NE–SW-oriented basins, and they commonly offset the basins or limit their extents along-strike. Multiple strands of transfer faults are often organised into broad-scale transfer zones that also seem to follow the pre-existing basement structure (Welsink et al., 1989).

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Figure 10.3 Summary diagram of events in the Newfoundland–Iberia rift. Data are arranged by conjugatemargin pairs from south to north in the rift. The third pair of columns in the figure correlates with the transect of drill sites shown in Fig. 10.2. The row labeled Profile identifies seismic reflection profiles illustrated in Fig. 10.4. Gray tones show principal periods of rifting; the barred gray areas represent uncertain Late Triassic and Late Jurassic rifting on the Iberia margin (Murillas et al., 1990) and post-Aptian detachment of basin fill related to early extension in the Labrador Sea on the northern Newfoundland margin (e.g., Grant et al., 1988). Vertically lined intervals indicate unconformities, most of which probably developed because of tectonic uplift associated with rifting. “E” at bottom indicates evaporites deposited in shallow rift basins during the Late Triassic to Early Jurassic. Minor magmatism occurred at scattered locations during both Late Triassic to Early Jurassic and latest Jurassic to Early Cretaceous rifting. The SENR at the southern margin of Newfoundland Basin was a major locus of volcanism in the Barremian–Aptian (Tucholke and Ludwig, 1982). Numbers by magmatic occurrences in the four right-hand columns indicate ODP drill sites. Open arrows labeled with associated M-series magnetic anomalies, earliest seafloor spreading proposed by various authors. Solid arrows, interpreted establishment of normal seafloor spreading, coincident in time with the breakup unconformity. S–B, Salar–Bonnition Basin; NB, Newfoundland Basin.

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The Grand Banks basins accumulated continental to shallow-marine, siliciclastic ‘red-bed’ sediments during Carnian-Norian time; these deposits included and were succeeded by evaporites in the earliest Jurassic (Hettangian-Sinemurian) (Jansa and Wade, 1975; Jansa et al., 1980). Igneous activity was very limited during this rift phase. It is documented primarily by a few diabase dykes and mafic basalts in Newfoundland and Nova Scotia and by rare occurrences of basalts in offshore wells (Fig. 10.3) (Sinclair, 1988). Offshore Iberia, there is some evidence for lithospheric extension before Triassic time. A metagabbro sampled at ODP Site 1067 in the southern Iberia Abyssal Plain (Fig. 10.2) contains earliest Permian (270  3 Ma) zircons, and it is interpreted as having been underplated at the base of the crust during early extension on this margin (Manatschal et al., 2001). Onshore, the Iberian Basin in central-eastern Spain was also undergoing extension at this time (Arche and Lopez, 1996). Other significant basins developed along the west Iberia margin, including (from south to north) the Lusitanian Basin, Porto Basin, and Galicia Interior Basin (Fig. 10.2) (Pinheiro et al., 1996). As outlined below, some of these basins began to form in the late Triassic, but their early history is not well known compared to that of the Grand Banks basins because they have been less explored for hydrocarbons. Lusitanian Basin is the principal basin on the shallow Iberia margin. It extends
250 km south to north, but its northern, offshore extent is uncertain (Wilson et al., 1989). Sediments along its NNE-trending axis of maximum subsidence are up to 4 km thick, contained between Hercynian basement to the east and basement horsts offshore to the west. The Triassic rift phase is represented by an unconformity-bounded sedimentary sequence which consists of Upper Triassic fluviatile sandstones and Upper Triassic to Hettangian dolomites and evaporites (Wilson et al., 1989). These sediments were deposited in grabens and half grabens (Leinfelder and Wilson, 1998) during a period of relatively diffuse extension. Porto Basin underlies the continental slope and outer shelf farther to the north. It is relatively narrow (
50 km) and is bounded to the east by the Porto-Tomar Fault and to the west by a horst block. Wells in southern Porto Basin have penetrated Triassic to Hettangian shallow-water sediments and evaporites that may be associated with rifting (Murillas et al., 1990). Porto Basin may represent the northernmost sub-basin of Lusitanian Basin. Galicia Interior Basin lies between Porto Basin and the basement highs of Galicia Bank (Fig. 10.2). The seismic stratigraphy of the basin was described by Murillas et al. (1990) but no well information has been published. Murillas et al. identified seismic units using seismic facies criteria and attempted to date them by correlation to sampling sites or wells in the adjacent Porto Basin. They presented several arguments for Triassic–Liassic initial rifting even though reflections that might indicate sediments of pre-Late Jurassic age are rarely observed. They

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also proposed that Lower and Middle Jurassic sediments were deposited in a tectonically quiet environment, but they could not exclude a phase of Late Jurassic extension. In contrast to this interpretation, a more recent seismic profile across the northern basin (along 42 400 N) shows no seismic interval interpreted to be older than Early Cretaceous (Valanginian; Pe´rez-Gussinye´ et al., 2003), even where the sediments are
6 km thick. Pe´rez-Gussinye´ et al.’s basal sedimentary sequence (Formation 5) overlies a strongly block-faulted basement that they interpret as crystalline continental crust. Their results do not support, but neither do they conclusively refute, the contention of Murillas et al. (1990) that the basin first extended in Triassic time.

Epeirogenic subsidence phase Extension largely ceased on the Grand Banks during Early and Middle Jurassic time, and the platform subsided while it accumulated shallow-water carbonates (Fig. 10.3) (Tankard and Welsink, 1987). A similar lack of extension characterised the broader rift during this period. Samples from dredges (Boillot et al., 1979; Dupeuble et al., 1987) and ODP Sites 639, 901, and 1065 (Boillot et al., 1987; Sawyer et al., 1994; Whitmarsh et al., 1998) indicate that thin Tithonian–Berriasian shallow-water carbonates were deposited widely on Galicia Bank, and Tithonian shallow-water siliciclastic mudrocks accumulated on the now deeply subsided continental crust south of Galicia Bank beneath the southern Iberia Abyssal Plain. These sediments document a period of quiescence before the final phase of rifting that led to continental breakup. An exception to this quiescence was a phase of faulting and basin subsidence that occurred in Lusitanian Basin south of Nazare´ Fault (Fig. 10.3) (Carvalho et al., 2005; Rasmussen et al., 1998). Igneous activity during this epeirogenic phase was minor and was restricted to the southern end of the rift (Fig. 10.3; Pinheiro et al., 1996; Sinclair, 1988).

Late Jurassic to Early Cretaceous extension A second phase of the Grand Banks–Iberia extension occurred in the Late Jurassic through Early Cretaceous (Fig. 10.3). This phase was more prolonged than the first phase, and it encompassed breakup of continental crust from latest Jurassic to Hauterivian–Barremian time, first in the southern part of the rift and then in the northern, Galicia Bank–Flemish Cap segment. The latter part of this extensional phase also included exhumation of apparently sub-continental mantle to as late as earliest Albian time off Iberia (Whitmarsh and Wallace, 2001) as well as exhumation of mantle of presently uncertain affinity off Newfoundland, possibly as late as the end of Aptian time (Tucholke et al., 2007). Except for an apparently plume-related magmatic event that affected the southern two-thirds of the rift in Barremian–Aptian time (see ‘Late-stage magmatism’ later in this section), the rifting seems to have been accompanied by only minor, scattered magmatism (Fig. 10.3). In the northern half of the rift, there is direct evidence for intrusion of small gabbroic bodies into peridotite basement at ODP sites 637, 1070, and 1277.

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On the Grand Banks, extension appears to have persisted with variable intensity from Oxfordian to Aptian time, with the most intense rifting occurring in the Kimmeridgian to Tithonian (Driscoll et al., 1995; Tankard and Welsink, 1987). Unconformities that developed in the Jeanne d’Arc Basin in late Barremian/early Aptian and late Aptian time are interpreted by Driscoll et al. (1995) to have been formed during the final rift events on the main Grand Banks, and they suggested that these events immediately preceded the onset of seafloor spreading to the east (see next section on ‘Transition to seafloor spreading’ for discussion of the onset of normal seafloor spreading). Beginning in late Barremian time, NE–SW–directed extension apparently associated with early rifting of the Labrador Sea also affected the northern margin of the Grand Banks (Tankard and Welsink, 1987; Welsink et al., 1989). This extension cut across pre-existing normal faults at high angles, reactivated transfer faults as normal faults in the northern Jeanne d’Arc and Flemish Pass Basins, and detached the basin fill; the extension appears to have persisted at least to the end of Albian time (Fig. 10.3) (Enachescu, 1988; Grant et al., 1988). Offshore southwestern Iberia, the NE-trending Alentejo Basin underlies the continental slope (Fig. 10.1) (Mougenot et al., 1979; Mougenot, 1988). It contains up to 3 s (two-way time) of sediment over Jurassic acoustic basement and it appears to have developed during the Late Jurassic rift phase. The relationship between this basin and Lusitanian Basin to the north is unknown. In the Lusitanian Basin, the strongest extension occurred during Oxfordian to early Kimmeridgian time, and it was followed by more limited rifting in Tithonian to Barremian time (Carvalho et al., 2005; Rasmussen et al., 1998; Wilson et al., 1989). Middle Oxfordian-Berriasian sediments show dramatic thickness variations; the southern part of the main basin divided into the Bombarral salt-withdrawal sub-basin and into a number of possibly short-lived (1–2 m.y.; Leinfelder and Wilson, 1998), rapidly subsiding half-grabens that were bounded by major N–S normal faults (Carvalho et al., 2005). Subsidence analyses at 26 wells confirm these observations and show that the cumulative preCretaceous extension in Lusitanian Basin was greater south of the Nazare´ fault (b ¼ 1.17) than north of it (b ¼ 1.03) (Stapel et al., 1996). NE–SW faults also were present, and at least some experienced Late Jurassic normal, and probably transtensional, motion (Wilson et al., 1989). Farther to the north, in southern Porto Basin, wells show a Jurassic section up to 1500 m thick that accumulated as the basin opened (Murillas et al., 1990). Galicia Interior Basin to the west is interpreted by Pe´rez-Gussinye´ et al. (2003) to have started rifting during the Early Cretaceous (Valanginian), with some continuing extension on the eastern margin of Galicia Bank in the Hauterivian to late Aptian. The latter phase culminated at the same time that extension ceased on the deep, western margin of Galicia Bank. On this western margin, ODP Leg 103 drilling results were interpreted to indicate the end of continental

