Magma Transport Pathways in Large Igneous Provinces: Lessons From

1 Magma Transport Pathways in Large Igneous Provinces: Lessons from

2 combining Field Observations and Seismic Reflection Data

3

4 1*Craig Magee ([email protected]), 2,3Richard E. Ernst

5 ([email protected]), 4James Muirhead

6 ([email protected]), 1Tom Phillips (Phillips, Thomas

7 ([email protected]), 1Christopher A-L Jackson ([email protected])

8

9 *corresponding author

10

11 1Basins Research Group, Department of Earth Science and Engineering, Imperial College

12 London, Prince Consort Road, London, SW7 2BP, UK

13 2Deparment of Earth Sciences, Ottawa, ON, K1S 5B6, Canada

14 3Faculty of Geology and Geography, Tomsk State University, Tomsk 634050, Russia

15 4 Department of Earth Sciences, Syracuse University, Syracuse, New York 13210, USA

16

17

18 Abstract

19 Large Igneous Province (LIP) formation involves the generation, intrusion, and extrusion of

20 significant volumes (typically >1 Mkm3) of mainly mafic magma and is commonly

21 associated with episodes of mantle plume activity and major plate reconfiguration. Within

22 LIPs, magma transport through Earth’s crust over significant vertical (up to tens of

23 kilometres) and lateral (up to thousands of kilometres) distances is facilitated by dyke swarms

24 and sill-complexes. Unravelling how these dyke swarms and sill-complexes develop is

25 critical to: (i) evaluating the spatial and temporal distribution of contemporaneous volcanism 26 and hydrothermal venting, which can drive climate change; (ii) determining melt source

27 regions and volume estimates, which shed light on the mantle processes driving LIP

28 formation; and (iii) assessing the location and form of associated economic ore deposits.

29 Here, we review how seismic reflection data can be used to study the structure and

30 emplacement of sill-complexes and dyke swarms. We particularly show that seismic

31 reflection data can reveal: (i) the connectivity of and magma flow pathways within extensive

32 sill-complexes; (ii) how sill-complexes are spatially accommodated; (iii) changes in the

33 vertical structure of dyke swarms; and (iv) how dyke-induced normal faults and pit chain

34 craters can be used to locate sub-vertical dykes offshore.

35

36 Introduction

37 Large Igneous Provinces (LIPs) comprise large volume (>0.1 Mkm3 but frequently >1

38 Mkm3), predominantly mafic (-ultramafic) intrusion networks, extrusive sequences (e.g.

39 flood basalts), and crustal magmatic underplate emplaced either over a relatively rapid

40 timeframe (<5 Myr) or through multiple, short-lived pulses over a few 10’s of millions of

41 years (Ernst 2014; Ernst and Youbi 2017; Coffin and Eldholm 2005; Coffin and Eldholm

42 1994; Bryan and Ferrari 2013; Bryan and Ernst 2008). LIPs occur in intraplate continental or

43 oceanic settings, and many appear to result from the impingement of a mantle plume at the

44 base of the lithosphere (e.g. Ernst 2014; Ernst et al. 1995; Coffin and Eldholm 1992).

45 Transport of magma through Earth’s crust during LIP development is facilitated by sheet-like

46 conduits including dykes, sills, and inclined sheets. Dykes are traditionally considered to be

47 the most important component of LIP magma plumbing systems and typically form

48 impressive swarms that can extend laterally for >2000 km (Fig. 1) (e.g. Ernst 2014; Ernst and

49 Baragar 1992; Ernst et al. 2001; Ernst et al. 1995; Halls 1982). Giant circumferential (sub-

50 circular to elliptical) swarms up to 2000 km in diameter have also been recently recognized 51 (e.g. Buchan and Ernst, 2018a, b). However, field observations reveal that sill-complexes, i.e.

52 extensive interconnected networks of sills and inclined sheets (e.g. the Siberian Trap sill-

53 complex is exposed over >1.5 × 106 km2), form the primary volumetric elements of many LIP

54 plumbing systems (Fig. 1) (e.g. Svensen et al. 2012; Leat 2008; Elliot 1992; Elliot and

55 Fleming 2004; Burgess et al. 2017). Analysing and integrating our understanding of these

56 two key parts of LIP plumbing systems (dyke swarms and sill-complexes) is crucial to: (i)

57 estimating LIP melt volumes, which provide critical insight into mantle processes; (ii)

58 reconstructing palaeogeographies, particularly during the Precambrian (e.g. Ernst et al. 2013;

59 Bleeker and Ernst 2006); (iii) locating economic exploration targets associated with LIP

60 formation (e.g. Ernst and Jowitt 2013; Jowitt et al. 2014); (iv) quantifying LIP-related

61 greenhouse gas production, and its role in driving mass extinction (e.g. Ernst and Youbi

62 2017; Svensen et al. 2004); and (v) understanding the emplacement mechanics and surface

63 expression of sheet intrusions, which can allow similar features to be identified and

64 interpreted on other planetary bodies (e.g. Head and Coffin 1997; Ernst 2014).

