# The Iceland Microcontinent and a continental Greenland-Iceland-Faroe Ridge

Gillian R. Foulger, Tony Doré, C. Henry Emeleus, Dieter Franke, Laurent Geoffroy, Laurent Gernigon, Richard Hey, Robert E. Holdsworth, Malcolm Hole, Ármann Höskuldsson, et al.

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► To cite this version:

Gillian R. Foulger, Tony Doré, C. Henry Emeleus, Dieter Franke, Laurent Geoffroy, et al. The Iceland Microcontinent and a continental Greenland-Iceland-Faroe Ridge. Earth-Science Reviews, 2020, 206, 10.1016/j.earscirev.2019.102926. insu-03685642

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https://insu.hal.science/insu-03685642v1

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Distributed under a Creative Commons Attribution 4.0 International License
THE ICELAND MICROCONTINENT AND A CONTINENTAL GREENLAND-
ICELAND-FAROE RIDGE

Gillian R. Foulger<fnref n="1" />, Tony Doré<fnref n="2" />, C. Henry Emeleus<fnref n="1" />,§, Dieter Franke<fnref n="3" />, Laurent Geoffroy<fnref n="4" />,
Laurent Gernigon<fnref n="5" />, Richard Hey<fnref n="6" />, Robert E. Holdsworth<fnref n="1" />, Malcolm Hole<fnref n="7" />, Ármann
Höskuldsson<fnref n="8" />, Bruce Julian<fnref n="1" />, Nick Kusznir<fnref n="9" />, Fernando Martinez<fnref n="10" />, Ken J.W. McCaffrey<fnref n="1" />,
James H. Natland<fnref n="11" />, Alex Peace<fnref n="12" />, Kenni Petersen<fnref n="13" />, Christian Schiffer<fnref n="1" />, Randell
Stephenson<fnref n="14" /> & Martyn Stoker<fnref n="15" />.

<footnote n="1">¹ Department of Earth Sciences, Durham University, Science Laboratories, South Rd. DH1 3LE, UK</footnote>

<footnote n="2">² Statoil UK Ltd., Statoil House, 11a Regent Street, London SW1Y 4ST, UK</footnote>

<footnote n="3">³ Bundesanstalt für Geowissenschaften und Rohstoffe (Federal Institute for Geosciences and Natural Resources), Germany</footnote>

<footnote n="4">⁴ Université de Bretagne Occidentale, Brest, 29238 Brest, CNRS, UMR 6538, Laboratoire Domaines Océaniques, 29280 Plouzané, France</footnote>

<footnote n="5">⁵ Norges Geologiske Undersøkelse (NGU), Geological Survey of Norway, Leiv Erikssons vei 39, N-7491 Trondheim, Norway</footnote>

<footnote n="6">⁶ Hawaii Institute of Geophysics and Planetology, School of Ocean and Earth Science and Technology, University of Hawaii, Honolulu, HI 96822 USA</footnote>

<footnote n="7">⁷ Department of Geology & Petroleum Geology, University of Aberdeen, Aberdeen AB243UE, UK.</footnote>

<footnote n="8">⁸ Háskóli Íslands (University of Iceland), Surlugötu 7, 101 Reykjavík, Iceland</footnote>

<footnote n="9">⁹ School of Environmental Sciences, University of Liverpool, Jane Herdman Building, Liverpool L69 3GP, UK</footnote>

<footnote n="10">¹⁰ Hawaii Institute of Geophysics and Planetology, School of Ocean and Earth Science and Technology, University of Hawaii, Honolulu, HI 96822 USA</footnote>

<footnote n="11">¹¹ Rosenstiel School of Marine and Atmospheric Science, University of Miami, Miami FL 33149, USA</footnote>

<footnote n="12">¹² Department of Earth Sciences, Memorial University of Newfoundland, St. Johns, Newfoundland, Canada, A1B 3X5; currently at School of Geography and Earth Sciences, McMaster University, 1280 Main Street West, Hamilton, Ontario, Canada L8S 4K1</footnote>

<footnote n="13">¹³ Department of Geoscience, Aarhus University, Høegh-Guldbergs Gade 2, DK-8000 Aarhus C., Denmark</footnote>

<footnote n="14">¹⁴ School of Geosciences, Geology and Petroleum Geology, Meston Building, King's College, University of Aberdeen, Aberdeen AB24 3UE, UK</footnote>

<footnote n="15">¹⁵ Australian School of Petroleum, University of Adelaide, Adelaide, South Australia 5005 Australia</footnote>

<footnote>§ deceased</footnote>
1 *Abstract*

2 The breakup of Laurasia to form the Northeast Atlantic Realm disintegrated an
3 inhomogeneous collage of cratons sutured by cross-cutting orogens. Volcanic rifted margins
4 formed that are underlain by magma-inflated, extended continental crust. North of the
5 Greenland-Iceland-Faroe Ridge a new rift-the Aegir Ridge-propagated south along the
6 Caledonian suture. South of the Greenland-Iceland-Faroe Ridge the proto-Reykjanes Ridge
7 propagated north through the North Atlantic Craton along an axis displaced ~150 km to the
8 west of the rift to the north. Both propagators stalled where the confluence of the
9 Nagssugtoqidian and Caledonian orogens formed an ~300-km-wide transverse barrier.
10 Thereafter, the ~150 x 300-km block of continental crust between the rift tips-the Iceland
11 Microcontinent-extended in a distributed, unstable manner along multiple axes of extension.
12 These axes repeatedly migrated or jumped laterally with shearing occurring between them in
13 diffuse transfer zones. This style of deformation continues to the present day in Iceland. It is
14 the surface expression of underlying magma-assisted stretching of ductile continental crust
15 that has flowed from the Iceland Microplate and flanking continental areas to form the lower
16 crust of the Greenland-Iceland-Faroe Ridge. Icelandic-type crust which underlies the
17 Greenland-Iceland-Faroe Ridge is thus not anomalously thick oceanic crust as is often
18 assumed. Upper Icelandic-type crust comprises magma flows and dykes. Lower Icelandic-
19 type crust comprises magma-inflated continental mid- and lower crust. Contemporary magma
20 production in Iceland, equivalent to oceanic layers 2-3, corresponds to Icelandic-type upper
21 crust plus intrusions in the lower crust, and has a total thickness of only 10-15 km. This is
22 much less than the total maximum thickness of 42 km for Icelandic-type crust measured
23 seismically in Iceland. The feasibility of the structure we propose is confirmed by numerical
24 modeling that shows extension of the continental crust can continue for many tens of millions
25 of years by lower-crustal ductile flow. A composition of Icelandic-type lower crust that is
26 largely continental can account for multiple seismic observations along with gravity,
27 bathymetric, topographic, petrological and geochemical data that are inconsistent with a
28 gabbroic composition for Icelandic-type lower crust. It also offers a solution to difficulties in
29 numerical models for melt-production by downward-revising the amount of melt needed.
30 Unstable tectonics on the Greenland-Iceland-Faroe Ridge can account for long-term tectonic
31 disequilibrium on the adjacent rifted margins, the southerly migrating rift propagators that
32 build diachronous chevron ridges of thick crust about the Reykjanes Ridge, and the tectonic
33 decoupling of the oceans to the north and south. A model of complex, discontinuous
34 continental breakup influenced by crustal inhomogeneity that distributes continental material
35 in growing oceans fits other regions including the Davis Strait, the South Atlantic and the
36 West Indian Ocean.
<table>
  <tbody>
    <tr>
      <td>37</td>
      <td>1</td>
      <td>Introduction</td>
      <td>5</td>
    </tr>
    <tr>
      <td>38</td>
      <td>2</td>
      <td>Continental breakup forming the Northeast Atlantic Realm</td>
      <td>6</td>
    </tr>
    <tr>
      <td>39</td>
      <td>2.1</td>
      <td>High-velocity lower crust</td>
      <td>8</td>
    </tr>
    <tr>
      <td>40</td>
      <td>2.2</td>
      <td>Seafloor spreading north and south of the Greenland-Iceland-Faroe Ridge</td>
      <td>10</td>
    </tr>
    <tr>
      <td>41</td>
      <td>2.2.1</td>
      <td>North of the Greenland-Iceland-Faroe Ridge</td>
      <td>11</td>
    </tr>
    <tr>
      <td>42</td>
      <td>2.2.2</td>
      <td>South of the Greenland-Iceland-Faroe Ridge</td>
      <td>11</td>
    </tr>
    <tr>
      <td>43</td>
      <td>3</td>
      <td>The Greenland-Iceland-Faroe Ridge</td>
      <td>12</td>
    </tr>
    <tr>
      <td>44</td>
      <td>3.1</td>
      <td>Crustal structure</td>
      <td>13</td>
    </tr>
    <tr>
      <td>45</td>
      <td>3.2</td>
      <td>The Faroe-Shetland basin—a bellwether of GIFR tectonic instability</td>
      <td>15</td>
    </tr>
    <tr>
      <td>46</td>
      <td>4</td>
      <td>A new model for the Greenland-Iceland-Faroe Ridge</td>
      <td>16</td>
    </tr>
    <tr>
      <td>47</td>
      <td>4.1</td>
      <td>Mass balance</td>
      <td>18</td>
    </tr>
    <tr>
      <td>48</td>
      <td>4.2</td>
      <td>Problems and paradoxes solved</td>
      <td>19</td>
    </tr>
    <tr>
      <td>49</td>
      <td>5</td>
      <td>Thermo-mechanical modeling</td>
      <td>21</td>
    </tr>
    <tr>
      <td>50</td>
      <td>6</td>
      <td>Geochemistry</td>
      <td>24</td>
    </tr>
    <tr>
      <td>51</td>
      <td>6.1</td>
      <td>Composition of the melt source</td>
      <td>24</td>
    </tr>
    <tr>
      <td>52</td>
      <td>6.2</td>
      <td>Temperature of the melt source</td>
      <td>25</td>
    </tr>
    <tr>
      <td>53</td>
      <td>6.3</td>
      <td><sup>3</sup>He/<sup>4</sup>He</td>
      <td>26</td>
    </tr>
    <tr>
      <td>54</td>
      <td>7</td>
      <td>Discussion</td>
      <td>27</td>
    </tr>
    <tr>
      <td>55</td>
      <td>7.1</td>
      <td>The Greenland-Iceland-Faroe Ridge</td>
      <td>27</td>
    </tr>
    <tr>
      <td>56</td>
      <td>7.2</td>
      <td>Crustal flow</td>
      <td>28</td>
    </tr>
    <tr>
      <td>57</td>
      <td>7.3</td>
      <td>Magmatism</td>
      <td>28</td>
    </tr>
    <tr>
      <td>58</td>
      <td>7.3.1</td>
      <td>The concept of the North Atlantic Igneous Province</td>
      <td>28</td>
    </tr>
    <tr>
      <td>59</td>
      <td>7.3.2</td>
      <td>Magma volume</td>
      <td>29</td>
    </tr>
    <tr>
      <td>60</td>
      <td>7.3.3</td>
      <td>The chevron ridges</td>
      <td>30</td>
    </tr>
    <tr>
      <td>61</td>
      <td>7.3.4</td>
      <td>The North Atlantic geoid high</td>
      <td>31</td>
    </tr>
    <tr>
      <td>62</td>
      <td>7.3.5</td>
      <td>Regions analogous to the GIFR</td>
      <td>31</td>
    </tr>
    <tr>
      <td>63</td>
      <td>7.3.6</td>
      <td>Regions analogous to the NE Atlantic Realm</td>
      <td>32</td>
    </tr>
    <tr>
      <td>64</td>
      <td>8</td>
      <td>Conclusions</td>
      <td>33</td>
    </tr>
    <tr>
      <td>65</td>
      <td></td>
      <td></td>
      <td></td>
    </tr>
    <tr>
      <td>66</td>
      <td></td>
      <td></td>
      <td></td>
    </tr>
  </tbody>
</table>
67 List of acronyms. See also Table 2.

68

69 GIFR - Greenland-Iceland-Faroe Ridge

70 JMMC - Jan Mayen Microplate Complex

71 SDR - Seaward-dipping reflector

72 NVZ - Northern Volcanic Zone

73 EVZ - Eastern Volcanic Zone

74 WVZ - Western Volcanic Zone

75 HVLC - High-velocity lower crust

76 $T_P$ - potential temperature

77 $V_P$ - compressional (P-) wave velocity

78 REE - rare-Earth element

79 SCLM - sub-continental lithospheric mantle

80

81 Keywords: Atlantic; Iceland; continental breakup; tectonics; Icelandic-type crust; SDRs;

82 geochemistry; geophysics.

83

84
85 **1 Introduction**

86 The NE Atlantic Realm, the region north of the Charlie Gibbs Fracture Zone, including the
87 seas and seaboards west of Greenland, has persistently resisted attempts to account for many
88 of its features in terms of conventional plate tectonics. Although the region figured
89 prominently in the development of the spectacularly successful continental drift and plate
90 tectonic theories, e.g., with the discovery of symmetrical magnetic anomalies across the
91 Reykjanes Ridge, it has also defied predictions made by this theory that are successful in
92 most other areas. This is particularly true along the Greenland-Iceland-Faroe Ridge (GIFR)
93 where the crust is typically 30 km thick and the bathymetry a full kilometer shallower than is
94 expected by cooling and subsidence models for oceanic-crust [Detrick *et al.*, 1977]. These
95 observations cannot be satisfactorily explained simply as conventional sea-floor-spreading
96 with a larger-than-typical magmatic rate at Iceland (Figure 1).

97 It is ironic that, despite the GIFR region not fitting the simple plate tectonic theory, it played
98 an important role in development of that theory. In the early 20th century Iceland attracted the
99 attention of Alfred Wegener who, as part of his theory of continental drift [Wegener, 1915],
100 predicted that Greenland and Scandinavia were separating at 2.5 m/a. Although his estimate
101 of rate was two orders of magnitude too large, his general theory was correct. Wegener was
102 influenced by arguments that a land bridge, postulated on biogeographical grounds to have
103 connected Europe and America, was inconsistent with isostasy. At the time, such land
104 bridges were widely invoked to explain the similarity, at some times in geological history,
105 between biota on opposite sides of wide oceans. Wegener recognized that biogeographical
106 observations worldwide could be explained by continental drift without land bridges.

107 Following acceptance of continental drift, the land bridge theory was essentially dropped.
108 Ironically, the NE Atlantic is perhaps the only place in the world where a long, ocean-
109 spanning land bridge did actually exist [Ellis & Stoker, 2014] (Section 3). The reason why
110 such a bridge existed, even when the ocean had attained a width of over 1000 km, is one
111 feature of many of the NE Atlantic that has, to date, not been satisfactorily explained.

112 A model for development of the NE Atlantic Realm that can account for these and all other
113 observations in a holistic way is required. Models that involve simple palinspastic
114 reconstructions of Laurasian super-continent breakup, and assume a bimodal crustal
115 composition (continental or oceanic) with sharp boundaries, are insufficient [Barnett-Moore
116 et al., 2018; Nirrengarten et al., 2018]. Models that explain the quantity, distribution and
117 petrology of igneous rocks in an *ad hoc* fashion are not forward-predictive and cannot
118 account for observations such as the close juxtaposition of volcanic and non-volcanic
119 margins, high-velocity lower crust (HVLC), frequent ridge jumps, and southward-
120 propagating rifts on the Reykjanes Ridge [Hey et al., 2010; Peron-Pinvidic & Manatschal,
121 2010]. Nor can such models, a century after Wegener's work, explain why a land bridge
122 spanned the NE Atlantic Ocean until it had attained a width of ~1,000 km, and why 40% of
123 its length remains subaerial to the present day as the island of Iceland.

124 In this paper we develop such a model. We propose that the currently ~1,200 km wide
125 Greenland-Iceland-Faroe Ridge (GIFR) formed by magma-assisted continental extension
126 facilitated by ductile crustal flow, in a similar fashion to magmatic passive margins. The
127 extraordinary width of the GIFR was enabled by the inclusion of a ~45,000 km² block of
128 continental crust which we term the Iceland Microcontinent. The lower part of the ~30 km
129 thick GIFR crust is magma-dilated continental mid- and lower crust. Surface extension has
130 been taken up on the GIFR by distributed, migrating rifts with shear between them
131 accommodated diffusely. Continental material is dispersed throughout the GIFR and sea-
132 floor spreading has not yet been established on a single, stable rift. Complete continental
133 breakup has thus still not fully occurred at this latitude.

134 Our paper is structured in the following way. First, we describe the unusual setting and
135 complex history of breakup of the NE Atlantic Realm that predicated the subsequent
136 complexities (Section 2). We then summarize structural and tectonic observations from the
137 GIFR and the adjacent Faroe-Shetland basin (Section 3). In Section 4 we present our new
138 model for the structure and evolution of the GIFR. Section 5 presents a numerical thermo-
139 mechanical simulation that illustrates the model is physically viable given reasonable
140 geological assumptions and Section 6 shows that it is consistent with the petrology,
141 geochemistry and source potential temperatures of NE Atlantic igneous rocks. Finally, in
142 Section 7, we discuss wider implications and analogous regions elsewhere in the oceans.