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rifting close to the Aptian/Albian boundary, and this was assumed to be coeval with the initiation of seafloor spreading (Boillot and Winterer, 1988). This event postdates by
10 m.y. the synrift emplacement of a 122 Ma (late Barremian) chlorite schist and gabbro (Scha¨rer et al., 2000). There is no direct evidence of pre-Cretaceous extension on the western Galicia Bank margin, although Mauffret and Montadert (1987) suggested that the margin was affected by Late Jurassic (Oxfordian-Kimmeridgian) extension and they mapped inferred pre-Lower Cretaceous sediments up to 2.5 km thick there. From analysis of drilling results and seismic stratigraphy of the basal sedimentary sections on the deep-water Newfoundland and Iberia margins, Tucholke et al. (2007) concluded that Late Jurassic to Early Cretaceous extension in the rift was concentrated in two episodes that culminated near the Berriasian–Valanginian and Hauterivian–Barremian boundaries (Fig. 10.3). They also proposed that the entire system experienced elevated extensional stress normal to the rift axis (i.e., ‘in-plane tensile stress’) for at least 30 m.y. up until the Aptian–Albian boundary, at which time normal seafloor spreading became well established and inplane tensile stress was significantly reduced. According to this hypothesis, this rift-wide event at the end of Aptian time correlates on the Iberia margin to a prominent, deep-basin reflection previously identified as the ‘breakup unconformity’ (e.g., Boillot and Winterer, 1988; Mauffret and Montadert, 1987) and on the Newfoundland margin to the widespread and highly reflective U reflection (Tucholke et al., 1989; Tucholke et al., 2007) (Figs. 10.4–10.6).

Transition to seafloor spreading The question of when, where, and how continental rifting evolved into normal seafloor spreading in this non-volcanic rift is complex, and answering it requires an understanding of the underlying and changing geological processes. The geological record of the transition is manifested in three zones across each pair of conjugate segments of the Newfoundland and Iberia margins: (1) extended continental crust, (2) a transitional zone, and (3) a presumed normal oceanic crust (i.e., ‘Penrose ophiolite’-type crust consisting of a volcanic layer overlying sheeted dykes and a gabbroic layer 3; Penrose Conference Participants, 1972). Unfortunately, without more extensive basement sampling than is presently available, it is difficult to be certain where normal oceanic crust first became a consistent and permanent feature in the rift. We summarise existing constraints on each of the three zones below. The zone of extended, seaward-thinning continental crust is closest to the margins and contains a record of late-stage rifting leading up to breakup of the continental crust. The thinning typically occurs over a distance <50–70 km, with crust at the seaward edge of the zone being
5–7 km thick, and usually it is accompanied by tilted normal fault blocks. Under the continental slopes and rises near the rift margins, larger fault blocks are 10–20 km wide over continental crust that is 20–30 km thick (Fig. 10.4). Farther seaward, wherever crustal blocks

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are 7 km thick, normal faults imaged on seismic profiles are clearly listric, closer together, and may even penetrate the mantle (Iberia margin, Fig. 10.4B and C; see also Fig. 10.4A in Chian et al., 1999). Seaward of this is a wide transition zone (TZ) where basement exhibits considerable variation in tectonic structure, development of magnetic anomalies, and seismic velocities (see Section 10.6). The genesis of the TZ is disputed, but

Figure 10.4 (See legend next page)

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Figure 10.4 Simplified interpretations of conjugate multichannel seismic-reflection (MCS) profiles across the Newfoundland and Iberia margins. All profile locations are shown in Fig. 10.2. Interpreted syn- and post-rift sediments are shaded. The reflections labeled U off Newfoundland and BU off Iberia mark the deep-water breakup unconformity at the Aptian/Albian boundary. Vertical exaggeration is 4.5 in (A) and is
10 at the basement surface in (B–D). (A) Conjugate profiles (in depth) across Flemish Cap and Galicia Bank, juxtaposed at magnetic anomaly
M0. The Flemish Cap profile is from Funck et al. (2003) and Hopper et al. (2004). The Galicia Bank profiles are from Whitmarsh et al. (1996) and Pe´rez-Gussinye´ et al. (2003), with the S reflection from profile GP101 (Reston et al., 1996); note that there is an offset between the GP101 and ISE 17 profiles (see location in Fig. 10.2). Profiles in B, C, and D are in reflection time and are juxtaposed at anomaly
M1. (B) Conjugate profiles across southern Flemish Cap and southern Galicia Bank along the ODP drilling transect near the boundary between the central and northern rift segments (see also Figs. 10.5 and 10.6). At left is an interpretation of SCREECH line 2MCS (Shillington et al., 2004) with location of ODP Site 1276 and lithology at the bottom of the hole; ODP Site 1277, not illustrated, is on a peridotite ridge about 5 km east of anomaly M1 (see Fig. 10.6). At right is a composite seismic section (Sonne 16, JOIDES Resolution, Lusigal 12, OC 103) along the conjugate Iberia drilling transect, adapted from ODP Leg 173 Scientific Party (1998); drill sites are numbered. PR, peridotite ridge. (C) A pair of approximately conjugate profiles in the central rift segment. At left, Conrad MCS profile NB1 (Tucholke et al., 1989) about 150 km south of SCREECH Line 2MCS shows another view of Newfoundland basement structure and the overlying U reflection. At right is MCS line IAM9 about 50 km south of Lusigal 12 off Iberia, adapted from Pickup et al. (1996). PR, peridotite ridge. (D) Conjugate profiles in the southern rift segment. At left is Conrad MCS line NB19 in southern Newfoundland Basin (Tucholke et al., 1989). At right is MCS profile Lusitanie 86 over Tagus Abyssal Plain, from Mauffret et al. (1989). Note the marked asymmetries in basement depth and roughness between the Newfoundland and Iberia sides of the rift on all profiles.

Figure 10.5 Segment of Lusigal 12 multichannel seismic reflection profile along the Newfoundland–Iberia drilling transect on the Iberia margin showing lithology of basement cored at ODP drill sites or, where basement was not reached (Sites 901 and 1065), at the bottom of the hole. Location in Fig. 10.2. The breakup unconformity (BU) can be traced to DSDP Site 398 where it is dated as late Aptian to early Albian. PR3 and PR4 are peridotite ridges that have been traced along the Iberia margin in previous studies. ODP Site 1070 is on a peridotite ridge about 8 km west of the left edge of the profile. The positions of magnetic anomalies M5 (projected) and M3 are as identified by Russell and Whitmarsh (2003).