65 To date, our understanding of magma transport through LIP plumbing systems has

66 primarily been driven by the mapping of dyke swarm and sill-complex extents and

67 geometries using traditional field approaches and geophysical techniques (e.g. aeromagnetic

68 data), coupled with a range of petrological, geochemical, and geochronological analyses (see

69 Ernst 2014 and references therein). Furthermore, application of different modelling

70 approaches has helped evaluate emplacement mechanics of dyke swarms and sill-complexes

71 (e.g. Townsend et al. 2017; Rivalta et al. 2015; Bunger et al. 2013; Malthe-Sørenssen et al.

72 2004; Galland 2012; Kavanagh et al. 2018; Kavanagh et al. 2015). Combining these

73 techniques has provided profound insights into LIP melt conditions, magma evolution and

74 distribution, emplacement mechanics, age and duration of emplacement, and the tectonic and

75 climatic impacts of these processes (e.g. Svensen et al. 2012; Svensen et al. 2004; Cooper et 76 al. 2012; Ernst 2014; Jerram and Bryan 2015; Bryan et al. 2010; Ernst et al. 2005; Burgess et

77 al. 2017; Ernst and Youbi 2017). Despite the importance of understanding LIP dyke swarms

78 and sill-complexes, we know relatively little about their 3D structure because of the 2D

79 nature of Earth’s surface. This limitation means that both the lateral and vertical extents of

80 dyke swarms or sill-complexes can seldom be examined simultaneously. Furthermore, it is

81 commonly challenging to resolve individual intrusions and determine their connectivity using

82 geophysical data that capture the subsurface expression of these LIP plumbing systems (e.g.

83 gravity and magnetics). Without being able to visualise the true 3D geometry of LIP dyke

84 swarms and sill-complexes, it is difficult answer a wide range of questions, including: How

85 does magma travel such large lateral distances (>100 km)? Can magma transport channels

86 within LIP plumbing systems remain open and facilitate sustained flow for protracted periods

87 of time? How do dyke swarms and/or sill-complexes feed flood basalts? How volumetrically

88 important are intrusive networks within LIPs compared to their extrusive counterparts? What

89 is their relationship to underlying, deep crustal magmatic underplate? How do sheet

90 intrusions relate to stress field conditions?

91 Reflection seismology is the only technique that can image intrusion networks in 3D,

92 at a relatively high-resolution (i.e. metres to tens of metres) over vast areas, allowing

93 individual magma bodies and associated structures to be mapped (e.g. Smallwood and

94 Maresh 2002; Thomson and Hutton 2004; Planke et al. 2005; Cartwright and Hansen 2006;

95 Schofield et al. 2017; Magee et al. 2016; Planke et al. 2015). For example, mapping of

96 magma flow indicators across regionally extensive sill-complexes imaged in seismic

97 reflection data reveal that these intrusion networks can transport melt over significant vertical

98 and lateral distances (e.g. Magee et al. 2016; Schofield et al. 2017; Reynolds et al. 2017).

99 These magma flow observations are further supported by field, magnetic, and geochemical

100 analyses of LIP sill-complexes (e.g. Airoldi et al. 2016; Leat 2008; Aspler et al. 2002; Marsh 101 2004). One challenge with reflection seismology, which relies on measuring arrival times of

102 acoustic energy reflected back from geological features, is that the technique favours imaging

103 of sub-horizontal to moderately inclined structures (e.g. sills). Dykes and dyke swarms are

104 thus rarely imaged in seismic reflection data (e.g. Thomson 2007), except when they have

105 been rotated post-emplacement to shallower dips (e.g. Phillips et al. 2017; Abdelmalak et al.

106 2015). Several studies have also shown, however, that dykes can be detected in seismic

107 reflection data through the recognition of dyke-related host rock structures and/or

108 geophysical artefacts (e.g. Wall et al. 2010; Malehmir and Bellefleur 2010; Zaleski et al.

109 1997; Bosworth et al. 2015; Kirton and Donato 1985; Ardakani et al. 2017).

110 In addition to imaging the 3D structure of dyke swarms and sill-complexes, seismic

111 reflection data also provide unprecedented insight into the host rock deformation mechanisms

112 that spatially accommodate magma emplacement and associated hydrothermal venting (e.g.

113 Hansen and Cartwright 2006b; Magee et al. 2014; Magee et al. 2013; Reeves et al. 2018;

114 Trude et al. 2003; Jackson et al. 2013; Svensen et al. 2004; Hansen 2006). Being able to

115 examine how intrusion induces deformation and the style of structures formed is critical

116 because it provides insight into the processes driving magmatism (e.g. Pollard et al. 1983;

117 Ernst et al. 2001; Magee et al. 2017; Galland 2012; van Wyk de Vries et al. 2014; Pollard and

118 Johnson 1973). Furthermore, understanding how subsurface magma movement and

119 accumulation translates into ground deformation, through 3D analysis of ancient intrusions

120 and associated deformation, informs how surficial intrusion-induced structures on Earth and

121 other planetary bodies relate to the underlying plumbing system (e.g. Pollard et al. 1983;

122 Ernst et al. 2001; Magee et al. 2017; Galland 2012; Wyrick and Smart 2009; Wyrick et al.

123 2004; Manga and Michaut 2017).

124 We consider seismic reflection to be a powerful tool in the study of magma plumbing

125 systems, and one that is yet to be fully utilised by the LIP community. To highlight the 126 potential use of seismic reflection data, we briefly examine seismically imaged sill-

127 complexes and dyke swarms and discuss how insights into their 3D geometry inform our

128 understanding of LIP magma plumbing systems. We particularly aim to promote future

129 integration and collaboration between of sectors of the LIP community using reflection

130 seismology and those who apply more conventional analytical methods.