143 **2 Continental breakup forming the Northeast Atlantic Realm**

144 Opening of the NE Atlantic Realm in the early Cenozoic was not a simple, abrupt, isolated
145 event. It was the latest event in a >300 Myr period of episodic rifting and cooling that lasted
146 from the Late Palaeozoic through the Mesozoic. It affected a region extending some half the
147 circumference of the Earth and disassembled a heterogeneous patchwork of cratonic blocks
148 and orogens [Bingen & Viola, 2018; Gasser, 2014; Gee et al., 2008b; Peace et al., this
149 volume; Wilkinson et al., 2017].

150 Final breakup occurred by magma-assisted continental extension [Gernigon et al., this
151 volume; Lundin & Doré, 2005; Peace et al., this volume; Roberts, 2003; Roberts et al., 1999;
152 Skogseid et al., 2000; Soper et al., 1992]. The crust extended by tens or hundreds of
153 kilometers from the Rockall Trough to the Barents Sea [Funck et al., 2017; Gaina et al.,
154 2017; Skogseid et al., 2000; Stoker et al., 2017]. Pre-breakup magmatism occurred
155 throughout the region including in Britain, the Rockall Trough, East Greenland, the Faroe
156 Islands and small-volume, small-fraction, scattered fields found in west Greenland and
157 Newfoundland [Larsen et al., 2009; Peace et al., 2016; Wilkinson et al., 2017]. Final
158 development of the axes of breakup in the NE Atlantic was influenced by both the direction
159 of extensional stress, pre-existing structure, and magmatism [Peace et al., 2018; Peace et al.,
160 submitted; Schiffer et al., this volume].

161 Greenland is cross-cut by several orogens that continue across formerly adjacent landmasses.
162 Easterly orientated orogens include the Inglefield mobile belt in the north
163 (Paleoproterozoic—ca. 1.96 - 1.91 Ga), the central Greenland Nagssugtoqidian orogen
164 bounded to the north by the Disko Bugt suture and to the south by the Nagssugtoqidian front
165 (Paleoproterozoic—ca. 1.86 - 1.84 Ga), and the south Greenland Ketilidian orogen
166 (Paleoproterozoic—ca. 1.89 - 1.80 Ga) (Figure 2) [Garde et al., 2002; van Gool et al., 2002].
167 On the Eurasian continent the Nagssugtoqidian orogen is represented in Scotland as the
168 Lewisian gneiss (Laxfordian) and the Ketilidian orogen is represented in NW Ireland as the
169 Rhinns Complex (Figure 3).

170 The much younger Caledonian suture formed in the Ordovician-Devonian and closed the
171 Tornquist Sea and Iapetus Ocean to unite Laurentia, Baltica and Avalonia [Pharaoh, 1999;
172 Schiffer et al., this volume; Soper et al., 1992]. The Scottish Caledonides lie orthogonal to
173 the eastward continuation of the Nagssugtoqidian and Ketilidian orogens [Holdsworth et al.,
174 2018]. The western frontal thrust of this suture runs down east Greenland ~100 - 300 km
175 from the coast [Gee et al., 2008a; Haller, 1971; Henriksen, 1999; Henriksen & Higgins,

176 1976]. A dipping feature imaged seismically using receiver functions at ~40 - 100 km
177 beneath east Greenland is interpreted as a subducted slab, trapped in the continental

178 lithosphere, when the Caledonian suture finally closed (Figure 4) [Schiffer et al., 2014].

179 Residual Caledonian slabs beneath the region were predicted earlier by plate models for the
180 geochemistry of Icelandic volcanics [Foulger & Anderson, 2005; Foulger et al., 2005]. A

181 congruent structure—the Flannan reflector—has been imaged seismically beneath north
182 Scotland [Schiffer et al., 2015; Smythe et al., 1982].

183 The breakup phases that formed the oceans west and east of Greenland are described in detail
184 by Peace et al. [this volume], Gernigon et al. [this volume] and Martinez and Hey [this
185 volume]. It is summarized here and a brief chronology of the most significant events is given

186 in Table 1. The north-propagating mid-Atlantic Ridge reached the latitude of the future
187 Charlie Gibbs Fracture Zone in the Late Cretaceous (~86.3 - 83.6 Ma) and the Rockall

188 Trough formed (Figure 1). The rift then propagated west of present-day Greenland at ~63 Ma
189 forming magma-poor margins and opening the Labrador Sea [Abdelmalak et al., 2018; Keen
190 et al., 2018; Nirrengarten et al., 2018; Oakey & Chalmers, 2012; Roest & Srivastava, 1989].

191 Propagation proceeded unhindered across the Grenville and Ketilidian orogens and the North
192 Atlantic craton but stalled at the junction of the Nagssugtoqidian and Rinkian orogens

193 [Connelly et al., 2006; Grocott & McCaffrey, 2017; Peace et al., 2018; Peace et al.,
194 submitted]. There, the crust was locally thick [Clarke & Beutel, 2019; Funck et al., 2007;

195 Funck et al., 2012; Peace et al., 2017; St-Onge et al., 2009] and pre-existing subducted slabs
196 may also have been preserved in the lithosphere [Heron et al., 2019]. Rift propagation stalled

197 and the Davis Strait NNE-SSW sinistral, right-stepping transtensional accommodation zone
198 formed. This subsequently opened by magma-assisted continental transtension and

199 transpression. Further north Baffin Bay opened by a combination of continental extension
200 and possibly some subsidiary sea-floor spreading [Chalmers & Laursen, 1995; Chauvet et al.,

201 2019; Oakey & Chalmers, 2012; Suckro et al., 2012; Welford et al., 2018]. The Davis Strait

202 today is a 550-km wide shallow ridge of extended, magma-inflated, continental crust that
203 spans the ocean from Baffin Island to West Greenland [Dalhoff et al., 2006; Heron et al.,

204 2019; Schiffer et al., 2017].

205 At ~56 - 52 Ma rifting began to propagate east of Greenland forming the proto-Reykjanes

206 Ridge and a ridge-ridge-ridge triple junction at the location of the current Bight fracture zone
207 (Figure 1). Shortly thereafter, at ~50 - 48 Ma, the pole of rotation for Labrador Sea/Baffin

208 Bay opening migrated south by ~1000 km resulting in clockwise rotation of ~30 - 40° of the
209 direction of motion of Greenland relative to Laurentia [Oakey & Chalmers, 2012; Srivastava,

210 1978]. As a consequence the Labrador Sea/Baffin Bay plate boundary west of Greenland
211 became less favorable to extension [Gaina et al., 2017] and motion was progressively

212 transferred to the axis east of Greenland. From ~36 Ma, opening was taken up entirely in the
213 NE Atlantic [Chalmers & Pulvertaft, 2001; Gaina et al., 2017].

214 As was the case for breakup west of Greenland, development of the mid-Atlantic Ridge in the
215 NE Atlantic was strongly influenced by pre-existing structure [Schiffer et al., this volume].

216 The classic “Wilson Cycle” model suggests that continental breakup occurs along older

217 sutures [Ady & Whittaker, 2018; Buiter & Torsvik, 2014; Chenin et al., 2015; Krabbendam,
218 2001; Petersen & Schiffer, 2016; Vauchez et al., 1997]. The collage of cratons and cross-
219
220

cutting orogens that comprised the disintegrating Laurasian supercontinent had several sutures that influenced breakup.

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228
229
230

Development of the oceanic regions north and south of the GIFR described in more detail in Sections 2.2.1 and 2.2.2 is summarized here. North of the present GIFR, the axis of extension opened by southerly propagation within the Caledonian orogen. That orogen consists of overthrust stacks of nappes and sinistral shear zones including the Møre-Trøndelag Fault Zone (Norway) and the Walls Boundary-Great Glen Fault and Highland Boundary Fault (Scotland). These features may have controled the structures that opened [Dewey & Strachan, 2003; Doré *et al.*, 1997; Fossen, 2010; Peace *et al.*, this volume]. The new mid-Atlantic Ridge formed obliquely along the orogen, however, so the propagating rift tip eventually intersected its edge at the Caledonian Western Frontal Thrust [Schiffer *et al.*, this volume]. There, it stalled.

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237

South of the present GIFR the proto-Reykjanes Ridge propagated north from the Bight Fracture Zone, cut unhindered across the Ketilidian orogen as did the Labrador Sea rift, and split the North Atlantic craton. It arrived at the confluence of the transverse Nagssugtoqidian and Caledonian orogens at ~C21 (50 - 48 Ma) [Elliott & Parson, 2008] and stopped at a location ~300 km to the south and ~150 km to the west of the stalled, south-propagating ridge tip to the north. It was between and around these two stalled ridge tips that the GIFR formed, by magma-assisted deformation of the continental region between them.

238

## 2.1 *High-velocity lower crust*

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240
241
242
243
244

High velocity lower crust (HVLC) is widespread beneath the margins of the NE Atlantic and has very similar geophysical properties to the lower part of the Icelandic-type crust that underlies the GIFR [Bott, 1974; Foulger *et al.*, 2003]. Because of this, understanding the origin and composition of HVLC is key to unraveling the development and current structure of the NE Atlantic. In this section, we discuss in detail its geophysical characteristics and possible origins.

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250
251

Before oceanic crust began to form in the NE Atlantic Realm, wide rifted margins of stretched continental crust developed and in some areas were blanketed by thick sequences of seaward-dipping basalt flows (seaward-dipping reflectors—SDRs) [Á Horni *et al.*, 2016; Talwani & Eldholm, 1977]. Lithospheric necking occurred by normal faulting in the upper crust and distributed magma inflation and ductile flow in the mid- and lower crust (Figure 5). Multiple changes in extension direction complicated the final structure [Barnett-Moore *et al.*, 2018].

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The volcanic rifted margins may be divided into Inner-SDR and Outer-SDR regions [Planke *et al.*, 2000]. The Inner-SDRs comprise lavas up to 5 - 10 km thick that blanket heavily dyke-injected continental upper crust formed during the continental extensional necking phase [e.g., Benson, 2003; Geoffroy, 2005; Geoffroy *et al.*, 2015]. Beneath this the sill-injected lower crust exhibits high seismic velocities. Outer-SDRs sometimes lie seaward of these and directly overlie thinner HVLC, with seismic properties identical to the HVLC beneath the necked continental crust [Geoffroy *et al.*, 2015]. HVLC may also extend for up to 100 km beneath both the adjacent oceanic and continental domains [Funck *et al.*, 2016 and references therein; Rudnick & Fountain, 1995; Thybo & Artemieva, 2013].
261 HVLC typically has seismic velocities intermediate between those expected for crust and mantle. Constraints on its density are poor because the densities of the SDRs and underlying lower crust are not well known. Geoffroy *et al.* [2015] define two kinds of HVLC – LC1 and LC2. Typical working values for velocity and density are, for LC1 $V_P$ ~7.2 - 7.3 km/s and density 3000 - 3100 kg/m³, and for LC2 $V_P$ ~7.6 to 7.8 km/s and density 3200 - 3300 kg/m³ (Figure 5) [Bauer *et al.*, 2000; Geoffroy *et al.*, 2015; Schiffer *et al.*, 2016].

267 These geophysical properties are ambiguous regarding the composition, origin, and tectonic significance of HVLC. Possible lithologies include:

269 • Ultra-high-pressure granulite/eclogite crystalline basement representing exhumed continental mid- and lower crust [Abdelmalak *et al.*, 2017; Ebbing *et al.*, 2006; Gernigon *et al.*, 2004; Gernigon *et al.*, this volume; Mjelde *et al.*, 2013]. Such material can have both high $V_P$ (7.2 - 8.5 km/s) and high density [2.8-3.6 g/cm3; Fountain *et al.*, 1994]. It outcrops in the Norwegian Western Gneiss Region which continues beneath the North Sea and the platform east of the Møre basin. Its top surface may comprise old suture accommodation zones that controlled deformation prior to breakup;

277 • Syn-extension sill-intruded mid-to-lower continental crust. In wide-angle seismic lines most HVLC beneath Inner-SDRs present high-amplitude, folded reflectors disconnected from the deepest layered lower crust [Clerc *et al.*, 2015; Geoffroy, 2005; Geoffroy *et al.*, 2015]. Such deformation fits with seaward ductile flow of this layer;

281 • Exhumed and syn-rift serpentinized mantle. The HVLC beneath the mid-Norwegian early Cretaceous basins, the Labrador Sea, Baffin Bay, Rockall Trough and the Porcupine basin may be partially syn-rift serpentinized mantle exhumed beneath the axes of maximum extension [Keen *et al.*, 2018; Lundin & Doré, 2011; O'Reilly *et al.*, 1996; Peron-Pinvidic *et al.*, 2013; Reston *et al.*, 2001; Reynisson *et al.*, 2011]. It is directly observed at amagmatic margins, *e.g.*, the Iberian margin, where the serpentinization is thought to be caused by seawater infiltrating down crustal faults and reacting with exhumed mantle at shallow depths. The HVLC beneath the NE Atlantic SDRs lies under several kilometers of sediments and crust and it is unlikely that seawater can penetrate sufficiently deep to cause pervasive serpentinization beneath the basalt [Abdelmalak *et al.*, 2017; Gernigon *et al.*, 2004; Zastrozhnov *et al.*, 2018];

293 • Inherited serpentinized material. Water could have been sourced from inherited Caledonian or Sveconorwegian-Grenvillian mantle wedge material [Fichler *et al.*, 2011; Petersen & Schiffer, 2016; Schiffer *et al.*, 2016; Slagstad *et al.*, 2017]. The source of NE Atlantic basalts is known to be wet [Jamtveit *et al.*, 2001; Nichols *et al.*, 2002]. Pressure conditions corresponding to deep crust/shallow upper mantle depths and temperatures of 500 - 700°C should not be exceeded for serpentinite to exist [*e.g.*, Petersen & Schiffer, 2016; Ulmer & Tromms Dorff, 1995]. Numerical modeling confirms that such material can be preserved in rifted margins [Petersen & Schiffer, 2016] and that its strength would be less than half that of dry peridotite [Escartin *et al.*, 2001];

303 • Mantle infiltrated with gabbroic melt. Such material has been observed at magma-poor margins [Lundin & Doré, 2018; Müntener *et al.*, 2010] and would have an average seismic velocity midway between that of mantle and gabbro ($V_P$ ~7 km/s);
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* Hybrid material comprising a mixture of some or all of the above on various scales.
For example, Schiffer *et al.* [2015] interpret HVLC bodies beneath east Greenland as Caledonian subduction material including eclogitized mafic crust. What appears geophysically to be a continuous layer might also vary laterally in composition—a classic example of geophysical ambiguity [Mjelde *et al.*, 2002].

An interpretation of HVLC as underplated material, i.e., high-temperature melt that accumulated during initial opening of the NE Atlantic [Eldholm & Grue, 1994b; Mjelde et al., 1997; Mjelde et al., 2002; Mjelde et al., 1998; Thybo & Artemieva, 2013] is challenged by key geophysical and structural observations from the outer Vøring basin. There, Cretaceous deformation was partly controlled by the top of a HVLC dome before the main magmatic event in the Late Paleocene-Early Eocene, suggesting that the dome may predate breakup magmatism by at least 15 - 25 Myr [Abdelmalak et al., 2017; Gernigon et al., 2004; Gernigon et al., 2006].

In summary, the provenance of the HVLC underlying the Outer SDRs is ambiguous but it likely includes a large proportion of continental crust. As a consequence the exact locations of the outer limits of continuous offshore continental material (the continent-ocean boundary) is poorly known in some areas [Bronner et al., 2011; Eagles et al., 2015; Gernigon et al., 2015; Lundin & Doré, 2018; Schiffer et al., 2018]. Continental crust may grade into thick oceanic crust via a magmatic transition zone tens of kilometers wide of stretched, intruded continental crust—the continent-ocean transition [Eagles et al., 2015; Eldholm et al., 1989; Gernigon et al., this volume; Meyer et al., 2009]. The width of the continent-ocean transition may be partly controlled by the degree of stretching with narrow extensional zones forming where new rifts follow pre-existing fabric, and wide zones where rifts cross-cut tectonic fabric [Buck, 1991; Dunbar & Sawyer, 1988; Harry et al., 1993; Schiffer et al., this volume].

Full rupture of the crust leading to region-wide sea-floor spreading may be discontinuous, diachronous and segmented [Elliott & Parson, 2008; Guan et al., 2019; Manton et al., 2018; Schiffer et al., this volume; Theissen-Krah et al., 2017]. Continental fragments trapped between pairs of volcanic rifted margins and transported into the new ocean to form “C-blocks” may be widespread (Figure 5; Section 4) [Geoffroy et al., 2015; Geoffroy et al., 2015; Theissen-Krah et al., 2017]. Continental crust may also be distributed by igneous mulling as seen in the southern Jan Mayen Microplate Complex (JMMC; Section 2.2.1), and by small-scale lateral rift migrations [Bonatti, 1985; Gernigon et al., 2012; Gillard et al., 2017]. Continental fragments may range in size from the 100-km scale down. Geophysical ambiguity and blanketing of microcontinents with lavas hinder mapping the full distribution of continental crust in the oceans. The eastern margin of the JMMC, for example, is overlain by SDRs and the subaerial part of the GIFR (i.e. Iceland) is blanketed with lavas younger than ~17 Ma [Breivik et al., 2012; Gudlaugsson et al., 1988]. Geochemistry can be used to complement geophysics by testing the viability of proposed HVLC petrologies (Section 6).