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Figure 10.6 Segment of SCREECH2 multichannel reflection profile through ODP Sites 1276 and 1277 in the Newfoundland Basin (Shillington et al., 2004). Location in Fig. 10.2. Locations of magnetic anomalies M3 and M1 are indicated. The U reflection is a high-amplitude, basinwide reflection that dates to the latest Aptian–earliest Albian and is equivalent to the breakup unconformity on the Iberia margin. At Site 1276, sediments at and below the reflection are intruded by Albian and early Cenomanian diabase sills (Tucholke et al., 2004; Hart and Blusztajn, 2006). Site 1277 drilled 38 m of serpentinised peridotite basement beneath
58 m of allochthonous basaltic, gabbroic, serpentinite, and sedimentary debris. Sedimentary sequences are from Tucholke et al. (2004): A – Dark green-gray to black shales (uppermost Aptian to lowermost Albian and older); B – Dark green-gray to black shales (Albian–Turonian); C – Red-brown to gray shales, sandstones, and mudstones (Turonian–middle Eocene); D – Gray-green mudstones forming an interval of sediment waves (middle Eocene to lower Oligocene); E – Probable Oligocene to Pliocene fan deposits (uncored) and F – Quaternary turbidites. Reflection AU marks the initiation of strong abyssal circulation in the North Atlantic.

resolving its origin is key to understanding evolution of the rift. We distinguish here, principally on the basis of evidence from the southern Iberia Abyssal Plain, a landward TZ (LTZ, older than anomaly
M3–M5) and a seaward TZ (STZ; anomaly
M3 to latest Aptian) (Figs. 10.2 and 10.4). The LTZ exhibits lowamplitude, linear magnetic anomalies that parallel the margin for up to tens of kilometres. Whether or not these anomalies were generated by seafloor spreading is a topic of debate. Srivastava et al. (2000) proposed that the LTZ basement on both the Newfoundland and Iberia margins is oceanic crust, on the basis of their picks of M-series magnetic anomalies as old as M20 (Tithonian) (Fig. 10.3). In contrast, in the central segment of the rift on the Iberia margin Whitmarsh and Miles (1995), Discovery 215 Working Group (1998), Dean et al. (2000), and Russell and Whitmarsh (2003) contend that the LTZ has no identifiable M-series magnetic anomalies older than M3–M5 (BarremianHauterivian). They and other workers conclude that the LTZ off Iberia is underlain by exhumed and variably serpentinised sub-continental mantle, termed the zone of exhumed continental mantle (ZECM) by Whitmarsh et al. (2001). Along most of the Newfoundland margin, a prominent magnetic low that probably is anomaly M3 is observed (Shillington et al., 2004; Tucholke et al., 1989). Younger anomalies M2–M0 also are identifiable adjacent to the high-amplitude

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‘J Anomaly’ on both sides of the rift (Sibuet et al., 2004; Srivastava et al., 2000). Given the general agreement that anomalies
M3 and younger can be identified and correlated along strike on both margins, it appears that this basement, which we associate with the STZ, has the magnetic signature of normal oceanic crust. Seismic observations also indicate that often the STZ has the velocity structure of normal oceanic crust (Dean et al., 2000; Lau et al., 2003; Nunes, 2002). There are, however, several problems with this interpretation. First, where the STZ basement has been drilled, that is near anomaly M1 in the Newfoundland Basin (Site 1277; Figs. 10.4 and 10.6) and near M2 in the Iberia Abyssal Plain (Site 1070; Fig. 10.4), apparently autochthonous basement is serpentinised peridotite like the exhumed mantle in the LTZ (Tucholke et al., 2004; Whitmarsh et al., 1998). Although the Site 1070 peridotites are intruded by E-MORB gabbroic dykes, they do not appear to have the geochemical signature of normal oceanic mantle. Second, the peridotite basement at Site 1070 was not exhumed until Albian time, at least 14 m.y. after it initially accreted at the anomaly-M2 plate boundary (Whitmarsh and Wallace, 2001). This exhumation brought rocks to the seafloor from at least
5–6 km depth and occurred
125 km ‘off-axis’ (assuming a half spreading-rate of 9 mm/a). Such large-scale off-axis extension is not known to occur on the flanks of a normal spreading ridge. Finally, if the STZ basement were normal oceanic crust, then the seafloor spreading that created it must have occurred simultaneously with the Hauterivian to Aptian continental rifting discussed above. This would mean that correlations commonly made between rift-basin unconformities and events such as ‘rift-onset’ or ‘breakup’ are incorrect and that development of such unconformities must have other explanations. If, on the other hand, we consider that normal seafloor spreading that produced crust approaching the Penrose model was not established until the end of Aptian time, and that up to this time the entire rift system experienced elevated inplane tensile stress (Tucholke et al., 2007), then the above problems are overcome. In particular, this timing for the initiation of normal seafloor spreading is consistent with the age of the previously defined breakup unconformity off Iberia (Boillot and Winterer, 1988; Mauffret and Montadert, 1987) and the correlative U reflection off the Grand Banks (Tucholke et al., 1989; Tucholke etal., 2007). Thus, we consider the possibility that the STZ actually extends
150–160 km seaward beyond anomaly
M3 to where basement has a latest Aptian– earliest Albian age. In Sections 9.5–9.7, we review the character of the LTZ/ZECM and STZ more fully by considering plate-kinematic constraints on breakup, by examining geological and geophysical data that constrain TZ structure and composition, and by considering models for its origin.

Late-stage magmatism In Barremian to early Aptian time (anomaly
M4 to younger than M0), there was extensive magmatism at the Mid-Atlantic Ridge (MAR) axis where it impinged on the southern end of the Newfoundland–Iberia rift at the Newfoundland–Gibraltar Fracture Zone (Tucholke and Ludwig, 1982). The

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volcanism appears to have been associated with a relatively short-lived mantle plume. On the North American plate, it constructed the massive basement edifice of the SENR (Figs. 10.1 and 10.2), as well as large, contiguous ridges that extended southwest along the MAR (J Anomaly Ridge, JAR) and northeast into the Newfoundland Basin (Tucholke et al., 1989). On the Iberia plate, the volcanism constructed a basement edifice conjugate to the SENR/JAR that comprises the Madeira-Tore Rise (MTR) (Fig. 10.1). The volcanic construction and associated thermal uplift elevated parts of the JAR, and probably parts of the SENR and MTR, to sea level (Tucholke and Ludwig, 1982), but the volcanism waned rapidly in Aptian time. Emplacement of melt from this plume in the southern two-thirds of the Newfoundland–Iberia rift most likely accounts for the presence of the high-amplitude magnetic ‘J Anomaly’, which is associated with anomalous magnetisation between about anomaly M1 and M0 (Rabinowitz et al., 1978). It may also have contributed to increasingly oceanic character of basement in this portion of the rift from anomaly
M4 to beyond anomaly M0, as suggested by some seismic refraction studies (e.g., Dean et al., 2000). A somewhat later magmatic event affected the TZ in the Newfoundland Basin but appears to have had no impact on the TZ of the Iberia margin. Drilling at ODP Site 1276 in the central Newfoundland Basin LTZ (Figs. 10.2, 9.4B, 10.6) penetrated diabase sills that were emplaced into lowermost Albian sediments above basement during Albian to early Cenomanian time (Tucholke et al., 2004; Hart and Blusztajn, 2006). Although the source of the magmatism is uncertain, it may be related to volcanism that formed the Newfoundland Seamounts. The age of the seamount volcanism is constrained by only a single isotopic date (97.7  1.5 Ma) from Scruncheon Seamount in the centre of the chain (Sullivan and Keen, 1977), which indicates volcanism there in the earliest Cenomanian. The Newfoundland Seamounts appear to be part of a trace of hotspot volcanism that subsequently formed the Milne Seamounts and the conjugate Azores–Biscay Rise near and at the mid-ocean ridge axis (Fig. 10.1) (Louden et al., 2004; Whitmarsh et al., 1982). The plate boundary did not migrate westward across this thermal anomaly until Campanian to early Eocene time, and the anomaly appears not to have propagated across the MAR into the Iberia plate. Thus, it would not have affected the Iberia TZ. This difference in Late Cretaceous magmatism may help explain asymmetries in sediment-unloaded basement depth between the conjugate Newfoundland and Iberia TZs, as discussed in ‘Conceptual models for asymmetry of the transition zones’, Section 10.7.

10.5 Plate kinematic reconstructions Accurate Early Cretaceous plate reconstruction between the Grand Banks and Iberia is effectively limited to the time of anomalies M0 to
M3 (Fig. 10.2) because the Cretaceous long normal-polarity interval immediately followed

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anomaly M0. There is also a lack of significant fracture zones to document plate motions at this time. Identification of anomaly M0 itself can be confusing because it is frequently and erroneously equated with the magnetic ‘J anomaly’, a high-amplitude anomaly generated around chrons M1–M0 during the magmatic event noted above (Rabinowitz et al., 1978; Tucholke and Ludwig, 1982). The high amplitude of the J anomaly diminishes markedly to the north near the southern margins of Flemish Cap (45 46 N) and Galicia Bank (41 42 N). Srivastava et al. (2000) proposed that M0 extends further north off Galicia Bank and Flemish Cap and Sibuet et al. (2004) have proposed that M0 continues in a northeast direction into the Bay of Biscay, although its presence in both these locations is uncertain (see ‘Northern segment’ in Section 10.6). Srivastava et al.’s (2000) reconstruction at M0 time provides the best current estimate of the fit between the Grand Banks and Iberia. A convincing constraint in this reconstruction is the close congruence of the estimated conjugate Iberia– Grand Banks–Eurasia RRR triple junctions based on magnetic anomalies (Sibuet et al., 2004). Further support comes from a measurement of uppermost mantle seismic anisotropy in the LTZ/ZECM off Iberia (Cole et al., 2002). There, the direction of fast seismic wave propagation can be interpreted as an estimate of the motion of Iberia relative to the Bay of Biscay triple junction if one postulates shear or flow of the uppermost mantle in this direction during extension of the LTZ/ZECM. The fast-propagation direction is 143 , which agrees with the direction of extension proposed by Sibuet et al. (2004) within 10 . Magnetic anomaly charts of the Grand Banks and west Iberia margins (Miles et al., 1996; Verhoef et al., 1996) reveal linear margin-parallel anomalies landward of M0 that individually can be followed for several tens of kilometres. However, it is debatable to what extent these margin-parallel magnetic lineations can be equated with seafloor-spreading anomalies and therefore can be used in plate reconstructions. Some proposed identifications require disproportionate changes in spreading rate and remanent magnetisation, and they are in locations where they do not overlie basement with the velocity structure of normal oceanic crust (Srivastava et al., 2000). Such identifications can have other, perhaps more plausible, geological explanations. In the southern Iberia Abyssal Plain, Russell and Whitmarsh (2003) identified M0–M3, and M5 only locally, whereas Srivastava et al. (2000) proposed anomalies M0–M17. Farther south, in the Tagus Abyssal Plain, Pinheiro et al. (1992) proposed that anomalies M0–M11 are present, whereas Srivastava et al. (2000) proposed anomalies M0–M20. The most straightforward conclusion that can be drawn from these studies is that there may be an increasing ‘oceanic’ component (i.e., an increasing proportion of melt) in LTZ/ZECM basement in a seaward direction (Russell and Whitmarsh, 2003; Whitmarsh and Wallace, 2001). Attempts to close Iberia and the Grand Banks at pre-M0 time lead to overlaps of continental crust of Flemish Cap and Galicia Bank in the northern part of the rift