131

132 Magma transport pathways in LIPs

133 Here, we summarise our current understanding of the roles that sill-complexes and dyke

134 swarms play in the development of LIPs using selected examples from field and seismic

135 reflection studies. In addition to examining the 3D structure of sill-complexes and dyke

136 swarms, we establish how this knowledge can be combined with analyses of host rock and

137 free surface deformation to unravel the emplacement mechanics of these intrusive networks.

138 LIP plumbing systems also comprise local, geometrically complicated intrusive complexes

139 (e.g. laccoliths and layered intrusions) and, whilst these features are poorly imaged in seismic

140 reflection data (e.g. Archer et al. 2005), their development can be linked to dyke swarm and

141 sill-complex emplacement (e.g. Menand 2008); we therefore briefly describe insights into

142 these intrusive complexes derived from field observations. Overall, this short review of

143 previous work illustrates how seismic reflection data has advance our understanding of LIP

144 magma plumbing systems, and provides a key tool for future studies addressing unresolved

145 questions related to LIP emplacement.

146

147 Field observations of sill-complexes

148 Sill-complex geometry and connectivity

149 Prior to the application of seismic reflection data to study sill-complexes associated with

150 LIPs, our understanding of these plumbing system components was largely driven by 151 analyses of the Ferrar and Karoo sill-complexes (Figs 1 and 2A) (e.g. du Toit 1920; Gunn and

152 Warren 1962; Grapes et al. 1974; Chevallier and Woodford 1999; Leat 2008; Elliot and

153 Fleming 2017; Elliot et al. 1999; Svensen et al. 2015; Polteau et al. 2008b; Polteau et al.

154 2008a). The Karoo sill-complex is relatively well-exposed across ~0.5 × 106 km2 in South

155 Africa (Chevallier and Woodford 1999). The Ferrar sill-complex can be traced across an

156 apparently linear belt that, based on palaogeographic reconstructions of the Early Jurassic,

157 likely extended for >4000 km from the Theron Mountains in Antarctica to Australia and New

158 Zealand (Fig. 2A) (Elliot and Fleming 2017). Both of the Ferrar and Karoo sill-complexes

159 consist of numerous sills, which have tabular or saucer-shaped morphologies, and inclined

160 sheets intruded into relatively unstructured sedimentary basins at depths of <5 km (e.g. Figs

161 2B and C) (e.g. Elliot and Fleming 2017; Muirhead et al. 2014; Airoldi et al. 2011; Svensen

162 et al. 2012; Galerne et al. 2011; Galerne et al. 2008; Chevallier and Woodford 1999). High-

163 precision U-Pb zircon and baddeleyite dating of the Ferrar and Karoo sill-complexes yield

164 ages of 182.779±0.033–182.59±0.079 Ma and 183±0.5–182.3±0.6 Ma, respectively,

165 indicating emplacement likely occurred simultaneously over <0.5 Myr (Svensen et al. 2012;

166 Burgess et al. 2015). Geochronological and geochemical data suggest a common source for

167 both the Ferrar and Karoo sill-complexes, perhaps a plume centre located in the Weddell Sea

168 (Encarnación et al. 1996; Elliot and Fleming 2000; Burgess et al. 2015). Linear dyke swarms

169 and picrite distribution associated with the Karoo sill-complex suggest the province may have

170 alternatively been sourced from onshore southern Africa in the Mwenezi (Nuanetsi) area

171 (Fig. 2A) (Hastie et al. 2014; Burke and Dewey 1973; Ernst and Buchan 1997a). Unlike

172 many LIPs (e.g. the Mackenzie LIP; Ernst and Baragar 1992; Baragar et al. 1996), no major

173 sub-parallel or giant radiating dyke swarms are associated with the Ferrar LIP (Figs 1 and

174 2A) (e.g. Elliot and Fleming 2017, 2004; Muirhead et al. 2014). Here, we briefly discuss how 175 observations from the Ferrar and Karoo sill-complexes shed light on magma transport within

176 these extensive LIP plumbing systems.

177 The Ferrar sill-complex is spectacularly exposed in the McMurdo Dry Valleys of

178 Antarctica, where networks of sills and inclined sheets can be observed and measured on >1

179 km-high cliff faces (Fig. 2B). These outcrops reveal a conduit system of interconnected,

180 stacked sills that facilitated magma ascent and possibly fed voluminous lava outpourings

181 (Elliot and Fleming 2004; Muirhead et al. 2014; Muirhead et al. 2012). Although these cliff

182 face outcrops limit access to the 3D geometry of individual sills, extensive exposures along

183 much of the Ferrar sill-complex length mean that magma transport within the system can be

184 regionally assessed (Fig. 2A) (Elliot and Fleming 2017). Key observations from the Ferrar

185 sill-complex include: (i) recognized sill-sill feeding relationships (e.g. Fig. 2B) (Muirhead et

186 al. 2012; Muirhead et al. 2014); (ii) decreasing Mg# and MgO contents consistent with

187 fractional crystallization away from the Weddell Sea region (Leat 2008; Elliot and Fleming

188 2000; Elliot et al. 1999); and (iii) the consistent orientation of broken bridge structure long

189 axes and magnetic lineations within sills and interconnected sheets across Antarctica (e.g. in

190 the Theron Mountains, Antarctica) (Hutton 2009; Dragoni et al. 1997; Airoldi et al. 2012).