## 2.2 *Seafloor spreading north and south of the Greenland-Iceland-Faroe Ridge*

Clear, well-mapped, linear magnetic anomalies reveal the contrasting histories of ocean opening north and south of the GIFR (Figure 6). Breakup did not occur simultaneously along the entire seaboard, as often assumed, but involved several isolated propagators and intermediate continental blocks [Elliott & Parson, 2008; Gernigon et al., this volume].
## 2.2.1 North of the Greenland-Iceland-Faroe Ridge

The earliest anomalies are likely associated with magma injection into extended continental crust. True sea-floor spreading on the Aegir Ridge began at ~54 Ma (C24r). It started at its northern end and propagated south to reach its full extent by ~52 Ma (Chron C23). Tectonic reorganization and fan-shaped spreading occurred about this ridge C22-C21 (~48 Ma) with spreading slower in the south than in the north (Table 1) [Gernigon et al., 2015].

Much if not all of the southern extension deficit was accommodated by diffuse, dyke-assisted crustal dilation in the continental crust immediately to the west. This region later became the southern JMMC [Brandsdóttir et al., 2015]. Crustal extension of up to 500% occurred forming mullioned crust [Gernigon et al., 2015; Schiffer et al., 2018]. Extension ultimately concentrated on the most westerly axis of dilation which developed into the Kolbeinsey Ridge. The first unambiguous magnetic anomaly formed there at ~24 Ma [C6/7; Blischke et al., 2017; Vogt et al., 1980].

The Aegir Ridge dwindled and became extinct a little after ~31 - 28 Ma [C12-C10; Gernigon et al., 2015] after which all spreading north of the GIFR was taken up on the proto-Kolbeinsey Ridge. This migration of the locus of extension likely occurred as a result of a tectonic reorganization that rotated the local direction of motion counter-clockwise [Gaina et al., 2017]. This would have rendered the southern part of the Aegir Ridge less favorable for spreading and encouraged extension on the proto-Kolbeinsey Ridge. That extension progressively detached the continental block and adjacent mullioned crust between the proto-Kolbeinsey Ridge and the Aegir Ridge to form the JMMC [Schiffer et al., 2018]. Opening of the Atlantic north of the GIFR (e.g. the Norwegian-Greenland Sea) thus occurred on a series of unconnected, sub-parallel, migrating, propagating rifts.

The northern part of the JMMC is a coherent microcontinent on the 100-km scale [Peron-Pinvidic et al., 2012]. SDRs formed on its eastern margin [Kodaira et al., 1998]. The crust that makes up its southern part is severely intruded continental crust with clear rift zones [Brandsdóttir et al., 2015]. The nature of its transition into Iceland is, however, unknown.

Despite developing in the highly magmatically productive environment of the early NE Atlantic the late Aegir Ridge was magma-starved and formed oceanic crust only 4 - 7 km thick [Breivik et al., 2006; Greenhalgh & Kusznir, 2007]. This contrasts with both the Kolbeinsey Ridge and the Reykjanes Ridge which are underlain by oceanic crust ~10 km thick. Extreme variations in magmatic rate over short distances are inconsistent with mechanisms of melt production that envisage extensive, coherent regions of influence and suggest, instead, local dependency on melt productivity [Lundin et al., 2018; Simon et al., 2009].

## 2.2.2 South of the Greenland-Iceland-Faroe Ridge

South of the GIFR, on the European side, poorly constrained, complicated magnetic anomalies SW of the Faroe Islands suggest early disaggregated sea-floor spreading. The first unambiguous and continuous spreading anomaly south of the Faroe Plateau formed at ~47 Ma (C21) [Elliott & Parson, 2008; Ellis & Stoker, 2014; Stoker et al., 2012]. On the Greenland side, the oldest linear magnetic anomalies produced by the proto-Reykjanes Ridge date from C24-22 (56 - 52 Ma), but they may represent rift-related basalt extrusion in the Outer-SDR region and not true oceanic spreading. Linear magnetic anomalies terminate
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along the SE Greenland margin, unlike the European side where they are continuous along the margin. This is consistent with early westward migration of the spreading axis. It finally stabilized along a zone ~150 km west of the Aegir Ridge.

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Extension proceeded normal to the strike of the Reykjanes Ridge and the continental edges until ~37 - 38 Ma (C17) when an abrupt counter-clockwise rotation of the direction of plate motion occurred. Spreading in the Labrador Sea then rapidly ceased (Table 1) [Gaina et al., 2017; Jones, 2003; Martinez & Hey, this volume]. The Bight ridge-ridge-ridge triple junction ceased to exist and the linear Reykjanes Ridge reconfigured to a right-stepping ridge-transform array such that the new ridge segments were normal to the new direction of plate motion.

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Subsequently, and up to the present day, the Reykjanes Ridge has been slowly migrated east by a series of small-offset, right-stepping propagators within the plate boundary zone that have eliminated the transforms [Benediktsdóttir et al., 2012; Hey et al., 2010; Martinez & Hey, this volume]. They originate at the GIFR and migrate south at rates of 10 - 25 cm/a, each slicing a few kilometers off the Eurasian plate and transferring it to the North American plate [Hey et al., 2016]. At least five and possibly as many as seven propagators [Jones et al., 2002] have now transferred a swathe of the Eurasian plate ~30 km wide to the North American plate between the GIFR and the Bight Fracture Zone [Benediktsdóttir et al., 2012].

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Progression of each propagator tip is associated with transient changes in thickness of ~2 ± 1 km in the oceanic crust formed. This has the curious consequence that the Reykjanes Ridge is flanked by diachronous “chevrons” (also called “V-shaped ridges”) of alternating thick and thin crust that are most clearly seen in the gravity field (Figure 7) [Vogt, 1971].

415

### 3 The Greenland-Iceland-Faroe Ridge

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The GIFR comprises a ~1,200-km-long, shallow, trans-oceanic aseismic ridge up to 450 km wide in the northerly direction (Figure 1). At present, 40% of it is exposed above sea level in Iceland. It is shallower than 600 m and 500 m deep offshore west and east Iceland respectively, ~1000 m shallower than the ocean basins to the north and south (Figure 8).

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The GIFR was subaerial along its entire length for most of the history of the NE Atlantic. Biogeographical evidence for plant and animal dispersal [Denk et al., 2011] and dating of the onset of overflow of intermediate- and deep waters between the Norway and Iceland basins [Ellis & Stoker, 2014; Stoker et al., 2005b] suggest that it formed a largely intact, trans-Atlantic land bridge (the Thulean land bridge) until ~10 - 15 Ma and that much survived above sea level longer than this. This leads to the surprising conclusion that the Thulean land bridge survived intact until the NE Atlantic Ocean had attained a width of ~1000 km.

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Magnetic anomalies on the GIFR are poorly defined, broader than classical oceanic spreading anomalies, and resemble more closely anomalies on the outer SDRs (Figure 6) [e.g., Gaina et al., 2017]. Very few can be clearly traced across the GIFR so the detailed history of breakup in this region cannot be deduced reliably. Previous interpretations have relied largely on extrapolation of anomalies to the north and south that are clear, assuming simple oceanic crustal accretion in the region between.

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Prior explanations for the poorly developed magnetic anomalies include repeated dyke intrusion into the same zone during more than one magnetic chron, re-magnetization by later intrusions, weathering, lateral migration of spreading centers and magmatism at multiple
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spreading centers [Bott, 1974]. Unclear anomalies are also expected because basalt extrusion was subaerial and flooded older lavas, and because the legacy magnetic data available are poor quality, limited, and poorly levelled.

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We concur with these suggestions but go further and propose that distinct, oceanic-type linear magnetic anomalies do not exist on the GIFR because it does not comprise oceanic crust formed by classical sea-floor spreading. Instead, much of it may consist of magma-dilated, ductile continental crust. Upper Icelandic-type crust [Bott, 1974; Foulger et al., 2003] corresponds to current basaltic production. Lower Icelandic-type crust corresponds to magma-inflated mid- and lower continental crust, the most likely lithology for the HVLC that is widespread beneath the NE Atlantic passive margins (Section 2.1).

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## 3.1 Crustal structure

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The GIFR has been the target of numerous refraction, wide-angle reflection, and passive seismic experiments [Foulger et al., 2003] as well as gravity, magnetic and magnetotelluric work [Beblo & Bjornsson, 1978; 1980; Beblo et al., 1983; Eysteinsson & Hermance, 1985; Hermance & Grillot, 1974; Thorbergsson et al., 1990]. It was the anomalous seismic nature of its crust that led to it being termed “Icelandic-type” [Bott, 1974; Foulger et al., 2003]. It features an upper crust with a thickness of ~3 - 10 km with high vertical velocity gradients, and a lower crust ~10 - 30 km thick with low vertical velocity gradients (Figure 9 and Figure 10) [Darbyshire et al., 1998a; Foulger et al., 2003; Holbrook et al., 2001; Hopper et al., 2003]. The lower crust has a $V_P$ of 7.0 - 7.3 km/s. Icelandic-type crust has, in recent years, usually been assumed to be anomalously thick oceanic crust with the lower crust equivalent to oceanic layer 3. That model became the default assumption after Bjarnason et al. [1993] reported a deep reflecting horizon at ~ 20 - 24 km depth beneath SW Iceland. It replaced an earlier model that interpreted the layer beneath the upper crust as anomalously hot mantle [Angenheister et al., 1980; Gebrande et al., 1980; Palmason, 1971; Tryggvason, 1962].

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The model that Icelandic-type lower crust is oceanic is inconsistent with other observations. Isostatic studies reveal the density of the lower crust to be ~3150 kg/m³, which is too high for it to be oceanic [Gudmundsson, 2003; Menke, 1999]. At the same time, its seismic velocity is too low for normal mantle peridotite. Models involving partial melt are ruled out by the low attenuation of seismic shear waves which suggests that Icelandic-type lower crust is no hotter than 800 - 900°C if it is peridotite [Sato et al., 1989] and 875 - 950°C if it is gabbroic [Menke & Levin, 1994; Menke et al., 1995].

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The theory that Icelandic lower crust is oceanic is largely based on interpreting deep seismic reflections as the Moho. However, such reflections can also be interpreted as sills intruded into continental lower crust. Refracted head waves are almost never observed in Iceland and the large amplitudes of reflections expected from a Moho are not observed in receiver functions [Du & Foulger, 1999; Du et al., 2002; Du & Foulger, 2001]. These properties are similar to those of HVLC beneath the Inner- and Outer SDRs of the continental margins [Mjelde et al., 2001]. The possible compositions and provenances of that material, discussed in Section 2.1, thus provide candidates for Icelandic-type lower crust.

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A serpentinized mantle origin for Icelandic-type lower crust is unlikely. If it were serpentinized mantle, ~20% of serpentinization of peridotite at ~1 GPa (~30 km depth) is required [Christensen, 2004]. Serpentinization in a rifting environment occurs from the top down and water is unlikely to be able to reach the mantle at the active rift zones of Iceland. If
480 it did, it would only be stable at temperatures < 700 °C or possibly < 500 °C [Tuttle & Bowen, 1958] and the Icelandic lower crust is hotter than this [Menke & Levin, 1994; Menke et al., 1995; Sato et al., 1989]. The only other possible way of serpentinizing the mantle is via fluxing from beneath. In the NE Atlantic such serpentinization could have occurred in the Caledonian suture [Fichler et al., 2011] and water is present in the source of basalts erupted in Iceland [Jamtveit et al., 2001; Nichols et al., 2002]. However, no peridotite xenoliths have been found in Iceland despite a century of extensive geological mapping and drilling, suggesting that, whereas serpentinite may exist beneath some rifted margins, it probably does not comprise lower Icelandic-type crust or HVLC beneath the adjacent volcanic margins.

489 Transitional crust comprising massively dyke- and sill-intruded, hyper-extended mid- and lower continental crust is the most likely composition for Icelandic-type lower crust, as it is for much of the HVLC beneath the volcanic margins. There is considerable support for this:

492 • The seismic velocity and density of continental lower crust match those of Icelandic-type lower crust. Continental lower crust is thought to comprise predominately mafic garnet-bearing granulites which have $V_p$ ~7.1 - 7.3 km/s and densities of 3000 - 3150 kg/m$^3$ [Rudnick & Fountain, 1995]. It may also contain minor components of metapelite, intermediate and felsic granulites and mafic melts that would reduce $V_p$ and density.

498 • The thickness of the brittle surface layer in Iceland and the viscosity of the underlying material have been constrained by geodetic studies of post-diking stress relaxation [Foulger et al., 1992; Heki et al., 1993; Hofton & Foulger, 1996a; b; Pollitz & Sacks, 1996] and post-glacial rebound [Sigmundsson, 1991]. The brittle surface layer is ~10 km thick, a value that is consistent with the maximum depth of earthquakes [Einarsson, 1991] and corresponds roughly to the upper crust from explosion seismology and receiver functions (Figure 9). The lower crust beneath has a viscosity of ~$10^{19}$ Pa s and is thus ductile.

506 • The Faroe Islands are underlain by continental crust topped by > 6 km of basalt [Bott et al., 1974; Ólavsdóttir et al., 2017]. Seismic data from the eastern part of the Iceland-Faroe Ridge detect stretched continental crust similar to that underlying the Rockall Bank where HVLC has been interpreted as inherited continental crust of Palaeo-European affinity [Bohnhoff & Makris, 2004].

511 • Palinspastic reconstructions of Iceland require up to 150 km of crust older than the surface lavas to underlie the island—the extreme westerly and easterly ~15-Ma palaeo-rift products are separated by ~450 km whereas only ~300 km of widening could have occurred at the ambient rate of 1.8 cm/a. Reassembly of the NE Atlantic Ocean also requires up to 150 km of continental crust (original unstretched width) to lie in the ocean [Blischke et al., 2017; Bott, 1985; Foulger, 2006; Gaina et al., 2009; Gaina et al., 2017; Gernigon et al., 2015]. A similar width is required by the original lateral offset of the tips of the Aegir Ridge and proto-Reykjanes Ridge. A southerly continuation of the JMMC beneath the GIFR would be a simple source of this material [Bott, 1985; Foulger & Anderson, 2005; Schiffer et al., 2018]. Icelandic-type crust also underlies the transitional region between the NE Icelandic shelf and the JMMC [Brandsdóttir et al., 2015].

523 • Magma-assisted extension at the far western and eastern ends of the proto-GIFR, outside of the axes of breakup, is predicted by stress modeling and may have fed additional continental crust into the developing GIFR.
526 • There are multiple lines of petrological and geochemical evidence for a component of
527 continental crust in Icelandic lavas, including Proterozoic and Mesozoic zircons
528 [Amundsen *et al.*, 2002; Foulger, 2006; Paquette *et al.*, 2006; Schaltegger *et al.*,
529 2002] elevated 87Sr/86Sr and Pb isotope ratios [Prestvik *et al.*, 2001] and extensive
530 silicic and intermediate rocks including rhyolite and icelandite—an Fe-rich form of
531 andesite (Section 6).

532 **3.2 The Faroe-Shetland basin—a bellwether of GIFR tectonic instability**

533 The Faroe-Shetland basin comprises the eastern extension of the GIFR, is thus sensitive to
534 tectonic activity in that zone, and has been unstable throughout the Palaeogene-early
535 Neogene [Stoker *et al.*, 2018; Stoker *et al.*, 2005b]. Key phases are summarized in Figure 11
536 and include the following.

537 • Paleocene (~63 - 56 Ma): The pre-breakup rifting phase (late Danian—Thanetian) was characterized by formation of a series of sag and fault-controlled sub-basins [Dean *et al.*, 1999; Lamers & Carmichael, 1999], coeval borderland uplift events (rift pulses) [Ebdon *et al.*, 1995; Goodwin *et al.*, 2009; Mudge, 2015] and rifting and extension accompanied by volcanism [Mudge, 2015; Ólavsdóttir *et al.*, 2017].

542 • Latest Paleocene (~56 - 55 Ma): Uplift [Ebdon *et al.*, 1995] and extrusion of syn-
543 breakup flood basalts and tuffs [Mudge, 2015] probably mark the onset of local,
544 discontinuous sea-floor spreading [Passey & Jolley, 2009].

545 • Early-Mid-Eocene (~54 - 46 Ma): The syn-breakup rift-to-drift transition continued
546 during the early/mid-Ypresian-early Lutetian [Stoker *et al.*, 2018]. Cyclical coastal
547 plain, deltaic and shallow-marine deposits attest to tectonic instability linked to
548 episodic uplift of the Munkagrunnur and Wyville Thomson ridges on the south flank
549 of the basin [Ólavsdóttir *et al.*, 2010; Ólavsdóttir *et al.*, 2013b; Stoker *et al.*, 2013].
550 Onset of continuous sea-floor spreading in the Norway basin (chron C21) was
551 accompanied by uplift events, continued growth of the Wyville Thomson and
552 Munkagrunnur ridges, and formation of inversion domes in the basin [Ólavsdóttir *et
553 al.*, 2010; Ólavsdóttir *et al.*, 2013b; Ritchie *et al.*, 2008; Stoker *et al.*, 2013; Stoker
554 *et al.*, 2018].