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(Fig. 10.2). Only if allowance is made for continental extension, particularly in the Galicia Interior Basin and west of Flemish Cap, can a tighter, but still approximate, closure be obtained (Srivastava and Verhoef, 1992; Sibuet et al., 2007). Positions of continental crust in the M0 reconstruction (Fig. 10.2) clearly indicate that LTZ/ZECM basement was being formed in the southern part of the rift while continental crust was still being extended in the northern part, consistent with the extension history of basins in Iberia (e.g., Murillas et al., 1990). Because there are no magnetic anomalies that have time significance in the 38 m.y.-long Cretaceous long normal-polarity interval between anomalies M0 and 34, and there is only sketchy information on fracture-zone trends in this interval, it is not possible to document Grand Banks–Iberia relative plate motions accurately for Aptian to early Campanian time. Thus, it is unknown whether the latestage Aptian rifting and the postulated normal seafloor spreading at the beginning of Albian time were associated with any significant change in plate motion.

10.6 Geological/geophysical constraints on the Early Cretaceous transition from rifting to seafloor spreading Rift segmentation It is convenient to divide the Newfoundland–Iberia rift into three pairs of conjugate segments on the basis of morphological and structural differences, and of differences in the timing of continental rifting noted above (Fig. 10.2): (1) a southern pair of segments that extend from the SENR north to the Newfoundland Seamounts and from Gorringe Bank to Estremadura Spur, (2) a central pair of segments that extend north of the first pair to the southern margins of Flemish Cap and Galicia Bank, and (3) a northern pair of segments that include Flemish Cap and Galicia Bank. We review the geological and geophysical constraints on the structure of these segments below.

Southern segment In this segment, the southern Salar–Bonnition Basin lies along the western margin of the rift (Fig. 10.2). This basin is bounded to the west and east by basement hinge zones that face one another, and it contains apparently thick evaporites of probable Triassic age (Figs. 10.4C and D) (Austin et al., 1989). The hinge zone at the east edge of the basin marks the seaward limit of clearly continental crust (Keen and deVoogd, 1988). The only other major rift basins (Whale, South Whale and Horseshoe) within this rift segment are more than 200 km to the west within the shallow Grand Banks platform. The intervening southeasternmost Grand Banks is underlain by essentially unstretched,
28-km-thick

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continental crust (Reid, 1994) that has been in a sub-aerial to shallow-shelf environment throughout the Mesozoic and Cenozoic (Jansa and Wade, 1975). The LTZ seaward of Salar–Bonnition Basin is about 100–140 km wide. Reid (1994) reported seismic refraction results along a transect extending from the shallow Grand Banks margin seaward into the LTZ at 43 450 N. The change in margin structure is abrupt, with most crustal thinning occurring over 30 km across Salar–Bonnition Basin. Immediately seaward in the LTZ, basement includes a
4 km-thick layer of velocity 7.2–7.6 km/s that overlies mantle and is beneath 2–3 km of lower-velocity material; these layers extend seaward for 50–60 km. Beyond this zone, but still within the LTZ, an anomalously thin (as little as 3 km) basement layer with low velocity (4.5–5 km/s) and high gradient overlies mantle with velocities of 7.7–7.9 km/s. In the conjugate LTZ under the central Tagus Abyssal Plain off Iberia, Pinheiro et al. (1992) observed a thin (
2 km) low-velocity layer overlying a layer in which velocity increases with depth from 7.6 to 7.9 km/s. The layers with velocities in the low- to mid-7 km/s range, both here and off Newfoundland, have been interpreted as serpentinised peridotite, with the degree of serpentinisation decreasing with depth (e.g., Pinheiro et al., 1992; Reid, 1994). The interpretation of Pinheiro et al. provided a plausible and more satisfactory interpretation of two refraction lines along the southern margin of the Tagus Abyssal Plain which Purdy (1975) was unable to interpret as a conventional reversed pair. It now appears that one line was shot mainly over the LTZ and the other line mainly over thinned continental crust. Various extents and ages (Jurassic and older) of ‘oceanic crust’ in the LTZ beneath the Tagus Abyssal Plain have been inferred but without any firm constraints. Several sub-parallel and relatively low-amplitude, linear, margin-parallel magnetic anomalies lie landward of anomaly M0, some of which intersect the refraction profile of Pinheiro et al. (1992). Pinheiro et al. modelled these anomalies and suggested that seafloor spreading started at anomaly M11 (late Valanginian), although the velocity structure interpreted as being caused by serpentinised peridotite basement there does not change across this interval. They also noted that farther landward, relatively high remanent magnetisations are required in magnetic models to fit the data, and they speculated that this may be explained by dykes intruded in thinned continental crust at about the time of breakup. Their results contrast with the suggestion of Mauffret et al. (1989) that an abandoned Late Jurassic (chron M21) spreading ridge underlies the abyssal plain. Farther east Mauffret et al. (1989) also noted a landwarddipping reflection in basement (Fig. 10.4D), analogous to a reflection noted by Keen and deVoogd (1988) which coincides with the hinge at the eastern edge of the conjugate Salar–Bonnition Basin, and they proposed that it marks the ocean-continent boundary. Pinheiro et al. (1992), however, suggested that the seaward edge of continental crust is
40 km farther seaward on the basis of the seismic refraction results reported by Purdy (1975).

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Magnetic anomalies in the conjugate LTZ off Newfoundland are low-amplitude (<100 nT) and are difficult to correlate or to relate to any seafloor-spreading model. The oldest seafloor-spreading anomaly that can be identified with some certainty is M3 at the seaward edge of the Newfoundland LTZ (Fig. 10.4) (Tucholke et al., 1989). Basement at the J anomaly (
M1–M0) rises more than two kilometres to form prominent isochron-parallel ridges near the SENR (Fig. 10.4D), but the amplitude of the ridges decreases northward within the rift segment (Tucholke et al., 1989). Similar ridges comprising the MTR extend north toward Tore Seamount on the Iberia side (Figs. 10.1 and 10.2). Both sets of ridges are attributed to Barremian–early Aptian magmatism noted in ‘Latestage magmatism’, Section 10.4.

Central segment In contrast to the southern segment, the thick continental crust of the central rift segment on the Newfoundland margin exhibits deep basins in the rift margins, including the Jeanne d’Arc, Carson, and Salar–Bonnition Basins. All these basins are filled with Cretaceous and older sediments. As in the southern segment, the Salar–Bonnition Basin is bounded by opposing hinge zones, with the seaward hinge marking the limit of clearly continental crust. Transfer zones accommodated differential extension within all the rift basins, but it is uncertain whether or how they reach seaward into the LTZ. Keen et al. (1977) postulated that the Newfoundland Seamounts formed above a pre-existing structural discontinuity across the LTZ, possibly the extension of a transfer fault such as the Avalon transfer zone (Fig. 10.2). If such a fault exists, its counterpart on the Iberia margin could be the Nazare´ Fault on the north side of Estremadura Spur, with the conjugate faults defining the boundary between the southern and central rift segments. The LTZ in the central rift segment on both margins is up to
170 km wide (Fig. 10.4C). On the Newfoundland side, Lau et al. (2003) reported that continental crust thins rapidly seaward beneath Carson Basin and is
7 km thick near the seaward hinge of Salar–Bonnition Basin. Farther seaward, they interpreted the velocity structure of the LTZ to indicate a
50-km-wide zone of thinned continental crust (3–5 km) over a thick (>5 km) layer of serpentinised mantle (7.5–8.0 km/s). This is followed seaward by an
80-km-wide zone with a thin (2 km) low-velocity layer over likely serpentinised mantle before basement with velocity structure more like that of oceanic crust is reached in the STZ seaward of anomaly
M3. Within the LTZ, but farther north, Srivastava et al. (2000) reported thin crust over probable serpentinised mantle (2 km of 4.4–5.5 km/s over 2 km of 7.1–7.3 km/s) from seismic refraction station Nwnb, much like the results of Lau et al. (2003). In the LTZ of the conjugate Iberia margin (IAM9, Fig. 10.4C), a 2.4-km-thick upper basement layer with a P-wave velocity of 4.5–7.0 km/s and a high velocity gradient (
1 s–1) merges into a lower layer 4 km thick with velocities of