191 These observations indicate that interconnected sills facilitated long distance magma

192 transport (e.g. Leat 2008; Elliot and Fleming 2017). However, it remains contentious as to

193 whether magma: (i) entered the sill-complex only at the plume head region (i.e. magma

194 flowed within the sill-complex for >4000 km; Leat 2008); or (ii) if dykes injected laterally

195 from the plume head region within the crust and periodically ascended to feed the sill-

196 complex (Elliot and Fleming 2017).

197 In contrast to the Ferrar sill-complex, exposure styles within the Karoo sill-complex

198 commonly allow the broad 3D geometry of individual intrusions to be assessed (e.g. Fig. 2C).

199 At deeper levels within the sill-complex, the sills are up to 200 m thick and tabular, with 200 those intruded closer to the palaeosurface, considered to be marked by the base of the

201 Drakensburg Group lavas, displaying saucer-shaped morphologies with thicknesses <100 m

202 (e.g. Fig. 2C) (Svensen et al. 2012); intrusion of these sills is predicted to have released at

203 least 27,400 Gt CO2 from hydrothermal vents, contributing to Toarcian global warming (e.g.

204 Svensen et al. 2007; Svensen et al. 2006; Jamtveit et al. 2004). The change in sill geometry

205 with proximity to the palaeosurface likely reflects the increasing interaction between the local

206 stress field around the magma body and the free surface (Malthe-Sørenssen et al. 2004). It is,

207 however, commonly difficult from field observations alone to determine how individual sills

208 are fed, i.e. are sills interconnected or are they fed by dykes? For example, within the vicinity

209 of the Golden Valley Sill, it may be assumed that the close spatial association of five major

210 sills and one major dyke indicates that the intrusions are interconnected (Fig. 2C) (Galerne et

211 al. 2011). However, Forward Stepwise-Discriminant Function Analysis using whole rock

212 major and trace element geochemical data, which distinguishes compositional populations,

213 indicates that different magma pulses can be tracked across the intrusions (Fig. 2C) (Galerne

214 et al. 2011). Three sills appear connected and fed by a single magma batch (i.e. the Golden

215 Valley, MV, and Glen sills), whereas the Morning Sun Sill, Harmony Sill and Golden Valley

216 Dyke are not connected and were fed by discrete magma batches (Fig. 2C) (Galerne et al.

217 2011). Furthermore, the compositional patterns are consistent with each of these magma

218 batches being derived from a common magma reservoir that becomes progressively more

219 contaminated by crustal host rocks (Neumann et al. 2011).

220

221 Surface expression of sills and sill-complexes

222 The growth of broadly sub-horizontal, tabular sills and laccoliths emplaced at shallow-levels,

223 particularly in sedimentary basins, is typically accommodated by overburden uplift (e.g. Fig.

224 3) (e.g. Jackson and Pollard 1988; Koch et al. 1981; Johnson and Pollard 1973; Pollard and 225 Johnson 1973; Morgan et al. 2008; Wilson et al. 2016; Galland 2012; Agirrezabala 2015; Le

226 Gall et al. 2010). This doming of the overburden, termed forced folding, commonly instigates

227 deformation of the free surface (e.g. Fig. 3) (e.g. Magee et al. 2017; Pollard and Johnson

228 1973). Although erosion and exposure of ancient sill-complexes involves removal of the

229 overburden, making it difficult to determine how entire sill-complexes are accommodated,

230 geodetic and remote sensing data from the Danakil Depression, Ethiopia document the

231 growth of forced folds above several sill intrusions (e.g. Fig. 3) (Magee et al. 2017). In

232 particular, analysis of ground deformation at and lava flow disposition around the Alu dome,

233 in the Danakil Depression, indicate that sill growth can occur over protracted timespans

234 through discrete injection and eruption cycles driving uplift and subsidence (Fig. 3) (Magee

235 et al. 2017). Identifying forced folds at the surface thus allows the geometry and locations of

236 subsurface sills and sill-complexes to be inferred (Magee et al. 2017; van Wyk de Vries et al.

237 2014). It should, however, be noted that inelastic deformation of the host rock can also

238 accommodate sill emplacement (e.g. fluidisation), which can partly or fully inhibit expression

239 of intrusions at the contemporaneous surface (e.g. Schofield et al. 2012c; Schofield et al.

240 2014).