555 • Late Paleogene-early Neogene (~35 - 15 Ma): The present-day basin physiography
556 was initiated in the latest Eocene/Early Oligocene with sagging leading to basin-ward
557 collapse of the margin west of Shetland [Stoker *et al.*, 2013]. Onlapping Oligocene
558 and Lower Miocene basinal sequences were deformed by compressional stresses and
559 widespread inversion and fold growth culminated in the early Mid-Miocene [Johnson
560 *et al.*, 2005; Ritchie *et al.*, 2008; Stoker *et al.*, 2005c].

561 • Mid-Miocene—Pleistocene (16/15 Ma - present): Basinal sedimentation was
562 dominated by deep-water deposits [Stoker *et al.*, 2005b] with Early Pliocene uplift
563 and tilting of the West Shetland and East Faroe margins accompanied by basinal
564 subsidence and reorganization of bottom current patterns [Andersen *et al.*, 2000;
565 Ólavsdóttir *et al.*, 2013b; Stoker *et al.*, 2005a; Stoker *et al.*, 2005b]. Mid-and Late
566 Pleistocene sedimentation was dominated by shelf-wide glaciations [Stoker *et al.*,
567 2005a].

568 In summary, the Faroe-Shetland basin has experienced persistent tectonic unrest from the
569 Paleocene to the Early Miocene (~63 - 15 Ma). This is reflected onland in the Faroe Islands
570 in Paleogene and younger faults and dykes that show progressive changes in the direction of
extension prior to and following NE Atlantic break-up [Walker et al., 2011]. This chronic
unrest likely reflects both instability on the GIFR to the west and the protracted breakup of
the wider NE Atlantic region.

## 4 A new model for the Greenland-Iceland-Faroe Ridge

In this section, we build on the background given above and propose a new working
hypothesis for development of the GIFR and how this affected the rest of the NE Atlantic
Realm. Numerical modeling of the processes we propose, and model fit with petrology and
geochemistry, are discussed in Sections 5 and 6.

As described above, the NE Atlantic Realm formed in a disorderly way as a consequence of
inherited strength anisotropy, coupled with frequent changes in the poles of rotation of sub-
regions [Hansen et al., 2009; Schiffer et al., 2018]. North of the GIFR, the Aegir rift opened
by southward propagation obliquely along the Caledonian orogen. It stalled at the western
frontal thrust and hooked around to the west (Figure 1). The Reykjanes Ridge to the south
stalled at the Nagssugtoqidian orogen, ~300 km south of the Caledonian frontal thrust and
~150 km west of the Aegir Ridge (Section 2.2). The Reykjanes Ridge and Aegir Ridge thus
formed a pair of propagating, approaching, laterally offset rifts. The broad barrier formed by
the Nagssugtoqidian and Caledonian orogens prevented them from propagating further and
conceivably eventually forming a continuous, conventional oceanic plate boundary.

As a consequence, the continental region between their tips, the ~300 x 150 km Iceland
Microcontinent, deformed by magma-assisted, distributed continental transtension and
developed into the GIFR as the ocean widened (Figure 12). The crust beneath the Iceland
Microcontinent and flanking areas thinned by ductile flow in its deeper parts. Extensive
magmatism built SDRs of the kind observed on the eastern margin of the JMMC and the NE
Atlantic rifted margins. Initially, the GIFR may have comprised an array of four passive
margins—one on each of the east Greenland and west Faroe margins, and one on either side of
the Iceland Microcontinent.

As the GIFR lengthened, and up to the present day, deformation persisted in a distributed
style along a series of ephemeral extensional rifts and diffuse, intermediate, poorly developed
shear transfer zones [Gerya, 2011]. The loci of extension repeatedly reorganized by migrating
laterally to positions that were stress-optimal and likely also influenced by pre-existing
structures in the underlying continental crust. Rifts that became extinct were transported
laterally out of the actively extending, central part. As it formed the GIFR was blanketed by
lavas in the style of volcanic-rifted-margins. Similar rift migrations also occurred in the
eastern Norway basin where the oceanic crust is thickest [Gernigon et al., 2012]. After ~48
Ma (C22) it seems that this style of extension persisted only on the GIFR. The permanent
disconnect between the Aegir Ridge and the Reykjanes Ridge and the low spreading rate in
the NE Atlantic (1 - 2 cm/a) would have further encouraged long-term diffuse deformation.

Figure 13 shows palinspastic reconstructions of the observed positions of active and extinct
rifts in the NE Atlantic Realm at various times. Swathes of extinct, short, NE-orientated
ridges similar to those that are currently active onland in Iceland are observed also in
submarine parts of the GIFR [Hjartarson et al., 2017]. There is insufficient observational data
at present to fully reconstruct the sequence of deformation on the GIFR because of the
blanketing lava flows and insufficient geophysical and geological research to date.
Nevertheless, in Figure 14 we attempt such a reconstruction by extrapolating in time from
known active and extinct rifts [Hjartarson *et al.*, 2017; Johannesson & Saemundsson, 1998].

A large block of crust older than the surface lavas is required by palinspastic reconstructions to lie beneath Iceland [Foulger, 2006]. Thus, much of the Iceland Microcontinent may still exist beneath Iceland and comprise a C-block [Geoffroy *et al.*, 2015; Geoffroy *et al.*, submitted]. C-blocks are expected to be flanked by Outer-SDRs and the geometry of some dykes and lava flows in Iceland resemble these [Bourgeois *et al.*, 2005; Hjartarson *et al.*, 2017]. It has long been speculated that the continental crust required by geochemistry to underlie Iceland (Section 6) comprises a southerly extension of the magma-dilated southern JMMC. Give the extent of continental crust that is required to underlie the GIFR it may be more appropriate to view the JMMC as an offshore extension of the Iceland Microcontinent.

Deformation on the GIFR cannot be described by traditional rigid plate tectonics and corresponding reconstructions. It corresponds to the case of multiple overlapping ridges, the limit of an extensional zone [Engeln *et al.*, 1988]. It may be likened to a lateral array of hyper-extended SDRs underlain by HVLC comprising heavily intruded, stretched, ductile continental crust. Repeated rejuvenation of the rift axes by lateral migration may have boosted volcanism. Westerly migrations may have induced extension to the north to concentrate in the westernmost axis of extension in the southern JMMC, leading to extinction of the Aegir Ridge and formation of the Kolbeinsey Ridge at ~24 Ma. That migration switched the sense of the ridges north and south of the GIFR from right-stepping to left-stepping.

Iceland is ~450 km wide in an EW direction and exposes ~40% of the GIFR (Figure 1). The oldest rocks found there to date are 17 Ma. There is no evidence, or reason to think, that the tectonic style on the GIFR was fundamentally different in the past from present-day Iceland. On the contrary, the similarity of the submarine GIFR synclines to structure on land in Iceland suggests that it was the same [Hjartarson *et al.*, 2017].

Onland in Iceland extension over the last ~15 Ma has occurred via multiple unstable, migrating, overlapping spreading segments connected by complex, immature shear transfer zones that reorganize every few Myr (Figure 15). These include the South Iceland Seismic Zone [Einarsson, 1988; 2008] and the Tjörnes Fracture Zone [e.g., Rognvaldsson *et al.*, 1998]. Both are broad, diffuse seismic zones that deform in a bookshelf-faulting manner and have not developed the clear topographic expression of faults that experience long-term repeated slip.

There is geological evidence in Iceland for at least 12 spreading zones (Table 2) of which seven are currently active, two highly oblique to the direction of plate motion, one waning, one propagating, two non-extensional and five extinct. At least five lateral rift jumps are known and a sixth is currently underway via transfer of extension from the Western Volcanic Zone (WVZ) to the Eastern Volcanic Zone (EVZ). Extension has always been concentrated in a small number of active, ephemeral rift zones at any one time (Figure 15).

There is no evidence that mature sea-floor spreading has yet begun anywhere along the GIFR. If such were the case it would be expected that all extension would be rapidly transferred to that zone and normal-thickness oceanic crust (i.e., ~6 - 7 km) would begin to form. Indeed, the fact that rift-zone migrations are still ongoing in Iceland suggest that this is not the case. There may be some narrow zones where embryonic sea-floor spreading began
but was abandoned due to subsequent lateral rift jumps, e.g., immediately east of the east Icelandic shelf, and in the deep channel in the Denmark Strait. Until this can be confirmed, however, it remains a possibility that full continental breakup has not yet occurred in the latitude band of the GIFR. This fundamentally challenges the concept that continental breakup has yet occurred in this part of the NE Atlantic.

## 4.1 Mass balance

The GIFR today is 1,200 km long with a lower crust generally ~20 km thick and maximally ~30 km beneath central Iceland (Figure 9). If a substantial part of this is continental, a large volume thus needs to be accounted for. Taking a present-day average breadth for the GIFR of ~200 km, the surface area is ~0.24 x 10⁶ km². If an average thickness of 15 km of continental material lies beneath, a volume of ~3.6 x 10⁶ km³ is required.

We propose that this material was sourced from the Iceland Microcontinent and flanking continental regions by ductile flow of mid- and lower crust. Ductile flow can stretch such crust to many times its original length (necking), draw in material from great distances, and maintain large crustal thicknesses. Numerical thermo-mechanical modeling (Section 5) confirms that these processes can account for the lower-crustal thicknesses proposed and even increase crustal thickness as material rises to fill the void created by rupture of the upper crust.

Flow is enabled by the low viscosity of the lower crust beneath Iceland. This has been shown to be 10¹⁸ - 10¹⁹ Pa s by GPS measurements of post-diking stress relaxation following a regional, 10-m-wide dyke injection episode in the Northern Volcanic Zone (NVZ) 1975 - 1995 [Bjornsson et al., 1979; Foulger et al., 1992; Heki et al., 1993; Hofton & Foulger, 1996a; b]. Numerical modeling of those data also showed that the surface, brittle layer was approximately 10 km thick. The low viscosity found for the lower crust was confirmed by measurements and modeling of the rapid isostatic rebound from retreat of the Weichselian ice cap in Iceland and melting of the Vatnajökull glacier in south Iceland [Sigmundsson, 1991]. [Sigmundsson, 1991].

As a prelude to numerical thermo-mechanical modeling we present here a simple mass-balance calculation. Inland in Greenland, receiver function studies indicate a Moho depth of ~40 km [Kumar et al., 2007]. The Caledonian crust of east Greenland is currently up to ~50 km thick [Darbyshire et al., 2018; Schiffer et al., 2016; Schmidt-Aursch & Jokat, 2005; Steffen et al., 2017]. A pre-breakup Caledonian crustal thickness of about 60 km and a post-breakup thickness of 30 km [Holbrook et al., 2001] is not unrealistic.

Beneath the Faroe-Shetland basin, crustal thinning left only a 10-km-thick crust while below the Faroe shelf and islands seismic data indicate basement modified by weathering, igneous intrusions and tuffs with a thickness of about 25 - 35 km [Richard et al., 1999]. Beneath the banks to the SW of the Faroe Islands the thickness of the subvolcanic crust is up to 25 km but it is as little as 8 km beneath the channels between them [Funck et al., 2008]. In the Faroe Bank Channel and the channel between George Bligh and Lousy Bank, in prolongation of the GIFR, the continental middle crust is almost completely gone and the lower crust is dramatically thinned. Initial and final thicknesses of 60 km and 15 km are reasonable.

Thinning of the mid- and lower crust of 30 km (Greenland) and 45 km (Faroe region) extending ~200 km along the margins and ~100 km inland could provide ~1.5 x 10⁶ km³ of
701 material. Assuming original northerly and easterly dimensions for the Iceland Microcontinent
702 of 300 km and 150 km respectively, and thinning from an original 60 km to 15 km, an
703 additional ~2 x 10⁶ km³ of material is accounted for. Together, this totals ~3.5 x 10⁶ km³ of
704 material, very close to the ~3.6 x 10³ km³ required.

705 This mass balance calculation illustrates simply that our model is reasonable. It also shows
706 that the Iceland Microcontinent can provide over half of the continental material required.
707 This suggests that the formation of such an unusually large microcontinent was likely a key
708 element in the development of this unique region.

709 **4.2 Problems and paradoxes solved**

710 The model we propose can account naturally for many hitherto unexplained observations
711 from the GIFR and surrounding regions, and it is supported by multiple lines of evidence. In
712 particular, it offers a solution to the decades-old problems of why the Thulean land bridge
713 existed, and the nature of Icelandic-type crust. Thus:

714 • A composition of Icelandic-type crust comprising magma-inflated continental crust
715 blanketed with lavas can explain the high topography and bathymetry of the GIFR
716 and its prolonged persistence above sea level.

717 • The assumption that the full thickness of Icelandic-type crust corresponds to melt has
718 been widely accepted ever since Bjarnason *et al.* [1993] reported a reflective horizon
719 at ~20 - 24 km depth beneath south Iceland which they interpreted as the Moho. That
720 model cannot, however, account for the absence of refracted seismic phases (Section
721 3.1) which is inconsistent with gabbroic crust overlying mantle with a step-like
722 interface velocity increase. The lack of such refractions is, however, consistent with
723 the reflective horizon being a sill-like structure within or near the base of magma-
724 inflated continental crust.

725 • Icelandic-type lower crust has a seismic velocity *V*<sub>P</sub> of 7.0 - 7.3 km/s and a density of
726 ~3150 kg/m³. No reasonable basaltic petrology is consistent with this [Gudmundsson,
727 2003; Menke, 1999], but these values fit a composition of magma-inflated continental
728 crust.

729 • A lower crust containing significant continental material solves the paradox of
730 magmatic production on the GIFR. Icelandic-type lower crust cannot be gabbroic
731 because a melt layer up to 40 km thick cannot be explained with any reasonable
732 petrology and temperatures (Section 6) [Hole & Natland, this volume]. If the melt
733 layer corresponds only to Icelandic-type upper crust plus magma inflating the lower
734 crust—possibly a total thickness of up to ~15 km—much less melt needs to be
735 explained.

736 • A substantial volume of continental material in the lower crust can explain why the
737 thicknesses of the upper and lower crustal layers on the GIFR are de-correlated
738 (Figure 9) [Foulger *et al.*, 2003; Korenaga *et al.*, 2002]. In particular, the lower crust
739 is thick throughout a NW-SE swathe across central Iceland where the upper crust is of
740 average thickness. In the far south, the upper crust has its maximum thickness but the
741 lower crust is unusually thin (Figure 9).

742 • MORB melt formed in the mantle below the crust passes through the latter, melting
743 fusible components to a high degree, boosting melt volume, and acquiring the
744 continental signature observed in Icelandic rocks including the geochemistry,
745 Proterozoic and Mesozoic zircons, and voluminous felsic and intermediate petrologies
746 (Section 6).

747 • The numerous, northerly trending synclines detected by seismology throughout
748 submarine parts of the GIFR are readily explained as volcanically active extensional
749 zones that were abandoned by lateral jumps and subsequently became extinct
750 [Hjartarson *et al.*, 2017] (Figure 8).

751 • If the JMMC is a northerly extension of the Iceland Microcontinent, the former may
752 have shared the tectonic instability of the GIFR, providing an explanation for why the
753 JMMC broke off east Greenland.

754 Our new model for the GIFR can account for the many unusual extensional, transtensional
755 and shear tectonic elements in the region. These include the curious distributed, bookshelf
756 mode in which shear deformation is taken up in Iceland in the South Iceland Seismic Zone
757 and the Tjörnes Fracture Zone [Bergerat & Angelier, 2000; Einarsson, 1988; Taylor *et al.*,
758 1994].

759 It can also account for the widespread hook-like tectonic morphology that resembles the tips
760 of overlapping propagating cracks (Figure 16). These suggest that short extensional elements
761 are abundant. The southernmost Aegir Ridge is hooked westward, mirroring the shape of the
762 Blosseville coast of Greenland and curving into the transverse Caledonian frontal thrust
763 [Brooks, 2011]. The extensional NVZ of Iceland curves westward at its northern end where it
764 links with the Kolbeinsey Ridge via the Tjörnes Fracture Zone. At its north end, the
765 Reykjanes Ridge hooks to the east where it runs onshore to form the Reykjanes Peninsula
766 extensional transform zone [Taylor *et al.*, 1994]. The direction of extension in the EVZ is
767 rotated ~35° clockwise compared with the NVZ as shown by both the strike of dyke- and
768 fissure swarms and current measurements of surface deformation made using GPS [*e.g.*, Perlt
769 *et al.*, 2008]. The southernmost tip of this propagating rift, the Vestmannaeyjar archipelago,
770 hooks to the west, complementing the east-hooking northern Reykjanes Ridge and Reykjanes
771 Peninsula Zone (Figure 16).

772 The contrasting tectonic morphology and behavior north and south of the GIFR are naturally
773 explained by tectonic decoupling by the GIFR that separates them. North of the GIFR the
774 boundary is dominated by spreading ridges orthogonal to the direction of extension, separated
775 by classic transform faults. To the south, the Reykjanes Ridge as a whole is oblique to the
776 spreading direction and devoid of transform faults. Numerous tectonic events occurred north
777 or south of the GIFR but not in both regions simultaneously [Gernigon *et al.*, this volume;
778 Martinez & Hey, this volume]. In Iceland, tectonic decoupling can explain the north-south
779 contrast in geometry, morphology and history of the rift zones and the north-south
780 asymmetry in geochemistry [*e.g.*, Shorttle *et al.*, 2013]. The latter may be important in
781 mapping the distribution of continental material beneath Iceland.