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7.6 km/s and a low velocity gradient (<0.2 s–1) (Dean et al., 2000). Moho reflections are weak or absent. The upper basement layer has unusually low reflectivity (Pickup et al., 1996). The reflection character and velocity structure probably reflect mantle serpentinisation that decreases with depth (Chian et al., 1999; Dean et al., 2000; Pickup et al., 1996). The LTZ here has been interpreted as a ZECM (Whitmarsh et al., 2001). Low-amplitude, margin-parallel magnetic anomalies indicate that basement magnetisations in the Iberia LTZ/ZECM are typically much lower than those of oceanic basement, while spectral properties of the anomalies suggest that most source bodies lie not at the top of acoustic basement but up to 6.5 km deeper (Russell and Whitmarsh, 2003). The source bodies are interpreted to be isolated, north-south-elongated mafic intrusions that were emplaced in exhumed mantle and that increase in volume seaward. The oldest magnetic anomalies identified are M3, and locally M5, at the seaward edge of the LTZ (Russell and Whitmarsh, 2003). On the Newfoundland side, M3 is the oldest anomaly that can be identified with some confidence at the seaward edge of the LTZ.

Boundary between the central and northern segments The boundary between the central and northern rift segments off Iberia is here defined as the limit of continental crust along the southern margin of Galicia Bank (NW–SE diamond pattern, Fig. 10.2). In detail, this boundary probably is irregular, consisting of a series of rift-parallel ridges that plunge to the south beneath, and eventually die out under, the southern Iberia Abyssal Plain; in this location they merge with the deep, smoother LTZ/ZECM (Russell and Whitmarsh, 2003). This zone is the locus of a transect of conjugate Newfoundland–Iberia drill sites that straddles the segment boundary. Because the zone is both complex and has been extensively studied by drilling and geophysical methods, we review it separately here. The Iberia side of the transect has been investigated more intensively (Fig. 10.5), and this is where a threefold zonation of thinned continental crust, a ZECM and a crust formed by earliest seafloor spreading were first proposed (Whitmarsh et al., 2001). The tectonic history of this area and a summary of the drilling results are described in greater detail by Whitmarsh and Manatschal in Chapter 9 (this volume). Off Iberia the continental crust tapers from 7 km to zero thickness westward along the transect over >50 km (Dean et al., 2000; Whitmarsh et al., 2000). Two deep-water fault blocks drilled by ODP in this zone (Sites 901 and 1065) are capped by upper Tithonian sediments (Figs. 10.4B and 10.6). At Site 901, the benthic foraminifers indicate neritic (<200 m) depths; the sedimentary facies are similar at both Sites 901 and 1065, indicating that both sites overlie continental crust. No subsidence history can be deduced from the cores because the relevant post-rift interval was not sampled adequately. Three other ODP boreholes, drilled <1400 m apart (Sites 900, 1067, and 1068), were located

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over a basement high (Hobby High) near where continental crust tapers to zero thickness. Among the cored rocks were lower-crustal metagabbro and amphibolites that were underplated in Late/post-Hercynian time; these rocks experienced ductile deformation at 161 Ma before being exhumed by rifting at <137 Ma. The main phase of continental extension here lasted from Tithonian (147 Ma) to at least late Berriasian time (<137 Ma); during this period the crust was stretched by
34 km, yielding an average extension rate of >3.5 mm/a (Whitmarsh et al., 2001). Seaward, the LTZ/ZECM is
80 km wide along the drilling transect (Figs. 10.4B and 9.5). Morphologically the acoustic basement may be divided into a western region of N-S-trending basement ridges and an eastern deep region (8–9 km below sea level) of relatively low-relief basement (Plate 1 of Chapter 9); both regions narrow northward. Highly serpentinised peridotite was cored at ODP sites 897, 899, 1068, and 1070 (Abe, 2001; He´bert et al., 2001; Whitmarsh and Wallace, 2001). Primary-phase chemistry and clinopyroxene trace-element compositions indicate that the ZECM is heterogeneous peridotite irregularly depleted by <10% partial melting and permeated by mafic melts (He´bert et al., 2001). Trace-element compositions are more like sub-continental or suprasubduction-zone mantle than abyssal (oceanic) mantle (Abe, 2001; He´bert et al., 2001). Serpentinisation began, at least locally, before exhumation of the peridotite to the seafloor (Skelton and Valley, 2000). The average extension rate between the first exhumation of mantle rocks at Hobby High and anomaly M3 was at least 7.2 mm/a (Whitmarsh et al., 2001). The LTZ/ZECM begins to merge into the STZ at about anomaly M3. Here, highamplitude, linear, margin-parallel magnetic anomalies are consistent with systematic accretion of basement at 10–14 mm/a, beginning locally no earlier than M5(R) (late Hauterivian) (Russell and Whitmarsh, 2003). Seaward of M3, the velocity structure gradually changes over
50 km to a structure characteristic of oceanic crust (Dean et al., 2000). The only drill site in this zone is ODP Site 1070, located near anomaly M2 (middle Barremian)
20 km seaward of a margin-parallel, basement peridotite ridge where ODP Site 897 was drilled (Whitmarsh and Wallace, 2001). This basement ridge can be traced more than 120 km northward along the western margin of Galicia Bank to near ODP Site 637 (Beslier et al., 1993). At Site 1070, upper Aptian sediments were found overlying depleted, serpentinised peridotite with gabbro veins and a small gabbro pegmatite body derived from an E-MORB source (Abe, 2001; He´bert et al., 2001). Although the gabbro pegmatite was emplaced at 127  4 Ma (Hauterivian-Barremian) (Beard et al., 2002), it was not exhumed through the plagioclase blocking temperature (
150 C, or
5–6 km sub-seafloor at a geothermal gradient of 25 –30 C/km) until 110.3  1.1 Ma (earliest Albian) (Whitmarsh and Wallace, 2001). Thus strong extension affected this basement near the time of the Aptian/Albian boundary. No basaltic rocks were encountered at the site. The mineralogy and the elemental and isotopic geochemistry of the gabbros

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are consistent with derivation from an enriched mantle source, possibly sub-continental mantle, and they mitigate against a depleted N-MORB asthenospheric source. Along the conjugate, Newfoundland half of the drilling transect, a long seismic refraction profile (Nunes, 2002) shows that continental crust thins seaward to a few kilometres thick at the Flemish Hinge (Fig. 10.4B). Farther seaward, in contrast to the rift segments to the south, the simplest velocity model that fits the data suggests the presence of thin (4–5 km) and laterally uniform oceanic crust (Nunes, 2002). This velocity structure extends both through the LTZ and seaward into the STZ at and beyond anomalies M3–M0, even though there is a dramatic seaward change from relatively smooth to high-amplitude basement topography beginning near anomaly M3 (Fig. 10.6). Nowhere along the profile is there a significant layer that shows the
7.2–7.8 km/s velocities commonly attributed to moderately serpentinised mantle. Poisson’s ratios of 0.28  0.01, derived from converted crustal shear-wave arrivals, are also observed throughout the LTZ and STZ (Nunes, 2002). Such ratios are consistent with basaltic oceanic crust but are outside the range normally associated with serpentinites. These results are puzzling because ODP Site 1277, drilled near anomaly M1 in the STZ (Fig. 10.6), cored serpentinised peridotite that is interpreted to be autochthonous basement (Tucholke et al., 2004). If the serpentinised peridotite is representative of a much wider region than the borehole, it has three important implications: (1) Poisson’s ratios cannot be relied upon to discriminate variably serpentinised peridotite from other types of crust, (2) the observed velocity structure across the TZ, although it is interpreted by normal convention to identify thin oceanic crust, actually can be created in serpentinised mantle, and (3) because the velocity structures of basement at Site 1276 (in the LTZ) and Site 1277 (STZ) are very similar, there is a strong possibility that the presently uncored basement at Site 1276 is serpentinised peridotite.