241

242 Imaging and identifying sill-complexes in seismic reflection data

243 Seismic reflection data have revolutionized our understanding of sill-complexes because

244 these intrusion networks can be imaged in 3D at resolutions on the order of tens of meters

245 (e.g. Magee et al. 2014; Smallwood and Maresh 2002; Cartwright and Hansen 2006;

246 Thomson and Hutton 2004; Planke et al. 2015). Mafic sills are typically well imaged in

247 seismic reflection data, particularly within sedimentary basins, where they are expressed as

248 positive polarity, high-amplitude reflections (Fig. 4) (e.g. Smallwood and Maresh 2002;

249 Thomson 2005; Planke et al. 2005; Planke et al. 2015). This high-amplitude, positive polarity 250 reflection configuration occurs because the high density and seismic velocity of mafic

251 igneous rocks, relative to the host sequence, produces a high acoustic impedance contrast at

252 intrusion contacts that reflects more acoustic energy (e.g. Brown 2004; Magee et al. 2015;

253 Eide et al. 2017); more felsic intrusions have lower densities and seismic velocities and thus,

254 depending on the host rock lithology, may not be easy to identify in seismic reflection data

255 (Mark et al. 2017). Mapping sills in seismic reflection data reveals that they can display a

256 range of geometries, from strata-concordant to saucer-shaped (Fig. 4A-C) (e.g. Planke et al.

257 2005). These sill-related reflections commonly display minor vertical offsets that radiate

258 outwards from the deepest portion of the sill (Fig. 4D), or where the sill appears connected to

259 an inclined sheet (e.g. Magee et al. 2014; Schofield et al. 2012a; Reynolds et al. 2017).

260 Comparisons between sills observed in outcrop and reflection seismic reveal that these minor

261 vertical offsets are likely intrusive steps or bridge structures (e.g. Figs 4A and D), which form

262 during and shed light on initial sheet propagation direction (Schofield et al. 2012c). Mapping

263 these seismically resolvable magma flow indictors thus provides critical insights into how

264 entire sill-complexes are emplaced (Schofield et al. 2017; Magee et al. 2014; Magee et al.

265 2016).

266 One limitation of seismic reflection data is that sills are typically imaged as tuned

267 reflection packages, whereby the discrete reflections emanating from the top and base of a sill

268 converge on their return to the surface and cannot be distinguished (Fig. 4) (Smallwood and

269 Maresh 2002; Thomson 2005; Magee et al. 2015; Rabbel et al. 2018; Eide et al. 2017). This

270 ‘tuning’ of sill reflections is controlled by the vertical resolution of the seismic reflection

271 data, itself a function of the dominant seismic frequency and seismic velocity of the host rock

272 interval (Brown 2004), and means that the true sill thickness cannot be extracted. Instead, the

273 maximum and minimum thickness of intrusions producing tuned reflection packages are

274 defined by the vertical resolution (or limit of separability) and the detection limit (or limit of 275 visibility), respectively (Brown 2004; Magee et al. 2015; Smallwood and Maresh 2002; Eide

276 et al. 2017; Rabbel et al. 2018). Intrusions with thicknesses above the vertical resolution of

277 the seismic reflection data will produce discrete top and base reflections.

278

279 North Atlantic Igneous Province (~61–50 Ma): Irish Rockall Basin

280 The North Atlantic Igneous Province formed ~61–50 Ma in response to the impingement of a

281 mantle plume (i.e. now centred on Iceland) and break-up of Pangaea (e.g. Storey et al. 2007).

282 Here, we specifically focus on a sill-complex imaged in seismic reflection data from the NE

283 Irish Rockall Basin and particularly consider how its emplacement deformed the host rock

284 and contemporaneous surface (Figs 1 and 5). Within a 748 km2 region of the sill-complex,

285 which is covered by 3D seismic reflection data, the imaged plumbing system comprises 82

286 seismically resolved, stacked saucer-shaped sills and inclined sheets (Figs 5A and B) (Magee

287 et al. 2014). Overall, the sill-complex was intruded into a ~2 km thick sequence of

288 Cretaceous marine shale and Paleogene volcaniclastic sandstone (Magee et al. 2014).

289 Individual intrusions have diameters of ~0.3–7.1 km and may transgress up to 1.79 km of

290 strata (Magee et al. 2014). Mapping of magma flow indicators (e.g. intrusive steps and bridge

291 structures; Fig. 4D) and comparison to vertical seismic sections reveal that individual

292 intrusions propagated away from apparent connection sites to underlying sills (Magee et al.

293 2014). These inferred flow patterns imply that individual sills are connected and facilitated

294 lateral magma flow, from a melt source to the NW likely corresponding to the Hebridean

295 Terrace Igneous Complex (i.e. an intrusive complex), across a broad area without dykes

296 (Magee et al. 2014).

297 The Irish Rockall Basin sill-complex is overlain by a series dome-shaped forced folds

298 (e.g. Fig. 5A) (Magee et al. 2014). Two types of forced fold can be recognized (Magee et al.

299 2014): (i) individual forced folds, the outline of which corresponds to the lateral tips of 300 underlying sills; and (ii) compound forced folds, which consist of and result from the

301 amalgamation of individual forced folds (Fig. 5C). These compound forced folds indicate that

302 the entire sill-complex was, at least partly, accommodated by overburden uplift (Magee et al.