782 Unstable tectonics on the GIFR can further explain the diachronous chevrons of alternating
783 thick and thin crust that form at the tips of propagators within the Reykjanes Ridge plate
784 boundary zone (Figure 7; Sections 2.2.2 and 7.3.3). The onset times of several of the most
785 recent of these propagators at the GIFR coincide with major ridge jumps in Iceland (Table 1).
786 These observations are consistent with the propagators being triggered by major tectonic
787 reorganizations on the GIFR. Several similar ridges are observed in the oceanic crust east of
788 the Kolbeinsey Ridge [Jones *et al.*, 2002]. The chronic instability of the mid-Norwegian shelf
789 and the adjacent Faroe-Shetland basin throughout the Palaeogene-earliest Neogene is also
790
791

accounted for [Ellis & Stoker, 2014; Gernigon et al., 2012; Stoker et al., 2018] (Figure 11)
(Section 3.2).

792

## 5 Thermo-mechanical modeling

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795
796
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798

We tested the plausibility of unusually prolonged survival of intact continental crust beneath the GIFR by modeling numerically the behavior under extension of structures characteristic of an ancient orogen such as the Caledonian and surrounding regions. The crust is required to have stretched to over twice its original width, retained a typical thickness of ~20 km, and persistently extended along more than one axis even up to the present day i.e. it underwent long-term, diffuse extension.

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807

We used a two-dimensional thermo-mechanical modeling approach [Petersen & Schiffer, 2016] to calculate the visco-elastic-plastic response of an ancient orogen under simple extension. Full details of our methodological approach along with petrologic, thermodynamic, rheological, thermal conductivity, radiogenic heat productivity, initial model state, boundary conditions and melt productivity are described in detail by Petersen et al. [2018]. The initial state for the model we use here differs from that used by Petersen et al. [2018] only in that a) a uniform adiabatic temperature with potential temperature $T_P=1325°C$ is assumed for the entire mantle, and b) there is no MORB layer at the upper/lower mantle boundary.

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817

Prior to continental breakup, crustal thickness and structure likely varied throughout the region, but precise details of the pre-rift conditions are not well known. Insights may be gained from well-studied, currently intact orogens. The Himalaya orogen, a heterogeneous stack of multiple terranes, entrained subduction zones, and continental material, is underlain by one or more fossil slabs trapped in the lithosphere. These locally thicken the crust and their lower parts are in the dense eclogite facies (Figure 17) [Tapponnier et al., 2001]. The Palaeozoic Ural Mountains preserve a crustal thickness of 50 - 55 km [Berzin et al., 1996]. The Caledonian crust is up to ~50 km thick under east Greenland [Darbyshire et al., 2018; Schiffer et al., 2016; Schmidt-Aursch & Jokat, 2005; Steffen et al., 2017] and ~45 km thick beneath Scandinavia [Artemieva & Thybo, 2013; Ebbing et al., 2012].

818
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822
823
824

The pre-breakup crust in the region of the future NE Atlantic comprised the south-dipping Ketilidian and Nagssugtoqidian orogens and the bivergent Caledonian orogen with east-dipping subduction of Laurentia (Greenland) and west-dipping subduction of Baltica (Scandinavia) (Figure 4). The GIFR thus formed over fossil forearc/volcanic front lithosphere that may initially have had a structure similar to that of the Zangbo Suture of the Himalaya orogen (Figure 17). North of the GIFR the supercontinent broke up longitudinally along the Caledonian suture where the crust was thinner.

825
826
827
828
829
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831
832

We modeled the Caledonian frontal thrust as an orogenic belt where the lithosphere contrasts with that of the flanking Greenland and Scandinavia areas in a) increased crustal thickness, and b) eclogite from fossil subducted slabs embedded in the lithospheric mantle (Figure 4) [Schiffer et al., 2014]. The eclogite is relatively dense, potentially driving delamination, but is rheologically similar to dry peridotite [Petersen & Schiffer, 2016 and references therein]. Additional weakening of the hydrated mantle wedge preserved under the suture would enhance the model behavior we describe below [Petersen & Schiffer, 2016]. For the mantle, we assume a pyrolite composition that is subject to melt depletion during model evolution.
833 Figure 18 shows our initial, simplified model setup [Petersen *et al.*, 2018]. Pre-rift continental crustal thickness is 40 km for the lithosphere adjacent to the orogen. The crust beneath the 834 200-km-wide orogen is 50 km thick and underlain by an additional 20-km-thick slab of 835 HVLC with an assumed mafic composition. Phase transitions and density are self- 836 consistently calculated from pressure/temperature conditions throughout the model such that 837 the topmost part of the body is above eclogite facies and the lower part is in the eclogite 838 facies and thus negatively buoyant. 839

840 Densities and entropy changes are pressure- and temperature-dependent and calculated using 841 Perple_X-generated lookup tables [Connolly, 2005] based on the database of Stixrude and 842 Lithogow-Bertelloni [2011]. We use a wide box (2000 km x 1000 km) to enable simulation 843 of considerable extension distributed over a broad region, a grid resolution of 2 km, and a run 844 time of 100 Myr at a full extension rate of 1 cm/a. Rifting of the lithosphere is kinematically 845 forced by imposing plate separation at a rate of 0.5 cm/a via outwards perpendicular 846 velocities at both left and right boundaries throughout the depth range of 0 - 240 km.

847 A multigrid approach is employed to solve the coupled equations for conservation of mass, 848 energy and momentum as described by Petersen *et al.* [2015; 2018]. For the continental crust, 849 we assume plagioclase-like viscous behavior [Ranalli, 1995]. The HVLC is assumed to 850 follow an eclogite flow law [Zhang & Green, 2007]. The upper mantle is assumed to follow a 851 combined diffusion/dislocation creep flow law [Karato & Wu, 1993]. The lower mantle, here 852 defined as the region where the pressure/temperature-dependent density exceeds 4300 kg/m³, 853 approximately corresponding to Ringwoodite-out conditions, is assumed to follow the linear 854 flow law inferred by Čížková *et al.* [2012].

855 As the structure extends, rifting develops in the broad region where the crust is thickest. The 856 Moho temperature is highest there (i.e. ~800°C; Figure 18 central panels) due to greater 857 burial and radiogenic heat production. During the first 10 Myr of widening, extensional strain 858 within the crust is laterally distributed due to the delocalizing effect of flow in the lower crust 859 [e.g., Buck, 1991]. Thinning of the mantle lithosphere is not counteracted by this effect and 860 within 10 Myr it has been thinned by a factor of up ~2. This results in onset of decompression 861 melting after ~12 Myr. At this point, the crust in the stretched orogen retains a large thickness 862 of 30 - 40 km. This contrasts with the sequence of events where the crust is thin and brittle 863 under which conditions decompression melting only onsets after breakup i.e. when complete 864 thinning of the continental crust occurs [Petersen *et al.*, 2018].

865 Thinning of the mantle lithosphere leads to lateral density gradients between the 866 asthenosphere and displaced colder lithospheric mantle that destabilize the lithospheric 867 mantle [Buck, 1986; Keen & Boutilier, 1995; Meissner, 1999]. Consequently, the mantle 868 lithosphere, including the already negatively buoyant HVLC, starts to delaminate at ~12 Myr 869 (Figure 18). As a result, asthenosphere at a potential temperature (*T*<sub>P</sub>) of ~1325°C and crust at 870 an initial temperature of ~600 - 800°C are rapidly juxtaposed. This leads to increased heat 871 flow into the crust which therefore remains ductile enough to flow and continues to extend in 872 the delocalized, “wide rift mode” of Buck [1991]. The loci of extension repeatedly migrate 873 laterally and this mode of deformation continues as long as lower crust is available. The loci 874 of extension only stabilize after ~70 Myr.

875 Figure 18 shows the predicted structure after 51 Myr, approximately the present day, and a 876 magnification of the 51-Myr lithology panel is shown in *. The continental crust is still intact 877 across the now 1,200-km-wide ocean and extending diffusely. Decompression melting is
occurring beneath two zones. The HVLC body has disintegrated and the largest pair of
fragments are 200 - 300 km in diameter. These, along with carapaces of lithospheric mantle,
have subsided to a depth near the base of the transition zone. Between them, a narrow arm of
mantle upwells and flattens at the base of the crust to form a broad sill-like body ~200 km
thick that underlies the entire ocean. Upwelling includes some material from just below the
transition zone as a consequence of the sinking HVLC displacing uppermost lower-mantle
material.

Full lithosphere breakup has not occurred by 51 Myr. It is imminent at 71 Myr (Figure 18).
The lower crust beneath distal areas flows into the thinning extending zone and only when
the supply of ductile lower crust is depleted does extension localize, leading to full
lithosphere rupture and sea-floor spreading. This may take several tens of millions of years.
In the case where the crust is thinner and/or HVLC is lacking modeling predicts transition to
sea-floor spreading after only a few million years.

The generic model presented here shows that it is possible for extending lithosphere to
remain for as long as 70 Myr in a delocalized ‘wide’ stretching mode [senso Buck, 1991]
with lower crust from distal areas flowing into the extending zone. Since the delamination,
the mantle lithosphere has effectively been removed and decompression melting occurs
where the mantle wells up beneath the rift system. Together, these interdependent processes
provide a physical mechanism for how continental crust could be preserved beneath the GIFR
despite more than 50 Myr of extension (Figure 20). At the same time, the model accounts for
the magmatism observed on the GIFR in that it predicts decompression melting in the mantle.
These melts rise, intrude and erupt, covering the continental crust as it stretches, and would
produce crust similar to the “embryonic” crust proposed to occur in the Norway Basin
[Geoffroy, 2005; Gernigon et al., 2012].

The predictions of our model for present-day structure compare well with seismic
tomography images (Figure 21). For example, a cross section through the full-waveform
inversion tomographic model of Rickers et al. [2013] shows several features that are in close
correspondence to those we predict. These include the flanking high-wave-speed bodies at
the bottom of the transition zone, a narrow, weak, vertical, low-wave-speed body between
them and a broader, stronger, low-wave-speed body in the top ~200 km underlying the entire
ocean. The high-wave-speed bodies correspond to the delaminated lithospheric mantle and
the low-wave-speed anomalies correspond to the temperature anomalies predicted by the
modeling (Figure 19). A mantle temperature anomaly of ~ 30°C is predicted beneath the
entire ocean down to ~ 200 km depth as a result of upper mantle upwelling. The seismic
anomaly could then be explained by this temperature anomaly and a resulting increase in the
degree of partial melt by up to 0.5% [Foulger, 2012]. Such a temperature anomaly is
consistent with the low values predicted by Ribe et al. [1995] who modeled the topography of
the region, and the petrological estimates of Hole and Natland [this volume]. Seismic
tomography images are notoriously variable in detail, in particular anomaly amplitudes
[Foulger et al., 2013], and we thus place most significance on the correspondence between
the shape of the predicted (Figure 19) and observed (Figure 21) anomalies.

Our results differ from those of existing mechanical models in that breakup of the continental
crust is more protracted [e.g., Brune et al., 2014]. For example, Huismans and Beaumont
[2011] showed that extension of lithosphere with relatively weak crust results in pre-breakup
wide-rift-mode extension for ~35 Myr. Our model differs from theirs by having dense HVLC
that delaminates as a consequence of rifting thereby increasing heat flow into the crust. This
924 enables wide rifting to persist for much longer than where no HVLC is present, even in the
925 absence of an especially weak lower crustal rheology.

926 The amount and detailed history of wide-mode extension is controlled by the thickness of the
927 initial crust and rheology-governing parameters such as initial thermal state, heat flow,
928 radiogenic heat production and crustal flow laws. We examined models that varied some of
929 these parameters to investigate the sensitivity of our results to the assumed initial conditions.
930 We modeled crustal thicknesses of 35 km for the region and 40 km for the orogen, and
931 lithosphere thickness of 100 km with quartzite-like crustal rheology. Similar results to those
932 described above were obtained. Factors we do not incorporate in our simple model that would
933 further encourage crustal stretching and delay breakup include increased basal heat flow
934 and/or internal heat production and inclusion of 3D effects that would permit ductile mid-
935 and lower crust to flow along the strike of the Caledonides towards the GIFR during
936 extension.

937 **6 Geochemistry**

938 All aspects of the petrology and geochemistry of igneous rocks in the NE Atlantic Realm are
939 consistent with a model where Icelandic lower crust contains a substantial amount of
940 continental crust. The geochemical and petrological work most powerful to test this model is
941 that which addresses the composition and potential temperature (T<sub>P</sub>) of the melt source.

942 **6.1 *Composition of the melt source***

943 The source of Icelandic lavas cannot be explained by mantle peridotite alone [e.g., Presnall &
944 Gudfinnsson, 2011]. A component of continental material is required and some studies have
945 presented evidence that this could be of Caledonian age [Breddam, 2002; Chauvel &
946 Hemond, 2000; Korenaga & Kelemen, 2000]. It could come from subducted slabs still
947 remaining in the shallow mantle, as has been proposed earlier [Foulger & Anderson, 2005;
948 Foulger et al., 2005]. The observations could also be explained by the upward flow of mantle
949 melt through a substrate of stretched, magma-inflated continental crust similar to some
950 HVLC beneath the passive margins.

951 *Titanium*: The petrology and geochemistry of igneous rocks along the mid-Atlantic ridge
952 change radically at the Icelandic margin. Low-TiO₂ basalts are found on the Reykjanes and
953 Kolbeinsey Ridges and in the rift zones of Iceland. These rocks do not follow the MORB
954 array of Klein and Langmuir [1987] but have the least Na₃ and Ti₈ of the entire global array.
955 These lavas are probably derived mostly from a peridotitic MORB source.

956 Basalts with high-TiO₂ and FeO(T) signatures occur in Iceland, Scotland, east and west
957 Greenland, but not on the Reykjanes or Kolbeinsey Ridges. These basalts cannot come from
958 MORB-source mantle—the source is required to have distinct Fe-Ti-rich material and other
959 important geochemical indicators such as REE that are not found in MORB-source mantle.
960 The extent of differentiation beneath Icelandic central volcanoes is also high enough to
961 produce abundant silicic lavas—icelandite and rhyolite—in association with the FeO(T)-
962 TiO₂-rich basalts. These rocks comprise 10% of the surface volcanics of Iceland but are not
963 present on the adjacent submarine ridges and are uncommon on all other oceanic spreading
964 plate boundaries.
A candidate for the source of such lavas is lower continental crust, possibly pyroxenite/eclogite arising from gabbro with elevated TiO₂ and FeO(T) at pressures in the eclogite facies, along with refractory sub-continental lithospheric mantle (SCLM), originally from the Greenland and European margins and still present in the central North Atlantic. A hybrid source of this sort can explain the diversity of Icelandic magmas [Foulger & Anderson, 2005; Foulger et al., 2005; Korenaga, 2004; Korenaga & Kelemen, 2000]. Because there is no major isotopic anomaly, the source cannot be very old. Candidate material is common in continental lithosphere. For example, xenolith suites of lower crustal cumulates from Permian lamprophyres in Scotland have the required characteristics [Downes et al., 2007; Hole et al., 2015].

The close proximity of the low- and high-Ti, high-FeO(T) basalts suggests that their different sources are physically close together. It is clear that these sources have been tapped since the opening of the NE Atlantic as they are also seen in the same successions in the Skaergaard intrusion in east Greenland [Larsen et al., 1999]. The high-TiO₂ and FeO(T) lavas found in Iceland are typical of lavas derived from sub-continental lithospheric mantle and pyroxenite. No more than 20 - 30% of pyroxenite in a hybrid source is required to explain the observations.

*Isotope ratios:* Elevated ⁸⁷Sr/⁸⁶Sr and Pb isotope ratios are found in basalts from east and southeast Iceland [Prestvik et al., 2001]. This has been interpreted as requiring a component of continental material in the source beneath Iceland. That component could come from crust or detached SCLM buried beneath surface lavas [Foulger et al., 2003].

*Zircons:* Archaean and Jurassic zircons with Lewisian (1.8 Ga) and Mesozoic (~126 - 242 Ma) inheritance ages have been reported from lavas in NE Iceland. This has been interpreted as indicating ancient continental lithosphere beneath Iceland [Paquette et al., 2006; Schaltegger et al., 2002]. A continental composition for Icelandic-type lower crust can explain these results.

*Water:* Water in basalt glass from the mid-Atlantic Ridge indicates elevated contents in the source from ~61°N across Iceland [Nichols et al., 2002]. The water contents are estimated to be ~165 ppm at the southern end of the Reykjanes Ridge, rising to 620 - 920 ppm beneath Iceland. Such a component, and other volatiles such as CO₂ [Hole & Natland, this volume] decrease the solidus of a source rock and increase the volume of melt produced for a given T<sub>P</sub> (Section 6.2).

## 6.2 Temperature of the melt source

The temperature of the melt source of Icelandic rocks is too low to be able to account for a 30-40-km-thick basaltic crust using any reasonable lithology [Hole & Natland, this volume]. It is therefore an inevitable conclusion that much of the lower crust beneath the GIFR must arise from a process other than high-temperature partial melting of mantle peridotite.