Northern segment The northernmost pair of conjugate rift segments differs markedly from the segments to the south because rifted continental crust under Galicia Bank extends much farther west and because the TZ is significantly narrower. In the northern part of the segment off Newfoundland (Fig. 10.4A), anomaly M3 as picked by Srivastava et al. (2000) coincides with thinned continental crust (Funck et al., 2003), so the LTZ is missing and the last extension of continental crust there postdates early Barremian time. At the conjugate location off Iberia, no anomalies are known in the age range of M3–M5, so the LTZ appears to be missing there too. Toward the south end of the segment, the LTZ widens to
60 km (Figs. 10.2 and 10.4B). Flemish Cap on the Newfoundland side of this segment has continental crust of normal thickness (
30 km) (Funck et al., 2003), but it is separated from the Grand Banks platform by thinned continental crust under Flemish Pass Basin and Flemish Cap Graben (Enachescu, 1987). The southeast margin of Flemish

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Cap has a well defined, seaward-dipping hinge zone (Figs. 10.2 and 10.4A) and there is no major rift basin seaward of the hinge. Seismic refraction results show that the crust of Flemish Cap thins over a distance of about 70 km to
5 km thick at the outer edge of the hinge. Seaward, a thin (
1–3 km), 55-km-wide basement layer with reflection character and velocity like that of oceanic crust reaches 25 km beyond anomaly M0, but this crust overlies a
3–5 km-thick 7.6–7.9 km/s layer that is interpreted to be serpentinised peridotite (Funck et al., 2003; Hopper et al., 2004). Farther seaward, a more normal oceanic crustal structure may be present, although it is not well constrained by the refraction data. On the conjugate Galicia Bank margin, a series of wide-angle seismic lines along an east-west transect showed that the continental crust thins seaward from
17 to
5 km over a distance of 80 km (Fig. 10.4A; Whitmarsh et al., 1996). This margin yielded the first clear evidence of continental fault blocks off west Iberia (e.g., Groupe Galice, 1979), probably in part because the segment was uplifted in Tertiary time and thus has been relatively sediment-starved. Weakly metamorphosed Paleozoic sediments were dredged from the most seaward fault block, about 30 km east of the peridotite ridge drilled at ODP Site 637 (Figs. 10.2 and 10.4A), and this block is clearly continental (Mamet et al., 1991). Subsidence history of continental fault blocks drilled during ODP Leg 103 (Sites 638–641) (Figs. 10.2 and 10.4A) was presented by Moullade et al. (1988), who inferred a 25 m.y. duration of extension. Whitmarsh et al. (1996) fitted these subsidence observations by assuming minor late Jurassic extension (b ¼ 1.08), followed by major Berriasian–Aptian extension (b ¼ 4.0), probably with a decreasing rate of extension over time. An important feature on many reflection profiles, unique to this segment, is the S reflection, which has been interpreted to be a low-angle detachment surface extending from mid-crust in the east to the base of the crust in the west (Fig. 10.4A) (Reston, 1996; Reston et al., 1996). The reflection is continuous for 50 km east of the peridotite ridge. Unger and Sawyer (1999) suggested that S is possibly only one of a series of three such anastomosing reflections in this area. Slip on the S detachment(s) probably occurred at the same time that extension was exhuming the peridotite ridge (Barremian-Aptian) (Boillot et al., 1989; Fe´raud et al., 1988). Two lines of evidence suggest that oceanic crust appears within
20 km west of the peridotite ridge on the Galicia margin (Fig. 10.4A). First, magnetic anomaly amplitudes on an E–W deep-towed magnetometer profile across the peridotite ridge are almost ten times greater west of the peridotite ridge than to the east. The anomalies were modelled by Sibuet et al. (1995), incorporating basement relief from a coincident multichannel seismic reflection profile, by juxtaposing a weakly magnetised peridotite ridge and landward-thickening thinned continental crust against strongly and uniformly magnetised crust to the west, which suggested that crust west of the peridotite ridge is oceanic. Subsequently,

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Srivastava et al. (2000) proposed that anomaly M0 is present along the same magnetometer profile. Second, the refraction results of Whitmarsh et al. (1996) show that the crust immediately west of the peridotite ridge is only 2.5–3.5 km thick but thickens westward to a normal oceanic thickness and velocity structure, as is observed under the southern Iberia abyssal plain seaward of anomaly
M3 (Dean et al., 2000). In contrast to this evidence, the M0 reversal was detected in a core of syn-rift sediments from ODP Site 641 (Fig. 10.4), some 40 m below (i.e.,
4 m.y. earlier than) the breakup unconformity at the Aptian/Albian boundary (Ogg, 1988). This implies that normal seafloor spreading did not commence between Galicia Bank and Flemish Cap until after the time of chron M0 and that Srivastava et al.’s (2000) anomaly M0, if correctly identified, is not developed in normal oceanic crust.

Significance of transition zone structure The principal characteristics of almost all basement in the LTZ (i.e., basement older than
M3–M5), including the more tightly defined LTZ/ZECM off Iberia, are (1) a landward edge abutting continental crust that has been block-faulted and thinned to about 5–7 km, (2) a lower layer, at least several kilometres thick, with velocities in the range of
7.0–7.8 km/s and low gradients, and (3) a
2–4 km-thick, 2.5–4.0 km/s upper layer with variable but commonly high velocity gradients. It has been postulated that once old cratonic crust is thinned by b ¼ 3–4, fluids can reach the underlying mantle, leading to serpentinisation and a weak rheology that promotes tectonic extension and mantle exhumation (Pe´rez-Gussinye´ and Reston, 2001). This is consistent with the observed amount of crustal thinning at the landward edges of the LTZ, and it is also consistent with the interpretation of the lower layer in the LTZs as serpentinised peridotite, with the degree of alteration decreasing downward into normal mantle. Thus, a dominant characteristic of the LTZ/ZECM appears to be exhumation of mantle to shallow levels (if not to the seafloor), which implies extreme tectonic extension and a very limited melt supply. In the 2–4 km thick upper layer of LTZ/ZECM basement, the wide range of velocities and reflection characteristics in seismic profiles (e.g., fault patterns, reflectivity) could be explained in several ways. They could represent real changes in composition, for example lateral changes between continental rocks and serpentinised mantle, both of which could contain varying amounts of intruded melt. Alternately, the layer could be entirely exhumed mantle peridotite that exhibits strong three-dimensional variations in the degree and depth distribution of serpentinisation. Despite the similar velocity structures of LTZ basement in the Newfoundland and Iberia conjugates, there are marked asymmetries in the two sides of the rift

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(Fig. 10.4). Most notably, the whole Newfoundland LTZ on average is 1 km shallower than that off Iberia. In addition, Newfoundland LTZ basement tends to have lower relief than that off Iberia. The latter observation implies that there was less brittle extension in the Newfoundland LTZ, although the basement roughness there may have been reduced by post-rift magmatism (see ‘Conceptual models for asymmetry of the transition zones’, Section 10.7). The STZ (basement younger than
M3) appears to be more complex than the LTZ. It is clear that mantle was exhumed in this zone because peridotite basement was cored in the only holes drilled there to date (ODP Site 1277 near anomaly M1 and ODP Site 1070 near M2); at Site 1070 this mantle appears to have continental affinity (Whitmarsh and Wallace, 2001). Paradoxically, however, the seismic refraction observations cited above show that the velocity structure becomes more like that of normal oceanic crust within this zone. Thus, it seems likely that a substantial part of the STZ basement, at least where sampled on basement highs, consists of exhumed (possibly sub-continental) mantle. Like the LTZ, it appears that the STZ was increasingly (albeit irregularly) intruded by melt as the rift widened, but unlike the LTZ the volume of melt was enough to give rise to magnetic anomalies that usually are considered characteristic of seafloor spreading. Depending on the volumes and distribution of the melt, it may be difficult to discriminate such intruded mantle from ‘normal’ oceanic crust in seismic refraction profiles, which typically average structure over several kilometres along the profiles. It is difficult to pinpoint the source of the melt. On one hand, the thermal evolution of the rift leads eventually to a melt-rich system of seafloor spreading (see ‘Relevant numerical models’, Section 10.7). On the other hand, from about anomaly M4 to a time somewhat younger than M0, the southern part of the rift was affected by melt input from the plume that formed the SENR and MTR volcanic edifices. Judging from the extent of the high-amplitude J anomaly, this effect reached at least to the northern edge of the central rift segment, but it is unclear whether it affected the northern segment. In order to resolve whether the introduction of melt into the rift increased continuously or was the product of the melt event centered on the J anomaly, high-quality reflection and refraction data will be required across the J anomaly and onto Albian-age crust, particularly in the southern part of the rift. Where is the seaward edge of the STZ, and to what extent might it be geologically and geophysically definable? It appears that the rift system experienced elevated in-plane tensile stress until the end of the Aptian (Tucholke et al., 2007), as manifested by extensional tectonism that exhumed mantle from depths as great as 20 km, and more than 100 km into the plate interior, up until this time (e.g., Fe´raud et al., 1988; Whitmarsh and Wallace, 2001). The classically defined breakup unconformity on the Iberia margin (BU in Figs. 10.4 and 10.5; Boillot and Winterer, 1988; Mauffret and Montadert, 1987), equivalent to the U reflection on the Newfoundland margin, is a major stratigraphic marker within the rift

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that dates to the Aptian/Albian boundary. We equate the development of this horizon with a release of in-plane tensile stress that accompanied the first well established seafloor spreading at the end of Aptian time. According to this interpretation, the BU and U reflections do not mark the breakup of continental crust; instead, they correlate with much later (>10 m.y.) breaching of probably sub-continental mantle lithosphere. If the release of tensile stress was caused by persistent influx of large volumes of melt at the plate boundary, we expect that the transition from the STZ to Penrose-type oceanic crust should be manifested in two ways. First, we expect consistent development of a relatively normal oceanic igneous sequence (volcanics over sheeted dykes and gabbros) seaward of the STZ. Second, the transition may also be detectable geochemically, for example by a compositional change from melt derived from sub-continental lithospheric mantle to melt derived from oceanic asthenospheric mantle. Both of these predictions can be tested by appropriate seismic and drilling experiments.