303 2014). Importantly, throughout the folded Cretaceous-to-Lower Eocene sequence and

304 particularly at the Top Palaeocene, seismic-stratigraphic onlap of strata onto the forced folds

305 is observed (Fig. 5A) (Magee et al. 2014). These onlap relationships form when sediment is

306 deposited around a topobathymetric high, demonstrating that the forced folds were expressed

307 at the contemporaneous surface (Magee et al. 2014). Furthermore, biostratigraphic dating of

308 the onlapping strata indicates that forced fold growth, and thus sill emplacement, occurred

309 incrementally over 15 Myr between ~65–54 Ma; prominent onlap onto the Top Palaeocene

310 horizon suggests a major pulse of sill-complex emplacement occurred at ~56 Ma (Fig. 5A)

311 (Magee et al. 2014). Recognition of these forced folds and onlap onto them indicates that the

312 surface expression of sill-complexes may be complicated, with individual injection events

313 shaping the deformation history of a broad area over a protracted period of time (e.g. Fig. 5C)

314 (Magee et al. 2014). The geometry and history of intrusion-induced forced folds currently

315 expressed at the surface of Earth or other planetary bodies may thus not relate to evolution of

316 a single intrusion, as is commonly assumed (Magee et al. 2014).

317

318 North Atlantic Igneous Province (~61–50 Ma): Faroe-Shetland Basin

319 Within the Faroe-Shetland Basin, to the north of the Rockall Basin, seismic reflection data

320 image a ~61–52 Myr old, extensive sill-complex (>22,500 km2) emplaced into Upper

321 Cretaceous shale (Figs 1 and 6) (Passey and Hitchen 2011; Schofield et al. 2017). Over 300

322 seismically resolved saucer-shaped sills, with diameters up to tens of kilometers, have been

323 mapped (Fig. 6) (Schofield et al. 2017). Reconstruction of flow patterns within the Faroe-

324 Shetland Basin sill-complex implies that magma was fed into the basin via dykes intersecting 325 basement-involved faults, from which sills propagated laterally for up to ~25 km into the

326 hanging wall (Fig. 6) (Schofield et al. 2017). Individual intrusions transgress up to 2 km of

327 the sedimentary succession and appear to feed overlying sills and lava flows (Fig. 6B)

328 (Schofield et al. 2017).

329

330 North Atlantic Igneous Province (~61–50 Ma): Vøring and Møre basins

331 Similar to the Irish Rockall and Faroe-Shetland basins, the Vøring and Møre basins offshore

332 Norway host an extensive sill-complex (i.e. >80,000 km2) emplaced during the break-up of

333 the North Atlantic (Fig. 7A) (e.g. Jamtveit et al. 2004; Skogseid et al. 1992; Planke et al.

334 2005; Schmiedel et al. 2017; Cartwright and Hansen 2006; Hansen and Cartwright 2006b, a;

335 Hansen et al. 2004). Associated with the sill-complex are >735 hydrothermal vent

336 complexes; i.e. pipe-like structures that emanate from the lateral tips of sills and extend

337 upwards to the syn-emplacement seabed, where they fed crater-, dome-, or eye-shaped vents

338 (e.g. Fig. 7B) (e.g. Hansen 2006; Svensen et al. 2004). Biostratigraphic dating suggests the

339 majority of the hydrothermal vent complex was active 55.8–55 Ma, coincident with the onset

340 of the 200 Kyr-long Eocene Thermal Maximum global warming event (Svensen et al. 2004).

341 It is estimated that sill-complex emplacement heated host rock organic shales, rapidly

342 releasing 3–0.3 × 1018 g of methane through the hydrothermal vent complex and thereby

343 instigating a global carbon isotope excursion of -1‰, which may have promoted the Eocene

344 Thermal Maximum (Svensen et al. 2004).

345

346 Warakurna LIP (~1070 Ma): Northern Yilgarn Craton, Western Australia

347 Within the ~3.05–2.62 Ga Youanmi Terrane in the Yilgarn Craton of Western Australia (Fig.

348 1), three 2D seismic reflection profiles image an interconnected network of highly reflective

349 sills that extend for several hundreds of kilometers (e.g. Fig. 8) (Ivanic et al. 2013a). In 350 contrast to the vast majority of sill-complexes described from seismic reflection data, which

351 occur in sedimentary basins (see Magee et al. 2016 and references therein), these sills within

352 the Youanmi Terrane were emplaced into crystalline rocks comprising greenstone belts,

353 granites, and granitic gneiss (Fig. 8) (Wyche et al. 2013). Radiometric dating of outcropping

354 intrusions within this sill-complex, including the Mount Homes Gabbro Sill that has a

355 baddeleyite SHRIMP U-Pb age of 1070 ± 18 Ma, suggests emplacement occurred during

356 formation of the ~1070 Ma Warakurna LIP (Wyche et al. 2013; Ivanic et al. 2013a; Ivanic et

357 al. 2013e; Wingate et al. 2004; Pirajno and Hoatson 2012). Basement sills are also potentially

358 associated with other LIP events, such as the ~1080 Ma Southwestern Laurentia LIP, the

359 ~1270 Ma Mackenzie LIP, and the ~1740 Ma Cleaver LIP (e.g. Mandler and Clowes, 1997;

360 Welford et al. 2004; Ernst, 2014). Although the relatively low resolution of the seismic

361 reflection data imaging this Warakurna sill-complex mean that its overall geometry and

362 contemporaneous surface expression cannot fully be assessed, these data indicate that

363 interconnected sill-complexes can facilitate long-distance lateral magma transport within

364 crystalline continental crust (Fig. 8) (Ivanic et al. 2013a; Ivanic et al. 2013e; Magee et al.

365 2016).