Geochemical work aimed at determining the potential temperature of NE Atlantic source rocks has used basalts from Iceland and high-MgO picrites from the Davis Strait [Clarke & Beutel, 2019; Hole & Natland, this volume]. The T<sub>P</sub> for the source of MORB is generally used as the standard against which other calculated mantle temperatures are compared. The currently accepted value of this is 1350±40°C (Table 3) [Hole & Natland, this volume].
1007 A large range of temperatures, $T_P$ = 1400 - 1583°C, has been suggested for the mantle
1008 beneath Iceland [Hole & Millett, 2016; Putirka, 2008]. The breadth of this range in itself
1009 indicates how difficult it is to derive a repeatable, reliable $T_P$ using geochemistry and
1010 petrology. Difficulties include the lack of surface samples that correspond to an original
1011 mantle melt—crystalline rocks essentially always contain xenocrysts, and no picritic glass
1012 has been found in the NE Atlantic Realm [Presnall & Gudfinnsson, 2007]. The unknown
1013 source composition also introduces uncertainty. The geochemistry of Icelandic lavas requires
1014 there to be a component of recycled surface materials in the source and variable volatile
1015 contents including water [Nichols et al., 2002]. Ignoring any of these unknowns causes
1016 estimates of $T_P$ to be erroneously high.

1017 Crystallization temperatures estimated from olivine-spinel melt equilibration, the so-called
1018 "aluminum-in-olivine" method, are independent of whole-rock composition. The
1019 temperatures yielded by this method are $T_P$ ~1375°C and ~100°C higher for the Davis Strait
1020 picrites (Table 3) [Hole & Natland, this volume]. A summary of global maximum
1021 petrological estimates of $T_P$ and ranges of olivine-spinel equilibrium crystallization
1022 temperatures for magnesian olivine are shown in Figure 22.

1023 Petrological estimates of the potential temperature $T_P$ of the source of basalts in the NE
1024 Atlantic Realm suggest upper-bound $T_P$ of ~1450°C for Iceland and ~1500°C for the picrites
1025 of Baffin Island, Disko Island and west Greenland [Hole & Natland, this volume]. There may
1026 thus have been a short-lived, localized burst of magma from a relatively hot source lasting ~2
1027 - 3 Myr when propagation of the Labrador Sea spreading center was blocked at the
1028 Nagssugtoqidian orogen, but there is no compelling evidence for a $T_P$ anomaly > ~100°C
1029 before or after this anywhere in the NE Atlantic Realm.

1030 The melt volume produced at Iceland has also been used as a constraint in models for $T_P$.
1031 That work has assumed that the full thickness of the 30-40-km-thick seismic crustal layer is
1032 melt produced by steady state fractional melting of a peridotite mantle source. Production of
1033 just 20 km of igneous crust would require a $T_P$ of ~1450 - 1550°C assuming a damp or dry
1034 peridotite source [Sarafian et al., 2017]. No credible lithology or temperature can explain the
1035 crustal thickness of ~40 km that has been measured for central Iceland [Darbyshire et al.,
1036 1998a; Foulger et al., 2003].

1037 Crustal thickness beneath the active volcanic zones of Iceland varies from ~40 km (beneath
1038 Vatnajökull) to ~15 - 20 km (beneath the Reykjanes Peninsula extensional transform zone)
1039 [Foulger et al., 2003]. If the full thickness of crust everywhere is formed from melting in the
1040 mantle, unrealistically large lateral variations in temperature of the source of ~150°C over
1041 distances of ~125 km would be required [Hole & Natland, this volume].

## 6.3 $^3He/$$^4He$

1043 Elevated $^3He/$$^4He$ values are commonly assumed to indicate a core-mantle boundary
1044 provenance for the melt source. This association was originally suggested when it was found
1045 that some lavas from Hawaii contain high-$^3He/$$^4He$ [Craig & Lupton, 1976]. It was reasoned
1046 that, over the lifetime of Earth, the $^3He/$$^4He$ of the mantle has progressively decreased from
1047 an original value of ~200 times the present-day atmospheric ratio (Ra) to ~8 ± 2 Ra—the
1048 value most commonly observed in MORB. It was subsequently assumed that a lava with
1049 $^3He/$$^4He$ much larger than 8 Ra must have arisen from a primordial source, isolated for
1050 Earth's 4.6 Ga lifetime, deep in the mantle near the core-mantle boundary.
1051 This theory has long been contested and it has been counter-proposed that the helium instead
1052 resided for a long time in depleted, unradiogenic materials such as olivine in the sub-
1053 continental lithospheric mantle [Anderson, 2000a; b; 2001; Anderson et al., 2006; Foulger &
1054 Pearson, 2001; Natland, 2003; Parman et al., 2005]. That theory would fit the high-³He/⁴He
1055 values reported from Iceland and the Davis Strait [Starkey et al., 2009; Stuart et al., 2003] if
1056 the deeper parts of the crust beneath these regions contain ancient material, as we propose in
1057 this paper.

1058 **7 Discussion**

1059 **7.1 The Greenland-Iceland-Faroe Ridge**

1060 The model presented here proposes that in general Icelandic-type upper crust is mafic in
1061 nature, equivalent to Layers 2 - 3 of oceanic crust, whilst Icelandic-type lower crust is
1062 magma-dilated continental crust. The pre-existing SCLM mostly delaminated during the
1063 stretching process (Section 5) (Figure 18). The melt layer thus comprises Icelandic-type
1064 upper crust plus the melt that intruded into the continental crust below as plutons, dykes and
1065 sills. The location of Iceland with respect to the east Greenland and Faroe Volcanic margins
1066 fits the model of Geoffroy et al. [2015; submitted] of a dislocated C-block (Figure 5).

1067 This new model contributes to the > 40-year controversy regarding whether the crust beneath
1068 Iceland is thick or thin. A “thin crust” model, generally assumed in the 1970s and 1980s,
1069 attributed Icelandic-type upper crust to the melt layer—the subaerial equivalent of oceanic
1070 crust—and the layer currently termed “Icelandic-type lower crust” to hot, partially molten
1071 mantle [Björnsson et al., 2005]. From the 1990s, long seismic explosion profiles using
1072 modern digital recording were shot and deep reflecting horizons were discovered. A “thick
1073 crust” model was then introduced that interpreted the layer previously thought to be hot,
1074 partially molten mantle as Icelandic-type lower crust, the equivalent of oceanic layer 3, and
1075 part of the melt layer.

1076 Our findings support the thin-crust model with the caveat that Icelandic-type lower crust is
1077 indeed crust, and not hot mantle as previously proposed, but it is magma-inflated continental
1078 crust. This model agrees with long-sidelined magnetotelluric work in Iceland which detects a
1079 high-conductivity layer at ~10 - 20 km depth. This layer was proposed to mark the base of
1080 the crust [Beblo & Bjornsson, 1978; 1980; Beblo et al., 1983; Eysteinsson & Hermance,
1081 1985; Hermance & Grillot, 1974]. High-conductivity layers are common in continental mid-
1082 and lower crust [e.g., Muñoz et al., 2008]. Explosion seismology and receiver functions find
1083 the thickness of Icelandic-type upper crust to be ~3 - 10 km (Figure 9; Figure 10) [Darbyshire
1084 et al., 1998b; Foulger et al., 2003] which is comparable with the crustal thicknesses beneath
1085 the Reykjanes Ridge and the Kolbeinsey Ridge if additional magma dilating the Icelandic-
1086 type lower crust is taken into consideration. This is nevertheless up to ~40% thicker than the
1087 global average of 6 - 7 km. Mantle fusibility enhanced by pyroxenite and water (Section 6), a
1088 moderate elevation in temperature (Section 6.2), and bursts of volcanism accompanying
1089 frequent rift jumps (Section 4) can account for the enhanced melt volumes.

1090 The plate boundary traversing the GIFR cannot be likened to a conventional spreading ridge
1091 with segments connected by linear transform faults as is commonly depicted in simplified
1092 illustrations. Historically, motion in the GIFR region was postulated to have been taken up on
1093 a classic ~150-km-long sinistral transform fault named the Faroe Transform Fault or the
1094 Iceland Faroe Fracture Zone [Bott, 1985; Voppel et al., 1979] and this idea was reiterated in
1095 subsequent work [e.g., Blischke et al., 2017; Guarnieri, 2015]. Locations proposed for this
1096 feature include the north edge of the Iceland shelf, central Iceland, and the South Iceland
1097 Seismic Zone [Bott, 1974].

1098 There is, however, no observational evidence for such a structure [Gernigon et al., 2015;
1099 Schiffer et al., 2018] and it does not, even to a first order, fit the observations on the ground.
1100 Only a GIFR that deforms as a broad zone of distributed extension and shear can account for
1101 the reality of the geology of Iceland and adjacent regions [Schiffer et al., 2018].

1102 Our model may provide a long-awaited explanation for why the JMMC broke off east
1103 Greenland. Westerly migration of axes of extension on the GIFR may have changed the stress
1104 field in the diffusely extending continental area to the north and encouraged extension there
1105 to coalesce on the single most westerly zone which thereafter developed into the Kolbeinsey
1106 Ridge.

1107 7.2 Crustal flow

1108 Ductile crustal flow has been incorporated into earlier numerical models of continental
1109 breakup. A ductile, low-viscosity layer that decouples the upper lithosphere from the lower
1110 was incorporated in models of extending continental lithosphere by Huismans and Beaumont
1111 [2011; 2014]. Such a layer enables ultrawide regions of thinned, unruptured continental crust
1112 to develop along with distal extensional (sag) basins. Crustal thicknesses are maintained by
1113 widespread lateral flow of mid- and lower-crustal material from beneath surrounding regions.
1114 Lower crust may well up, further delaying full crustal breakup.

1115 In our model, subsidence resulting from progressive thinning or delamination of the mantle
1116 lithosphere is mitigated by hot asthenosphere rising to the base of the crust. This abruptly
1117 raises temperatures, increasing heat flow and further encourages ductile flow. Low extension
1118 rates, such as have characterized the NE Atlantic, tend to prolong the time to breakup and
1119 encourage diffuse extension because ductile flow and cooling can continue for longer. The
1120 crust may stretch unruptured for tens of millions of years and widen by 100s of kilometers
1121 with axes of extension migrating diachronously and laterally across the extending zone. Only
1122 after eventual rupture of the continental lithosphere can sea-floor spreading begin. Until that
1123 occurs, geochemical signatures of continental crust and mantle lithosphere are expected in
1124 overlying magmas that have risen through the continental material.

1125 Depth-dependent stretching, in particular involving the lower-crustal ductile flow that we
1126 model in Section 5, is both predicted by theory [McKenzie & Jackson, 2002] and required by
1127 observations from many regions. These include amagmatic margins, the Basin Range
1128 province, western USA [Gans, 1987] and deformation at collision zones, e.g., the Himalaya
1129 and Zagros mountain chains [e.g., Kusznir & Karner, 2007; Royden, 1996; Shen et al., 2001].
1130 Lower-crustal flow is actually observed where such crust is exhumed to the surface, e.g., at
1131 Ivrea in the Italian Alps, where lower-crustal granulite intruded by mafic plutons is exposed
1132 [e.g., Quick et al., 1995; Rutter et al., 1993].
1133 **7.3 Magmatism**

1134 **7.3.1 The concept of the North Atlantic Igneous Province**

1135 The issues laid out in this paper bring into question the concept that the magmas popularly grouped into the North Atlantic Igneous Province (NAIP) can be viewed as a single magmatic entity [Peace et al., this volume]. The NAIP is generally considered to include the volcanic rocks in the region of the Davis Strait, the volcanic margins of east Greenland and Scandinavia, and the magmatism of the GIFR. These magmas are, however, only a subset of those in the region and many others are not typically included [Peace et al., this volume].

1141 These include melt embedded in the “amagmatic” margins of SW Greenland and Labrador, current volcanism at Jan Mayen, the Vestbakken Volcanic Province ~300 km south of Svalbard, conjugates in NE Greenland [Á Horni et al., 2016], magmatism at the west end of the CGFZ [Keen et al., 2014] and basaltic sills offshore Newfoundland detected in ODP site 210-1276 that are thought to extend throughout an area of ~20,000 km² [Deemer et al., 2010]. It is illogical to exclude these, especially since the Cretaceous(?) Anton Dohrn and Rockall seamounts are included in the NAIP [Jones et al., 1994].

1148 The grouping of a select subset of magmas in the NE Atlantic Realm into a single province is predicated on and reinforces, the concept that they all arise from a single, generic source. A model of such simplicity that fits all observations has been elusive for over half a century.

1151 The obvious solution, and one that can readily account for the observations, is a model whereby each magmatic event occurs in response to local lithospheric tectonics and melts are locally sourced.

1154 The same reasoning may well apply to other volcanic provinces, e.g., the South Atlantic Volcanic Province. Generally included in this are the Paraná and Etendeka flood basalts, the volcanic rocks of the Rio Grande Rise and the Walvis Ridge, the currently active Tristan da Cunha archipelago and even kimberlites and carbonatites in Angola and the Democratic Republic of the Congo [see Foulger, 2018 for a review]. These volcanic elements contrast with one another in the extreme and each most likely erupted in reaction to local tectonic responses to global events and processes, with magmas locally sourced.

1161 **7.3.2 Magma volume**

1162 Estimates for the total volume of the magma generally lumped together as the NAIP are 2 - 10 x 10⁶ km³ with a value of ~6.6 x 10⁶ km³ for the north Atlantic volcanic margins [Eldholm & Grue, 1994a]. Assuming these margins formed in ~3 Myr, Eldholm and Grue [1994a] calculate a magmatic rate of 2.2 km³/a and suggest the NAIP is one of the most voluminous igneous provinces in the world. That calculation assumes that the HVLC beneath the *Inner* SDRs is all igneous and formed contemporaneously with the volcanic margins. If this is not the case, the volume and magmatic rate for the north Atlantic volcanic margins must be downward-revised by up to 30%, i.e. to ~4.4 x 10⁶ km³ for volume and 1.5 km³/a for magmatic rate. Eldholm and Grue [1994a] furthermore estimate a magmatic rate of ~ 0.2 km³/a for Iceland. If the igneous crust on the GIFR is only 10 - 15 km thick, this rate must be downward-revised to 0.12 - 0.08 km³/a. The magmatic rate per rift kilometer would then be 2 - 3 x 10⁻⁴ km³/a compared with ~4.8 x 10⁻⁴ km³/a per rift kilometer for the global plate boundary.

1175
1176 These changes reconcile geological estimates with those derived from numerical modeling.

1177 Magmatism at the NE Atlantic rifted margins has been simulated using models of

1178 decompression melting in a convectively destabilized thermal boundary layer coupled with

1179 upper-mantle (“small-scale”) convection [Geoffroy *et al.*, 2007; Mutter & Zehnder, 1988;

1180 Simon *et al.*, 2009]. These models explore whether the volumes and volume rates can be

1181 accounted for simply by breakup of the 100-200-km-thick lithosphere without additional *ad*

1182 hoc processes. Current numerical models slightly under-predict traditional geological

1183 estimates but could be reconciled with estimates lowered to take into account a wholly or

1184 partially continental affinity of HVLC.

1185 More accurate estimates of volume could also explain the extreme variations in magmatic

1186 thickness over short distances required by assumptions of HVLC igneous affinity. For

1187 example, the radical contrast between the unusually thin (4 - 7 km) oceanic crust beneath the

1188 Aegir Ridge [Greenhalgh & Kusznir, 2007] and a ~30 km igneous thickness beneath the

1189 adjacent GIFR defies reasonable explanation but the problem vanishes if the latter

1190 assumption is dropped.

1191 **7.3.3 The chevron ridges**

1192 Lithosphere- and asthenosphere-related mechanisms compete to explain the chevron ridges

1193 that flank the Reykjanes Ridge [Hey *et al.*, 2008; Jones *et al.*, 2002]. Martinez and Hey [this

1194 volume] suggest that the required oscillatory changes in magmatic production result from

1195 axially propagating mantle upwelling instabilities that travel with ridge-propagator tips along

1196 the Reykjanes Ridge. These originate in Iceland and the gradient in mantle properties along

1197 the Reykjanes Ridge results in the convective instabilities migrating systematically south

1198 along the Ridge. Upwelling is purely passive and the propagators behave in a wave-like

1199 manner without the flow of actual mantle material along the Ridge. In this model, the

1200 transition from linear to ridge/transform staircase plate boundary geometry at ~37 - 38 Ma

1201 failed to eliminate the structure of the deeper asthenospheric melting zone and the Reykjanes

1202 Ridge is restructuring itself to realign over that zone.

1203 Several of the propagators onset at the GIFR in concert with tectonic reorganizations there

1204 (Table 1) [Benediktsdóttir *et al.*, 2012] inviting consideration of lithospheric triggers. The

1205 Reykjanes Ridge as a whole is oblique to the direction of plate motion but its axis comprises

1206 an array of right-stepping *en echelon* spreading segments, each of which strikes perpendicular

1207 to the direction of motion. Such fabric resembles a left-lateral transtension zone.