10.7 Extensional models and melt supply Relevant numerical models Generally, numerical models of the formation of rifted margins investigate the dynamic extension of the lithosphere and/or the processes that generate melt during extension. Ideally, any model specifically for the Grand Banks and/ or Iberia margins should address a large number of constraints based on observations. First-order constraints are the two-phase extensional history (Fig. 10.3 and Section 10.4) and the limited distribution and volume of melt within the TZs. Second-order constraints include the generally deeper basement on the Iberia side and the abundant evidence of low-angle detachment surfaces there. Few authors have attempted to address these problems for this margin pair. Tett and Sawyer (1996) computed dynamic models of the two-phase lithospheric extension. The most successful models, which incorporated symmetric rifting, predicted a total lack of magmatism and suggested that extension was significantly greater during the second, Late Jurassic–Early Cretaceous phase of rifting, which they assumed to last 35 m.y. The models did not predict the observation that continental crust is highly thinned over a broad area. Keen et al. (1994), working on the assumption that decompression melting of the mantle occurred during extension, computed the volume and distribution of melt over the Grand Banks from a grid of subsidence values estimated from bathymetry and sediment thickness. However, they showed (their Fig. 10.8) that most observed occurrences of Triassic to Cretaceous basaltic rocks in industry boreholes were not where predicted by their model, and they inferred that other processes may also be important. More recently, evidence has been adduced that the major rift phase off Iberia was much shorter than the 35 m.y. assumed by Tett and Sawyer (1996) or inferred from drill cores (10 m.y.) (see ‘Boundary between the central and

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northern segments’, Section 10.6). Wilson et al. (2001) suggested that rifting lasted even less than 5 m.y. Many models predict that during such a short rift episode, decompression melting of the asthenosphere will lead to the production of significant volumes of melt, contrary to observations off the Grand Banks and Iberia. Bowling and Harry (2001) attempted to solve this problem by suggesting a model in which decompression melting, because it was controlled by necking of a homogeneous lithosphere, was delayed until the last 10% of the rift episode. The model successfully produces broad rift margins underlain by highly attenuated continental crust. Unfortunately, the model does not differentiate, in time or space, between continental break-up and the onset of seafloor spreading (i.e., the model does not include a broad TZ or ZECM). Off Iberia and the Grand Banks, these clearly are quite separate events. In their model, a small volume of melt appears during the last stages of continental extension but only a few tens of kilometres landward of where seafloor spreading starts. Where the TZ/ZECM is narrow, as off Galicia Bank, the lack of significant syn-rift melt was modelled successfully by Whitmarsh et al. (1996) only by assuming that extension lasted at least 25 m.y. Along the Iberia drilling transect (Fig. 10.4B), an alternative solution was proposed by Minshull et al. (2001). They explored and rejected several reasons for the small volume of melt in the ZECM (e.g., lateral heat loss to the adjacent continental lithosphere, anomalously low mantle potential-temperature at the time of break-up, and depth-dependent stretching). They noted that because the thermal structure will affect mantle flow, the ZECM may result from a transitional stage between continental break-up and steady-state seafloor spreading in which the mantle flow pattern is quite different. They speculated that during the formation of the ZECM the thermal and viscosity structures evolved only slowly towards those of a mid-ocean ridge. The final stage of this process rapidly focused flow, once sufficient melting had been initiated, because of a large viscosity increase in the mantle as water was removed in the melt. On the basis of our present geological and geophysical observations of the LTZ and STZ, this model presents a reasonable fit.

Conceptual models for asymmetry of the transition zones Pure-shear lithospheric extension models or simple seafloor-spreading models create axisymmetric crust, and thus they cannot account for the marked asymmetry in basement depth and relief in the LTZ between Newfoundland and Iberia unless the asymmetry was caused by a process active on only a single margin after breakup. Here, we take the asymmetry observed along the drilling transect (Fig. 10.4B) as representative, and we consider four simplified kinematic models that could explain the asymmetry (Fig. 10.7). We assume in each case that the LTZ, and the STZ (not depicted), developed by mantle exhumation with an increasing melt component up to the onset of normal seafloor spreading at the end of Aptian time.

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Figure 10.7 Schematic models for the opening of the Newfoundland–Iberia rift during formation of the landward transition zones (LTZ), which may explain observed asymmetry between the conjugate margins. See text for discussion. (A) Asymmetric shear exposes mantle in the Iberia LTZ and leaves thin, buoyant continental crust in the Newfoundland LTZ. (B) Asymmetric shear exposes mantle in the LTZ on both margins and a rift axis then develops in the centre of the basin; buoyancy of Newfoundland crust is provided by melt impregnation. (C) Asymmetric shear evolves into seafloor spreading that forms the Newfoundland LTZ; the rift axis later jumps to the centre of the rift (FR, failed rift). (D) Symmetric rifting and exhumation of the mantle; buoyancy of the Newfoundland LTZ is provided by later melt impregnation. In all cases, Melt 1 is emplaced during formation of LTZ crust, and Melt 2 is post-rift magmatism probably associated with sill emplacement and formation of the Newfoundland Seamounts on the Newfoundland margin. Rift-scale detachment fault (later broken by plate flexure or in-plane tensile stress) is shown by a bold black line in (A–C). Magmatic intrusions (e. g., gabbroic plutons and sills) and volcanics are in black.

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Model 1: Asymmetric extension of continental crust Although it seems unlikely that very thin continental crust (or fragments of such crust) would be widely distributed in the LTZ/ZECM, we cannot yet entirely discount this idea because the geological interpretation of the velocity structure of the uppermost few kilometres of basement is ambiguous and typically is averaged laterally over at least several kilometres. Furthermore, extreme extension of continental crust above mobile serpentinites, possibly by multiple generations of faults, is not unrealistic. The simple-shear model in Fig. 10.7A proposes that the shallower and smoother LTZ basement off Newfoundland is an upper plate consisting of very thin continental crust in a rift-scale asymmetric detachment system (e.g., Tankard and Welsink, 1987). The lower, Iberia plate would be exhumed lower continental crust and mantle, consistent with the sub-continental mantle origin of peridotites interpreted there (Whitmarsh et al., 2001). In this model, less dense continental crust would account for the shallower basement of the LTZ off Newfoundland compared to that of the denser, exhumed mantle on the Iberia margin. Interestingly, basement in the Newfoundland LTZ does not seem to show significant brittle extension in reflection profiles (as evidenced by low basement roughness and absence of low-angle detachments), and there is little evidence for lower continental crust in the LTZ/ZECM off Iberia. How, then, could extreme thinning of the continental crust be accomplished? A possible explanation is that lower crust was removed by ductile flow during early stages of rifting, with relatively little accompanying brittle extension in the upper crust (Driscoll and Karner, 1998; Whitmarsh et al., 2001). Additionally it is possible, although considered unlikely, that the Newfoundland LTZ was affected by multiple generations of faults, only the most recent of which is recognisable (Fig. 10.4), with the end product being thin continental crust with a moderately low-amplitude surface. The Newfoundland LTZ might also have been widely affected by emplacement of melt. For example, melt extracted from the rising and decompressing lower-plate mantle could have permeated the crust of the Newfoundland upper plate (Melt 1, Fig. 10.7A) while leaving the exhumed Iberia mantle relatively meltfree. It is also known from drilling at ODP Site 1276 that diabase sills were intruded in the Newfoundland LTZ during the Albian to early Cenomanian (Melt 2), up to at least 30 m.y. after the underlying basement was formed (Tucholke et al., 2004; Hart and Blusztajn, 2006). Either of these effects could have helped to smooth the basement relief. However, intrusion of mafic melts into thin continental crust would have tended to reduce the buoyancy of the Newfoundland LTZ and thus reduce the depth asymmetry with the conjugate LTZ/ZECM off Iberia.

Model 2: Asymmetric extension in an amagmatic rift In this simple-shear model (Fig. 10.7B), mantle was exhumed from beneath the Newfoundland margin along a west-dipping detachment (e.g., Tankard and

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Welsink, 1987; Whitmarsh et al., 2001) to form LTZ/ZECM basement. Extension then shifted to the middle of the rift just before anomaly
M3, splitting the mantle exposure into separate LTZs on the two margins and heralding development of the STZ. The shallower basement of the Newfoundland LTZ could be explained by buoyancy due to emplacement of melt into the exhumed mantle either (a) generated in the decompressing footwall as it was exhumed (Melt 1), or (b) associated with the post-rift sill emplacement already noted (Melt 2). The first style of melt emplacement would probably require mantle heterogeneity to explain magmatism in the Newfoundland LTZ and the relative lack of magmatism in the Iberia LTZ/ZECM. This model predicts that the basement in the Newfoundland LTZ is sub-continental mantle.