366

367 Field observations of dyke swarms

368 Dyke swarm geometry and connectivity

369 Key to the mapping of dyke swarms has been the combination of field observations and

370 regional and local, high-resolution aeromagnetic maps (e.g. Ernst 2014; Cooper et al. 2012;

371 Reeves 2000; Glazner et al. 1999; Buchan and Ernst 2013; Buchan and Ernst 2004).

372 Aeromagnetic mapping of dyke swarms has proved particularly important because these data

373 allow dykes to be traced beneath areas of thin sedimentary cover (Fig. 9) (e.g. Cooper et al.

374 2012; Mäkitie et al. 2014; Buchan and Ernst 2004). An excellent example of how 375 aeromagnetic data can illuminate otherwise unseen dyke swarm extents and properties is

376 provided by the Tellus regional airborne geophysical survey of Northern Ireland, which

377 imaged several Palaeocene dyke swarms related to the formation of the North Atlantic

378 Igneous Province (Fig. 9) (Cooper et al. 2012). Subtle changes in dyke orientation, cross-

379 cutting relationships, and variable offset across strike-slip faults revealed by the Tellus

380 aeromagnetic data highlight dyke swarm emplacement occurred periodically during north-

381 south Alpine-related compression (Fig. 9) (Cooper et al. 2012). Overall, field and

382 aeromagnetic mapping reveal that individual dykes within dyke swarms can display a range

383 of lengths, from metres to thousands of kilometres, and thicknesses of typically 10–50 m, but

384 up to >100 m and >1 km in the case of dyke-like layered intrusions. It should, however, be

385 noted that it can be challenging to determine dyke lengths if dykes are segmented and linked

386 across en echelon offsets; i.e. the actual length of the dyke could be much greater than the

387 length of measured segments (e.g. Ernst 2014; Ernst et al. 1995; Townsend et al. 2017). In

388 particular, whilst the en echelon arrangement of dyke segments is clear in some swarms (e.g.

389 700 km long 1140 Ma Great Abitibi dyke of the Superior craton), connections between

390 segments can be difficult to recognise within dense swarms.

391 Geochemical and palaeomagnetic studies suggest that individual dykes within a

392 swarm may represent a single, unique magma pulse (Ernst 2014). For example, the 700 km

393 long 1140 Ma Great Abitibi dyke of the Superior craton displays distinct and internally

394 consistent ratios of mantle incompatible elements (i.e. expressed as parallel curves on multi-

395 element diagrams), which are slightly but clearly different from the shapes of multi-element

396 diagrams for other nearby dykes within the swarm (Ernst 2014). It is likely that individual

397 dykes within a swarm are injected from a magma reservoir located in the plume centre

398 region, probably a layered mafic-ultramafic intrusion, in separate injection events (e.g.

399 Baragar et al. 1996; Blanchard et al. 2017). 400 Whilst sill-complexes are typically restricted to sedimentary basins, dykes and dyke

401 swarms can vertically span lower- to supra-crustal structural levels and extend along-strike

402 across various rock types. Dyke swarm geometries are thus primarily controlled by regional

403 stress conditions, whereby dykes typically orient orthogonal to σ3 and in the σ1-σ2 plane. A

404 range of overall dyke swarm geometries have been identified (Fig. 10): (i) giant radiating

405 swarms that have radii up to 2000 km, diverge away from the plume centre, and attain widths

406 of hundreds of kilometres (Ernst et al. 2001; Ernst and Buchan 1997a); (ii) linear dykes

407 swarms, which are commonly associated with overlying rift zones, extend along-strike for up

408 to >2000 km, and have widths of tens of kilometres (e.g. Ernst and Buchan 1997a; Ernst,

409 2014); and (iii) giant circumferential swarms that have diameters of up to 2000 km (Buchan

410 and Ernst, 2018a, b). All three dyke swarm geometries can develop around and appear to

411 emanate from a mantle plume centre (Ernst et al. 2001). Although the plan-view geometry of

412 a dyke swarm can be assessed at the level of exposure and under thin cover using field

413 observations and remote sensing data, it is difficult to estimate dyke height and determine

414 whether dyke swarm width varies with height. Modelling by Wilson and Head (2002)

415 suggests that the bottom of laterally emplaced dykes could extend into the upper lithospheric

416 mantle. However, this interpretation differs from studies that suggest a lower density of dykes

417 emplaced in high grade terranes, compared lower grade terranes, probably reflects the

418 presence of a lower boundary to laterally propagating dykes within the lower crust (Fahrig,

419 1987). Here, we provide a brief summary of several major dyke swarms.

420 Linear dyke swarms for many LIPs develop when uplift induced by mantle plume

421 impingement localise mechanical (i.e. faulting) and dyke-driven extension of Earth’s crust

422 within narrow rift zones that extend away from the plume centre, potentially forming a rift-

423 rift-rift triple junction (Fig. 10A). For example, a major linear dyke swarm parallels the Red

424 Sea rift for ~2500 km (Fig. 1) (Bosworth et al. 2015). The dyke swarm and rifting are 425 associated with the Afro-Arabian LIP, which is in turn related to the Afar plume (Fig. 1). The

426 ~30 Ma Ethiopian flood basalts and related magmatism was focused above the inferred

427 location of the Afar plume and triple junction rifting extends along the Red Sea, the Gulf of

428 Yemen, and southward along the East Africa rift system (Ebinger and Sleep 1998). Initial

429 flood basalt magmatism occurred at 30 Ma (Hofmann et al. 1997), but the major pulse of rift-

430 parallel diking in the Red Sea occurred at ~23 Ma (e.g. Bosworth et al. 2015).