1208 The diachronous chevron crustal fabric began to form at ~37 - 38 Ma when the Reykjanes

1209 Ridge changed from a linear to a ridge-transform configuration with a ~30° counter-

1210 clockwise rotation in the direction of plate motion (Section 2.2.2) [Gaina *et al.*, 2017]. From

1211 25-15 Ma slow, counter-clockwise rotation of the extension direction continued and from 15

1212 Ma - present it rotated back [Gaina *et al.*, 2017]. Slow counter-clockwise migration of the

1213 spreading direction would gradually hinder strike-slip motion on the transform segments and

1214 encourage evolution toward extension with a minor left-lateral shear overprint. Very slow

1215 changes in the direction of extension might be insufficient to trigger a sudden and major

1216 reorganization but enough to bring about the slow plate-boundary evolution observed.

1217 Regardless of whether a lithosphere- or asthenosphere-related mechanism is responsible for

1218 the chevron ridges, it is clear that shallow processes control them as their southerly

1219 propagation was temporarily blocked by several previously existing transform faults north of
1220 the present reorganization tip near the Bight transform fault. Furthermore, if their inception is
1221 related to tectonic reorganizations on the GIFR, then conversely the time at which the
1222 propagators set off from the GIFR could indicate the times of first-order tectonic
1223 reorganizations on the GIFR.

1224 The transform-eliminating rift propagators of the Reykjanes Ridge are unique in their degree
1225 of development but not entirely unknown elsewhere. Examples outside the NE Atlantic
1226 include a southward propagator eliminating a transform formerly at 21°40′N on the mid-
1227 Atlantic Ridge [Dannowski *et al.*, 2011] and propagators on the faster-spreading (~100
1228 km/Myr) NE Pacific plate boundary that eliminated the Surveyor, Sila, Sedna and Pau
1229 transforms [Atwater & Severinghaus, 1989; Hey & Wilson, 1982; Shih & Molnar, 1975].
1230 Propagating small-scale convective instabilities have also been postulated to form volcanic
1231 ridges and seamount chains that flank parts of the East Pacific Rise in a direction parallel to
1232 plate motion [e.g., Forsyth *et al.*, 2006].

1233 **7.3.4 The North Atlantic geoid high**

1234 The GIFR sits at the apex of a ~3000-km-long bathymetric and geoid high (up to ~4000 m
1235 and 80 m respectively) that stretches from the Azores to the Jan Mayen Fracture Zone
1236 [Carminati & Doglioni, 2010; King, 2005; Marquart, 1991]. Without this high the Thulean
1237 land bridge and Iceland would not have been subaerial. Globally, the only other comparable
1238 geoid high extends through Indonesia and Melanesia and to the Tonga Trench. Major geoid
1239 highs with lower amplitudes or smaller spatial extents are associated with the SW Indian
1240 Ocean and the Andean mountain chain.

1241 The geoid highs associated with Indonesia, Melanesia and Tonga, and the Andean mountain
1242 chain are a consequence of accumulations of dense subducted slabs. The geoid high of the
1243 north Atlantic corresponds closely to the pre-breakup Caledonian orogen plus the south
1244 European/North African Hercynian orogen (Figure 23). A possible explanation for part of the
1245 geoid anomaly is thus residual, dense, subducted Caledonian and Hercynian slabs along with
1246 continental lower crust and mantle lithosphere distributed in the shallow mantle. Henry Dick
1247 and colleagues have long argued that the petrology and geochemistry of magmas on the SW
1248 Indian ridge require SCLM in the melt source [Cheng *et al.*, 2016; Dick, 2015; Gao *et al.*,
1249 2016; Zhou & Dick, 2013]. That ridge is the current locus of extension between Africa and
1250 Antarctica which separated as part of Pangaea breakup beginning in the Jurassic. By analogy
1251 with the NE Atlantic, continental material might also remain in the mantle beneath the ocean
1252 there and the SW Indian geoid high might thus be explained in a similar way to that of the
1253 north Atlantic.

1254 **7.3.5 Regions analogous to the GIFR**

1255 There are clear parallels between the GIFR and the Davis Strait. The structure and tectonic
1256 development of the latter show similar characteristics to the GIFR but to a less extreme
1257 degree (Figure 24). The Davis Strait is colinear with the GIFR and both function as
1258 transtensional shear zones. Its primary feature is the long Ungava Fault Complex [Peace et
1259 al., 2017]. This is underlain by ~8 km of oceanic crust beneath which is ~8 km of HVLC
1260 with Vp up to 7.4 - 7.5 km/s [Chalmers & Pulvertaft, 2001; Funck et al., 2006; Funck et al.,
1261 2007; Srivastava et al., 1982] and density of 2850 - 3050 kg/m³ [Suckro et al., 2013]. These
1262 values are similar to those of Icelandic-type lower crust.
1263 Like the GIFR, the bathymetric high that contains the 550-km-long Davis Strait is elongated
1264 in the direction approximately perpendicular to plate motion. It has water depths of < 700 m,
1265 contrasting with the adjacent > 2000-m-deep Labrador Sea and Baffin Bay. At the Davis
1266 Strait north-propagating rifting stalled at the confluence of the Nagssugtoqidian and Rinkian
1267 orogens and continued displaced by several hundred kilometers in a right-stepping sense. In
1268 the case of the GIFR, both north- and south-propagating oceanic rifting stalled at the
1269 confluence of the Nagssugtoqidian and Caledonian orogens.

1270 The Jan Mayen Fracture Zone formed where a major, pre-existing transverse structure
1271 formed a barrier to the south-propagating Mohns Ridge. It was also an episodic transtensional
1272 structure, has a history of migration of the locus of deformation, bathymetric highs and
1273 unusual volcanism, e.g., on the island of Jan Mayen and in the submarine Traill Ø and Vøring
1274 Spur igneous complexes [Gernigon et al., 2009; Kandilarov et al., 2015]. Continental crust is
1275 possibly trapped between parallel segments of the Zone.

1276 7.3.6 Regions analogous to the NE Atlantic Realm

1277 The history, structure, tectonics and petrology of the NE Atlantic Realm are unusually
1278 complex but it represents an extreme example and not a unique case. Other regions that show
1279 similar features suggest that the style of breakup it exemplifies is generic. The NE Atlantic
1280 Realm may owe its extremity to the facts that the NE Atlantic was formed by two opposing
1281 propagators that stalled at a barrier, an unusually large microcontinent was captured, and the
1282 spreading rate was and is exceptionally slow.

1283 The South Atlantic Igneous Province also includes regions of shallow sea-floor, anomalously
1284 thick crust, anomalous volcanism and continental crust distributed in the ocean. It has a
1285 history of stalled spreading-ridge propagation, coincidence with a major pre-existing
1286 transverse structure and both shear and extensional deformation in a zone several hundred
1287 kilometers broad in the direction perpendicular to plate motion [Foulger, 2018; Kusznir & al.,
1288 2018]. Graça et al. [2019] recently presented evidence that the Rio Grande Rise, which
1289 contains continental material [Santos Ventura et al., 2019], and parts of the Walvis Ridge
1290 were once joined, but split apart by at least four ridge jumps. Such a process is very similar to
1291 that which we propose for the GIFR.

1292 The Lomonosov Ridge in the Arctic ocean can be viewed as an incipient microcontinent.
1293 West of India, the Laxmi basin comprises a pair of aborted conjugate volcanic passive
1294 margins with Outer SDRs that appear to be underlain by HVLC and flank an intra-oceanic
1295 microcontinent—a C-block [Geoffroy et al., submitted; Guan et al., 2019; Nemčok & Rybár,
1296 2017]. The Seychelles region in the West Indian Ocean, the Galapagos Islands region in the
1297 east Pacific [Foulger, 2010, p 100-101] and the Shatsky Rise [Korenaga & Sager, 2012;
1298 Sallares & Charvis, 2003] all display analogous features. Regions currently in the process of
1299 breaking up in a similar mode include the Afar area [Acton et al., 1991], the Imperial and
1300 Mexicali Valleys and Baja California (California and Mexico). The abundance of continental
1301 crustal fragments in the oceans is becoming increasingly clear, with much originating at
1302 locations where continental breakup was complicated by lithospheric heterogeneities.

1303 Despite the very different structure and context, tectonics comparable to those observed on
1304 the GIFR and in Iceland are also observed on the East Pacific Rise (EPR). There, “dueling”
1305 overlapping propagating ridge pairs with intermediate bookshelf shearing build ridge-
1306 perpendicular and ridge-oblique zones of crustal complexity (Figure 25) [Perram et al.,
1307 1993]. In oceanic settings overlapping ridge tips tend to form where lithosphere is weak and to migrate along-strike. Overlapping spreading centers are kinematically unstable and the tips inevitably fail episodically and are replaced by new ones. An unusual facet of the development of the GIFR that is not reported from the East Pacific Rise is the switching of the sense of overlap when the Aegir Ridge was replaced by the Kolbeinsey Ridge.

1312 Comparable styles of deformation are also observed in the Japan-, Manus-, Lau- and Mariana Trench back-arc basins [Kurashimo et al., 1996; Martinez et al., 2018; Taylor et al., 1994]. 1313 1314 Beneath back-arc basins the hydrous mantle environment above the dewatering slab does not become dehydrated and the attendant increase in viscosity tends to localize upwelling melt. 1315 1316 As a result, extension does not become focused in a single rift zone but remains distributed 1317 between multiple, parallel rifts. Magnetic anomalies are disorganized and water also reduces 1318 the solidus, increasing melt production [Dunn & Martinez, 2011; Martinez et al., 2018].

1319 On land, similar petrologies, including high-TiO₂ basalts, association with abundant rhyolite, 1320 and likely provenance of the source in subcontinental material are observed in flood basalts 1321 that erupted through continental lithosphere. These include the Central Atlantic Magmatic 1322 Province [Peace et al., this volume], the Deccan traps and the Columbia River Basalts.

1323 **8 Conclusions**

1324 Our main conclusions may be summarized:

1325 1. Disintegration of the Laurasian collage of cratons and orogens to form the Labrador Sea, 1326 Baffin Bay and the NE Atlantic Ocean lasted several tens of millions of years and 1327 occurred piecewise and diachronously via rift propagation.

1328 2. The GIFR formed where the south-propagating Aegir Ridge and the north-propagating, 1329 Reykjanes Ridge stalled at the junction of the Nagssugtoqidian and Caledonian orogens. 1330 The intervening ~300-km wide (northerly) and ~150-km long (easterly) continental 1331 block, the Iceland Microcontinent, along with flanking areas, extended by distributed, 1332 magma-assisted continental extension via multiple parallel migrating rifts with diffuse 1333 shear zones between them. The continental crust was capped by surface lavas. It stretched 1334 to form the 1000-km long Thulean continental land bridge which was not overrun by 1335 oceanic waters until ~10 -15 Ma.

1336 3. Magma-assisted continental extension was enabled by ductile flow of low-viscosity mid- 1337 and lower crust.

1338 4. Icelandic-type crust comprises the 3 - 10 km thick upper crust, equivalent to oceanic 1339 layers 2 - 3, underlain by lower crust up to ~ 30 km thick comprising magma-inflated 1340 continental crust.

1341 5. The melt layer that caps the GIFR comprises the Icelandic-type upper crust plus magma 1342 injected into the Icelandic-type lower crust, and has a total thickness of ~10 - 15 km.

1343 6. The petrology and geochemistry of Icelandic lavas is consistent with inclusion of a 1344 component from underlying continental crust.

1345 7. A largely continental Icelandic-type lower crust is consistent with the fact that no 1346 reasonable models of temperature or mantle petrology can generate the ~40 km of melt 1347 necessary to explain its entire thickness as wholly oceanic.

1348 8. The chevron ridges that flank the Reykjanes Ridge form in association with small-offset 1349 propagators initiated by tectonic reorganizations on the unstable GIFR.
1350 9. The GIFR tectonically decouples the oceanic regions to the north and south.

1351 10. The continuity of continental crust beneath the GIFR means that, at this latitude, Laurasia

1352 still has not yet entirely broken up. An implication of this is that the GIFR could be

1353 considered to be a new kind of plate boundary.

1354 11. A model whereby continental breakup is characterized by diachronous rifting, strong

1355 influence from pre-existing structures, distributed continental material in the new oceans,

1356 and anomalous volcanism matches many other oceanic regions.

1357

1358 Acknowledgments

1359 Erin K. Beutel and D. Barrie Clarke made valuable contributions to the discussions that led to

1360 this manuscript. CS's postdoctoral fellowship at Durham University was financed by the

1361 Carlsberg Foundation. F. Martinez and R. Hey are supported by US NSF grants OCE-

1362 1154071 and OCE-1756760. A. Peace's postdoctoral fellowship at Memorial University of

1363 Newfoundland was funded by the Hibernia project geophysics support fund. M. Stoker

1364 acknowledges the award of Visiting Research Fellow at the Australian School of Petroleum.

1365
1366 Table 1: Chronology of major tectonic events in the NE Atlantic. Timescale after Gradstein *et al.* [2012]. Chevron ridge # after Jones *et al.* [2002].

1368

<table><thead><tr><th>Date (Ma)</th><th>Chevron ridge #</th><th>Magnetic chron</th><th>Event</th></tr></thead><tbody><tr><td>58-57</td><td></td><td>C26</td><td>Beginning of opening of the Labrador Sea.</td></tr><tr><td>56-52</td><td></td><td>C24-22</td><td>First magnetic anomaly on the proto-Reykjanes Ridge.</td></tr><tr><td>54</td><td></td><td>C24r</td><td>Beginning of opening of the North Atlantic Ocean on the Aegir Ridge and west of the Lofoten margin.</td></tr><tr><td>54-ca 46</td><td></td><td>C24-21</td><td>Rift to drift transition, Faroe-Shetland and Hatton margins.</td></tr><tr><td>54.2-50</td><td></td><td></td><td>Spreading propagated from the Greenland Fracture Zone south to the Jan Mayen Fracture Zone.</td></tr><tr><td>52</td><td></td><td>C23</td><td>Aegir Ridge reaches its maximum southerly extent.</td></tr><tr><td>50-48</td><td></td><td>C21</td><td>~30-40° clockwise rotation of direction of plate motion.</td></tr><tr><td>48</td><td></td><td>C22-21</td><td>Onset of fan-shaped spreading about the Aegir Ridge. Pulse of extension in the southern JMMC. No major change south of the GIFR.</td></tr><tr><td>40</td><td></td><td>C18</td><td>Counter-clockwise rotation of direction of plate motion.</td></tr><tr><td>38-37</td><td>7</td><td>C17</td><td>Reykjanes Ridge becomes stair-step. First chevron ridge begins to form.</td></tr><tr><td>36</td><td></td><td>C13</td><td>Cessation of spreading in the Labrador Sea.</td></tr><tr><td>33-29</td><td></td><td>C12-10</td><td>Counter-clockwise rotation of direction of plate motion.</td></tr><tr><td>31-28</td><td>6</td><td>C12-10</td><td>Extinction of the ultra-slow Aegir Ridge. Second chevron ridge begins to form about the Reykjanes Ridge</td></tr><tr><td>24</td><td></td><td>C6/7</td><td>First unambiguous magnetic anomaly about the Kolbeinsey Ridge.</td></tr><tr><td>15-10</td><td></td><td>C5A/C5</td><td>Breaching of the Thulean land bridge.</td></tr><tr><td>14</td><td>5</td><td></td><td>Rift jump in Iceland from North West Syncline to Snæfellsnes Zone and Húnaflói Volcanic Zone, propagator "Loki" starts to travel south down Reykjanes Ridge forming third chevron ridge.</td></tr><tr><td>9</td><td>4</td><td></td><td>Propagator "Fenrir" starts to travel south down Reykjanes Ridge forming fourth chevron ridge.</td></tr><tr><td>7</td><td>3</td><td></td><td>Extinction of Snæfellsnes Zone, propagator "Sleipnir" starts to travel south down Reykjanes Ridge forming fifth chevron ridge.</td></tr><tr><td>5</td><td>2</td><td></td><td>Propagator "Hel" starts to travel south down Reykjanes Ridge forming sixth chevron ridge.</td></tr><tr><td>2</td><td>1</td><td></td><td>EVZ in Iceland forms, propagator "Frigg" starts to travel south down Reykjanes Ridge forming seventh chevron ridge.</td></tr></tbody></table>
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<table><caption>Table 2: Rift zones indicated by geological observations on land in Iceland.</caption><thead><tr><td>Name</td><td>Acronym</td><td>Tectonic status</td></tr></thead><tbody><tr><td>North West Syncline</td><td>NWS</td><td>Extinct</td></tr><tr><td>Austurbrún Syncline</td><td>AS</td><td>Extinct</td></tr><tr><td>East Iceland Zone</td><td>EIZ</td><td>Extinct</td></tr><tr><td>Snæfellsnes Zone</td><td>SZ</td><td>Oblique, non extensional</td></tr><tr><td>Húnaflói Volcanic Zone</td><td>HVZ</td><td>Extinct</td></tr><tr><td>Mödrudalsfjallgardar Zone</td><td>MZ</td><td>Extinct</td></tr><tr><td>Reykjanes Peninsula Zone</td><td>RPZ</td><td>Oblique, extensional</td></tr><tr><td>Western Volcanic Zone</td><td>WVZ</td><td>Active, waning</td></tr><tr><td>Hofsjökull Zone</td><td>HZ</td><td>Active, very short</td></tr><tr><td>Northern Volcanic Zone</td><td>NVZ</td><td>Active</td></tr><tr><td>Öræfajökull-Snæfell Zone</td><td>ÖVZ</td><td>Active, non extensional</td></tr><tr><td>Eastern Volcanic Zone</td><td>EVZ</td><td>Active, propagating</td></tr></tbody></table>

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Table 3: Potential temperatures required to produce 20 km of melt for various source compositions [from Hole & Natland, this volume].