Model 3: Asymmetric accretion of oceanic crust In this model (Fig. 10.7C), asymmetric simple-shear extension as in model 2 first exposed sub-continental mantle, now observed in the Iberia LTZ/ZECM, but extension then evolved into seafloor spreading (Melt 1) that formed the Newfoundland LTZ. Like model 2, this requires a later, eastward jump of the rift axis prior to anomaly
M3 to centre the plate boundary in the basin. Buoyancy of the oceanic crust could explain the shallower basement of the Newfoundland LTZ compared to the Iberia LTZ/ZECM (serpentinised peridotite). This oceanic crust might also be intruded and its upper surface smoothed by post-rift extrusives (Melt 2). This model is consistent with geophysical and drilling observations of exposed sub-continental mantle in the Iberia LTZ/ZECM, and with oceanic crust interpreted from velocity data in the Newfoundland LTZ (Nunes, 2002). However, it does not explain the lack of an axisymmetric set of pre-M3 magnetic anomalies in the Newfoundland LTZ, and the rift-axis jump to the centre of the basin might be considered fortuitous. The model predicts that any mantle exposed within the Newfoundland LTZ should have oceanic affinities.

Model 4: Symmetric accretion of mantle The final model (Fig. 10.7D) calls for symmetric extension, with the LTZ/ZECM on both the Newfoundland and Iberia margins being primarily sub-continental mantle. From the LTZ to the STZ, the exhumed mantle would have decreasing sub-continental affinity (Whitmarsh et al., 2001) and would be increasingly intruded by melt until seafloor spreading became fully established at the end of Aptian time. Post-rift magmatism only on the Newfoundland side (Melt 2, associated with Site 1276 sill emplacement) would be required to explain the elevation and smoothness of the Newfoundland LTZ.

Rift-model summary While it is not possible to fully credit or discount any of these models, it seems very likely that the LTZ/ZECM on each of the conjugate margins is substantially sub-continental mantle. This mantle could have been exhumed either

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symmetrically or asymmetrically, and its surface might include continental fragments, although it is unlikely that it contains extensive areas of very thin continental crust. The LTZ basement contains relatively deep intrusive bodies. The STZ on the conjugate margins represents slowly increasing emplacement of melt at higher levels in exhumed mantle, although melt emplacement probably was irregular in space and time. Neither the landward nor the seaward parts of the TZs were created by normal seafloor spreading to form oceanic crust that resembles the Penrose model. We suggest that LTZ/ZECM and STZ formation should be considered to be a separate phase of ‘transitional extension’ between the end of continental rifting and the full establishment of normal seafloor spreading.

10.8 End of rifting and post-rift sedimentary history Release of in-plane tensile stress when normal seafloor spreading began at the end of Aptian time probably resulted in slight tilting of the shallow-water margins and relative sea-level regression (Tucholke et al., 2007). Combined with late Aptian low eustatic sea-level (Haq et al., 1988), this resulted in erosion of the shallow-water margins and basin flooding by turbidites that lap onto the deepwater breakup unconformity (U and BU, Figs. 10.4–10.6). During the Albian, these sediments accumulated at rates of 50–100 m/m.y., and rates decreased thereafter as the margins subsided and eustatic sea-level rose (Sibuet and Ryan, 1979; Tucholke et al., 2004). Drilling results on both margins document that the early post-rift sedimentary sequences are largely comparable to the formally defined formations of the western North Atlantic Ocean (Jansa et al., 1979). They consist of Albian-Cenomanian ‘black shales’ (cf. Hatteras Formation), deposited under low-oxygen to anoxic conditions, and Turonian-Paleocene multicoloured shales (cf. Plantagenet Formation) deposited on an oxygenated seafloor. A prominent feature of all the Cretaceous sedimentary sequences in the rift, except on elevated areas, is that they were dominated by fine to coarse material transported downslope in gravity flows. The Cenozoic sedimentary record shows more variation between the Newfoundland and Iberia margins as they became separated by the widening ocean basin. Nonetheless, the margins and deep-sea basins on both sides of the rift continued to accumulate large volumes of sediment transported downslope in gravity flows, culminating in the largely late-Pliocene to Pleistocene deposition of the modern Newfoundland and Iberia abyssal plains. A notable event that affected sedimentation patterns from about late middle Eocene onward was the development of strong abyssal circulation sourced by dense water sinking to the seafloor in the sub-Arctic seas (e.g., Tucholke et al., 2004). This event is marked by an

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erosional unconformity in the Newfoundland Basin (AU, Fig. 10.6) and in the western North Atlantic basin to the south. The deep circulation is westward intensified, so its effects (e.g., erosion, development of sediment waves) have been more marked on the Newfoundland margin than on the Iberia margin.

10.9 Conclusions Rifting between Newfoundland and Iberia extended
30–35 km-thick continental crust that was assembled during the Paleozoic Taconian, Salinic, Acadian, and Alleghenian orogenies that closed the Iapetus and Rheic oceans. Initial rifting in the Triassic to earliest Jurassic created large, deep rift basins in the Grand Banks (e.g., Jeanne d’Arc, Horseshoe, Salar–Bonnition Basins) and in Iberia (Lusitanian Basin), but it did not significantly thin the crust. The basins accumulated clastic red-bed sediments and evaporites beneath shallow seas. Following an episode of epeirogenic subsidence and carbonate accumulation, the continental platform was again extended in the latest Jurassic to Early Cretaceous, culminating in normal seafloor spreading near the Aptian/Albian boundary. Although the rift most likely experienced elevated in-plane tensile stress throughout latest Jurassic and Early Cretaceous, most extension of continental crust appears to have been concentrated near the end of the Berriasian and the end of the Hauterivian. The latter episode correlates with the separation of Flemish Cap and Galicia Bank at the northern end of the rift and with the introduction of a limited amount of melt into the rift at about chrons M5–M3 in the southern and central segments of the rift. In the southern two-thirds of the rift, from the Late Jurassic up to this time, probably sub-continental mantle was being exhumed to form a wide (up to 170 km) LTZ. Between anomaly M3 and the end of Aptian time, mantle exhumation continued, evolving into the STZ. This possibly sub-continental peridotite basement increasingly acquired the seismic and magnetic characteristics of oceanic crust as it was exhumed; it was increasingly intruded by melts, probably at shallower levels, but it did not acquire the character of Penrose-type oceanic crust. The melts could have been derived either from a mantle plume that affected the southern part of the rift from about chron M4 to younger than chron M0, or from enhanced axisymmetric late-stage upwelling of the mantle, or both. There is marked asymmetry in sediment-unloaded basement depth of the LTZ on the Newfoundland and Iberia margins. It is possible that the shallower basement on the Newfoundland margin can be explained by rift-scale simple shear with the Newfoundland LTZ comprising the upper plate, or by asymmetric accretion of ocean crust. Alternately, the rift may have evolved largely by pure shear, with buoyancy later added to the whole Newfoundland LTZ by melt intruded into the exhumed mantle. In either case, it appears that Albian to early Cenomanian magmatism, possibly associated with formation of the Newfoundland Seamounts, affected the Newfoundland margin but was not manifested on the Iberia margin.

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On each margin, the total width of the zone of exhumed mantle in the LTZ and STZ is on the order of 280 km in the southern part of the rift and
100–180 km in the northern rift between Flemish Cap and Galicia Bank. This expression of a ‘transitional phase’ of extension, occurring before the onset of normal seafloor spreading and presumably above rising oceanic asthenosphere, may be a common feature on non-volcanic rifted margins.

Acknowledgments B. E. Tucholke’s research leading to this synthesis was supported by U.S. National Science Foundation grants OCE-9819053 and OCE-0326714, by the Henry Bryant Bigelow Chair in Oceanography at Woods Hole Oceanographic Institution, and by the Ocean Drilling Program. R.B. Whitmarsh thanks the School of Ocean and Earth Science, National Oceanography Centre, Southampton for office facilities. Many of the results reported here from the Iberia margin depended on support from the U.K. Natural Environment Research Council, and the Ocean Drilling Program. The Ocean Drilling Program was sponsored by the U.S. National Science Foundation and participating countries under management of Joint Oceanographic Institutions (JOI), Inc. We thank S. Holbrook, J. Hopper, H. Lau, K. Louden, G. Manatschal, D. Sawyer, J.-C. Sibuet, D. Shillington, and H. van Avendonk for many insightful discussions over the years concerning the evolution of the Newfoundland–Iberia rift. We also thank D. G. Roberts and A. W. Bally for review and comments on the original draft of the manuscript. Contribution No. 11,375 of Woods Hole Oceanographic Institution. Note: This paper was submitted in November 2004 and accepted in revised form in May 2005. Further discussion, updates, and evolution of our interpretations as presented here can be found in Tucholke et al. (2007), Tucholke and Sibuet (2007), and Peron-Pinvidic et al. (2010).

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