1376

<table><thead><tr><th>Source composition</th><th><i>T</i><sub>P</sub> °C</th></tr></thead><tbody><tr><td>dry peridotite</td><td>1550</td></tr><tr><td>dry peridotite + 10% pyroxenite</td><td>1540</td></tr><tr><td>dry peridotite + 40% pyroxenite</td><td>1470</td></tr><tr><td>damp peridotite + pyroxenite</td><td>1450</td></tr><tr><td>damp peridotite</td><td>1450</td></tr><tr><td>pyroxenite</td><td>1325-1450</td></tr><tr><td>Baffin Island picrites (<i>T</i><sub>Ol-Sp</sub>)</td><td>1500</td></tr></tbody></table>

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<figure><img src="image_3.png" /></figure>

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Figure 1: Regional map of the North East Atlantic Realm showing features and places mentioned in the text. Bathymetry is shown in color and topography in land areas in gray.
BB: Baffin Bay, DAS: Davis Strait, DI: Disko Island, LS: Labrador Sea, CGFZ: Charlie-Gibbs Fracture Zone, BFZ: Bight Fracture Zone, RR: Reykjanes Ridge, IB: Iceland basin, DKS: Denmark Strait, SKI: Skaergaard intrusion, BC: Blosseville coast, KR: Kolbeinsey, Ridge, JMMC: Jan Mayen Microcontinent Complex, AR: Aegir Ridge, NB: Norway basin, WJMFZ, EJMFZ: West and East Jan Mayen Fracture Zones, JM: Jan Mayen, LB: Lofoten basin, VP: Vøring Plateau, VB: Vøring basin, MB: Møre basin, FI: Faroe Islands, SI: Shetland Islands, FSB: Faroe-Shetland basin, MT: Moine Thrust, GGF: Great Glen Fault, HBF: Highland Boundary Fault, SUF: Southern Upland Fault, MTFZ: Møre-Trøndelag Fault, Zone, HB: Hatton basin, RT: Rockall Trough, PB: Porcupine basin. Red lines: boundaries of the Caledonian orogen and associated thrusts, blue lines: northern boundary of the Hercynian orogen, both dashed where extrapolated into the younger Atlantic Ocean.
<figure><img src="image_4.png" /></figure>

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Figure 2: Schematic map of Greenland showing features referred to in the text. A: Arsuk, AS:

Ataa Sund, BS: Bylot Sund, DB: Disko Bugt, IBr: Inglefield Bredning, IBu: Inussulik Bugt,

IL: Inglefield Land, ITZ: Ikertoq thrust zone, KF: Karrat Fjord, KY: Kap York, M:

Midternæs, Na: Naternaq, NI: Nordre Isortoq, NS: Nordre Strømfjord, Nu: Nunatarsuaq, RB:

Rensselaer Bugt, S: Sermilik, T: Thule basin, TB: Tasiussaq Bugt, UF: Uummannaq Fjord

[from St-Onge *et al.*, 2009].
<figure><img src="image_5.png" /><figcaption>Figure 3: Reconstruction of the North Atlantic Realm at 1265 Ma. NAC: North Atlantic Craton, Nag: Nagssugtoqidian. Juvenile terranes: a: Sissimuit Charnockite, b: Arfersiorfic diorite, c: Ammassalik Intrusive Complex, d: South Harris Complex, e: Loch Maree Group, f: Lapland-Kola Granulite Belt, g: Tersk and Umba terranes [from Mason *et al.*, 2004].</figcaption></figure>

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<figure><img src="image_6.png" /></figure>

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Figure 4: Receiver function image of the crust and upper mantle under central east Greenland from Schiffer *et al.* [2014] showing the Central Fjord structure. A geological cross-section based on Gee [2015] is overlain showing the Caledonian nappes and foreland basin, and the Devonian basin. Younger sedimentary basins are from Schlindwein & Jokat [2000].

Extrapolated Moho depths are from Schiffer *et al.* [2016].
<figure><img src="image_7.png" /></figure>

Figure 5: Schematic diagram illustrating the generalized structure of Inner- and Outer-SDRs and a possible “C-Block” under Iceland. Outer-SDRs comprise thick subaerial eruptive layers underlain by hyper-extended middle crust and high-Vp mafic material of uncertain affinity but similar in structure to massively sill-intruded lower crust. Ductile flow and magma-assisted inflation can extend such crust to many times its original length. Material eroded from the underlying lithospheric mantle may be distributed in the direction of extension and incorporated in the underlying asthenosphere. LC1: sill-injected continent-derived ductile crust. LC2: highly reflective, undeformed layer, tectonically disconnected from LC1, and with much higher Vp (7.6-7.8 km/s) [adapted from Geoffroy et al., submitted].

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<figure><img src="image_8.png" /><figcaption>Figure 6: Magnetic anomalies in the North Atlantic Ocean [from Gaina *et al.*, 2017].</figcaption></figure>

1440

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1444
<figure><img src="image_9.png" /></figure>

Figure 7: Perspective view along the Reykjanes Ridge looking towards Iceland showing the flanking chevron ridges converging with the spreading axis. Fracture-zone traces delineating former transform faults that have been eliminated can be seen as oblique cross-cutting structures in the lower part of the figure. Submarine areas show satellite-derived Free Air gravity anomalies from Sandwell *et al.* [2014] with the land topography of Iceland superimposed.

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<figure><img src="image_10.png" /></figure>

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1459 Figure 8: The Greenland-Iceland-Faroe Ridge and surrounding areas showing bathymetry and tectonic features. JMMC: Jan Mayen Microcontinent Complex. Thick black lines: axes of Reykjanes and Kolbeinsey Ridges, thin gray lines on land: outlines of neovolcanic zones, dark grey: currently active extensional volcanic systems, dashed black lines: extinct rifts on land, thin black lines: individual faults of the South Iceland Seismic Zone (SISZ), white: glaciers. WVZ, EVZ, NVZ: Western, Eastern, Northern Volcanic Zones, TFZ: Tjörnes Fracture Zone comprising two main shear zones and one (dotted) known only from earthquake epicenters (see also Figure 15). Thick red lines: extinct rift zones from Hjartarson et al. [2017].

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<figure><img src="image_11.png" /><figcaption>Figure 9: Compilation of results from receiver function analysis in Iceland. Top: Depth to the base of the upper crust, middle: depth to the base of the lower crust, bottom: thickness of the lower crust [data from Foulger *et al.*, 2003].</figcaption></figure>

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<figure><img src="image_12.png" /><figcaption>Figure 10: Velocity-depth profiles showing the average one-dimensional seismic structure of Icelandic-type crust from explosion profiles shot in Iceland and in 10-Ma oceanic crust south of Iceland [Gebrande *et al.*, 1980]. Open-headed arrows, estimates of the base of the upper crust from various studies; solid-headed arrows, estimates of the base of the lower crust; M, proposed Moho identifications [from Bjarnason *et al.*, 1993; Flovenz, 1980; Foulger *et al.*, 2003; Staples *et al.*, 1997].</figcaption></figure>

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<figure><img src="image_13.png" /></figure>

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1491 Figure 11: Cenozoic tectonostratigraphy for the Faroe-Shetland basin. The compilation of the

1492 stratigraphy and Faroe-Shetland tectonics is based mainly on Stoker *et al.* [2013; 2018;

1493 2005b]. Additional information: ‘Stratigraphy’ and ‘Unconformity’ columns [Mudge, 2015],

1494 ‘Faroe-Shetland Tectonics’ column [Blischke *et al.*, 2017; Dean *et al.*, 1999; Ellis & Stoker,

1495 2014; Johnson *et al.*, 2005; Ólavsdóttir *et al.*, 2013a; Stoker *et al.*, 2012; Stoker *et al.*,

1496 2005a], timescale [Gradstein *et al.*, 2012].

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<figure><img src="image_14.png" /></figure>

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1501

1502 Figure 12: Schematic diagram illustrating the Iceland Microcontinent.
<figure><img src="image_15.png" /><figcaption>Figure 13: Locations of known extensional axes during opening of the NE Atlantic Ocean. Basemap and magnetic chrons (black lines) from GPlates using a Lambert Conformal Conic projection. Isochrons are from Müller *et al.* [2016]. Spreading ridges: Red—active, blue—extinct. Locations of some extinct offshore spreading axes are from Hjartarson *et al.* [2017] and Brandsdóttir *et al.* [2015]. Green: approximate boundary of Jan Mayen Microplate Complex. Areas where there is no direct evidence for rifts or spreading axes are left white.</figcaption></figure>

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<figure><img src="image_16.png" /><figcaption>Figure 14: Speculative reconstruction of the sequence of extensional deformation on the GIFR and surroundings. Outline of land areas and locations of known extensional axes are from Figure 13 with the latter shown as solid lines. Red: active, blue: extinct. Dashed lines show speculative positions of ridges at times when observational data are lacking. Green solid line: approximate boundary of Jan Mayen Microplate Complex. Pale blue dashed line: approximate boundary of Iceland Microcontinent. This, and the Jan Mayen Microplate Complex expand with time as a result of magma inflation and ductile flow.</figcaption></figure>

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<figure><img src="image_17.png" /></figure>

Figure 15: Map of Iceland from Einarsson [2008] showing earthquakes 1994-2007 from the database of the Icelandic Meteorological Office. Yellow: volcanic systems. The Tjörnes Fracture Zone comprises GOR: the Grímsey Oblique Rift, HFZ: the Húsavík-Flatey Zone, ER: the Eyjafjardaráll Rift, DZ: the Dalvík Zone. Other abbreviations are RR: Reykjanes Ridge, KR: Kolbeinsey Ridge, RPR: Reykjanes Peninsula Rift Zone (also known as the Reykjanes Peninsula extensional transform zone), WVZ: Western Volcanic Zone, SISZ: South Iceland Seismic Zone, EVZ: Eastern Volcanic Zone, CIVZ: Central Iceland Volcanic Zone, NVZ: Northern Volcanic Zone, SIVZ: South Iceland Volcanic Zone, Kr, Ka, H and L: the central volcanoes Krafla, Katla, Hengill and Langjökull, V: the Vestmannaeyjar archipelago.

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<figure><img src="image_18.png" /></figure>

1537
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1541 Figure 16: Similar to Figure 8 but showing additionally lines of curved sections of plate boundary that resemble curving, approaching crack tips (dashed magenta lines). Inset: expanded view of Vestmannaeyjar archipelago. Bold arrows: current direction of regional plate motion. For other details and abbreviations see caption of Figure 8.

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1547

HIMALAYA - TIBET

<figure><img src="image_19.png" /></figure>

1548

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1551 Figure 17: Schematic figure of the lithospheric structure of a well-studied, currently intact

1552 orogen—the Himalaya-Tibet orogen. Green: lithospheric mantle, red and pink: crust or

1553 intrusives, yellow and dark green: sedimentary basins. The orogen is underlain by an array of

1554 trapped fossil slabs that thicken the crust locally. Deeper parts of the slabs are in the dense

1555 eclogite facies and negatively buoyant [from Tapponnier *et al.*, 2001].

1556

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1558
<figure><img src="image_20.png" /></figure>

Figure 18: Simplified thermo-mechanical model of Cenozoic extension of the western frontal thrust of the Caledonian suture. Left panels: Lithology at selected times, thick black lines: minimum/maximum temperature profiles as a function of depth, thin black lines: isotherms from 1400°C with 100°C intervals. Upper left panel: initial model configuration. Central panels: density evolution, dashed black lines: isotherms from 0°C to 1400°C at 200°C intervals, full black lines: isotherms from 1450°C at 25°C intervals. Right panels: effective viscosity evolution.

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<figure><img src="image_21.png" /></figure>

Figure 19: Expanded view of the lithology panel for 50.6 Myr, from Figure 18. See that Figure for details.

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<figure><img src="image_22.png" /></figure>

Figure 20: Map view sketch of our model. Green: the Caledonian frontal thrust zone where the crust is relatively thick prior to breakup, arrows: lateral inflow of weak lower crust into the extending, thinning zone. The persistence of continental crust beneath the GIFR maintains a warm, weak lithosphere and encourages distributed deformation and lateral rift jumps to persist.
<figure><img src="image_23.png" /><figcaption>Figure 21: Cross section through the full-waveform inversion tomographic model of Rickers *et al.* [2013]. Colored dots are spaced at intervals of 1000 km. Compare with lower left panel of Figure 18.</figcaption></figure>

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<figure><img src="image_24.png" /><figcaption>Figure 22: Summary of global maximum petrological estimates of *T*<sub>P</sub> (diamonds ± 40°C; Herzberg & Asimow [2015]) and olivine-spinel equilibrium crystallization temperatures (T<sub>OI</sub>-S<sub>p</sub>) for magnesian olivine (dots). The lower light-blue shaded region represents the range of T<sub>OI</sub> for olivine which crystallized from near-primary magmas formed at ambient T<sub>P</sub> ~1350 ± 40°C (upper pink-shaded region). The horizontal dashed lines represent the maximum estimated T<sub>P</sub> for Iceland and West of Greenland (WoG; Disko Island, Baffin Island) from Hole and Natland [this volume]. Data sources for T<sub>OI-Sp</sub>: MORB, Gorgona komatiite and Archean komatiite: Coogan *et al.* [2014], British Palaeocene, Baffin Island, West Greenland (Disko Island): Coogan *et al.* [2014], Spice *et al.* [2016], Iceland: Matthews *et al.* [2016], Spice *et al.* [2016], Tortugal: Trela *et al.* (2017). Petrological estimates from Herzberg and Asimow [2008; 2015], Hole [2015], Hole and Millett [2016] and Trela *et al.* [2017].</figcaption></figure>

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<figure><img src="image_25.png" /></figure>

1612
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1615 Figure 23: Continents reassembled to 50 Ma with the location of the future NE Atlantic
centered over the present-day geoid high (red area). Thick white lines outline the Caledonian,
1616 Nagssugtoqidian, and Hercynian orogens. The area encompassed by these orogens is shaded
1617
1618 and corresponds to the majority of the region of the geoid high.
1619
1620

<figure><img src="image_26.png" /></figure>

1621
1622

1623 Figure 24: Structural map of the oceanic region west of Greenland showing Cretaceous basins and the extent of Paleogene volcanics, including inferred continuation as shown in Abdelmalak et al. [2018]. Different proposed continent-ocean boundaries are also shown.
1624
1625 The magnetic lineations and fracture zones are reproduced from Chalmers [2012]. BB: Baffin Bay, BI: Baffin Island, CFZ: Cartwright Fracture Zone, DS: Davis Strait, GR: Greenland, LFZ: Leif Fracture Zone, LS: Labrador Sea, UFZ: Ungava Fault Zone, LA: Labrador.
1626
1627
1628
1629 Elevation data are from Smith and Sandwell [1997].

1630
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1632

A

<figure><img src="image_27.png" /></figure>

B

<figure><img src="image_28.png" /></figure>

1633

<figure><img src="image_29.png" /></figure>

1634

1635

1636 Figure 25: Schematic diagrams showing models of spreading ridge evolution observed on the

1637 East Pacific Rise. Top panels: ridges in the region 26°S - 32°S between the Easter and Juan

1638 Fernandez microplates. Parallel, solid lines: active ridges, parallel dashed lines: extinct

1639 ridges. The structure is modeled as a brittle layer overlying and weakly coupled with an

1640 underlying ductile layer. Deformation in this layer is shown by shading with gray indicating

1641 uniform motion and white indicating little or no motion. Arrows show displacement.

1642 Extension occurs in the overlap zone on curved, overlapping ridges that progressively

1643 migrate outward, are removed from the magma supply, become extinct, and are replaced by

1644 new ridges. Distributed bookshelf faulting occurs in the overlap zone [from Martínez *et al.*,

1645 1997]. Bottom panels: Model for the evolution of the East Pacific Rise at 20°40'S showing a

1646 possible origin of rotated blocks. Shading indicates magnetization polarities. The ridge tips

1647 alternate between propagation and retreat, leading to the term “dueling propagators” [from

1648 Perram *et al.*, 1993].

1649

1650
1651

1652 *References*

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1654 North Atlantic Igneous Province, in *The North-East Atlantic Region: A Reappraisal of*
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1656 Pinvidic, J. Hopper, M. S. Stoker, C. Gaina, H. Doornenbal, T. Funck and U. Árting,
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1658 Abdelmalak, M. M., J. I. Faleide, S. Planke, L. Gernigon, D. Zastrozhnov, G. E. Shephard, and R. Myklebust (2017), The T-Reflection and the Deep Crustal Structure of the Vøring Margin, Offshore mid-Norway, Tectonics, 36, 2497-2523.

1661 Abdelmalak, M. M., S. Planke, S. Polteau, E. H. Hartz, J. I. Faleide, C. Tegner, D. A. Jerram, J. M. Millett, and R. Myklebust (2018), Breakup volcanism and plate tectonics in the NW Atlantic, Tectonophysics.

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