431 If extension at the level of dyke emplacement is instead accommodated solely by

432 intrusion, giant radiating dyke swarms may form, whereby dyke-spacing increases away from

433 the plume centre (e.g. Fig. 10B). The 1267 Ma Mackenzie dykes exhibit a radiating pattern

434 that spans about 100° arc and extends at least 2400 km from the inferred plume centre at the

435 focus of the swarm (Fig. 1) (e.g. Baragar et al. 1996; Buchan and Ernst 2013; Ernst and

436 Bleeker 2010). Average dyke widths are ~30 m and magnetic fabric measurements

437 demonstrate vertical magma flow dominated within ~500 km of the plume centre (Ernst and

438 Baragar, 1992). All dykes analysed beyond ~500 km from the plume centre (i.e. sample sites

439 at 600, 800, 1000, 1400 and 2100 km) display magnetic fabrics consistent with lateral

440 injection (Ernst and Baragar 1992). Extrusive lavas of the Coppermine River flood basalts are

441 only present in isolated remnants and their original extent remains uncertain (Ernst, 2014).

442 Associated sills are present in scattered sedimentary basins and it is predicted that some of

443 these sill-complexes and also layered mafic intrusions were fed by the radiating dyke swarm

444 (Ernst and Buchan, 1997b).

445 Giant circumferential dyke swarms likely originate from and define the outer edge of

446 the mantle plume head, which flattens against the base of the lithosphere, or from the edge of

447 magmatic underplate (e.g. Fig. 10C) (Buchan and Ernst, 2018b). The 1385 Ma Lake Victoria

448 giant dyke swarm has an arcuate geometry extending for ~180° around an arc, which has a

449 diameter of ~600 km (Fig. 1) (Mäkitie et al. 2014). It is predicated that this circumferential 450 pattern continues for the full 360 degrees, but this cannot be tested at present because no

451 aeromagnetic data and only limited geological data are available for the western region

452 (Mäkitie et al. 2014). The Lake Victoria swarm is interpreted to mark a plume centre for the

453 Kunene-Kibaran LIP that extends across the Congo craton (Ernst et al. 2013). The magmatic-

454 tectonic conditions responsible for the circumferential pattern are still being investigated

455 (Buchan and Ernst, 2018a, b).

456

457 Surface expression of dykes and dyke swarms

458 Dyke intrusion typically involves tensile failure and opening of the host rock (e.g. Rubin

459 1995; Rubin 1992; Baer 1991; Lister and Kerr 1991), although emplacement may also occur

460 via viscous indentation in some instances (e.g. Spacapan et al. 2017; Abdelmalak et al. 2012;

461 Donnadieu and Merle 1998; Guldstrand et al. 2017). Because the majority of dykes do not

462 breach the surface and instead arrest at depth, the act of generating space for dyke intrusion

463 by host rock extension or viscous indentation can induce deformation of the overlying rock

464 (e.g. Fig. 11) (e.g. Gudmundsson and Brenner 2004; Gudmundsson 2003; Mastin and Pollard

465 1988; Pollard et al. 1983; Donnadieu and Merle 1998; Abdelmalak et al. 2012; Guldstrand et

466 al. 2017; Kavanagh et al. 2018). For example, the development of an extension discrepancy

467 between the level of dyke intrusion and the overlying rock may be resolved by the formation

468 of dyke-induced normal faults that dip towards the upper dyke tip and bound dyke-parallel

469 graben (Fig. 11) (e.g. Mastin and Pollard 1988; Rubin 1992; Pollard et al. 1983; Hjartardóttir

470 et al. 2016; Hofmann 2013; Mège et al. 2003; Wilson and Head 2002). Dyke-induced normal

471 faults have been observed on Earth (e.g. Iceland and Hawaii) and other planetary bodies (e.g.

472 Mars and Venus) and are considered to accommodate the same extension as the underlying

473 dyke, i.e. their cumulative heave is expected to equal the dyke thickness (Fig. 11) (e.g.

474 Mastin and Pollard 1988; Pollard et al. 1983; Rubin 1992; Mège et al. 2003; Ernst et al. 2001; 475 Hjartardóttir et al. 2016; Muirhead et al. 2016). Because the geometry and growth of dyke-

476 induced normal faults is kinematically linked to dyke propagation, it is expected that the

477 surface expression of these faults can be used as a proxy for dyke parameters (e.g. Rowland

478 et al. 2007; Dumont et al. 2017; Dumont et al. 2016; Hofmann 2013; Mastin and Pollard

479 1988). To this end, four hypotheses have been developed to explain dyke-induced normal

480 fault growth and explain their surface expression (Fig. 11): (i) dyke-induced normal faults

481 nucleate at the surface and propagate downwards (Acocella et al. 2003); (ii) they propagate

482 upwards from the dyke tip (Grant and Kattenhorn 2004); (iii) a combination of both

483 (Rowland et al. 2007; Tentler 2005); or (iv) that they nucleate between and grow towards the

484 dyke tip and surface (Mastin and Pollard 1988). It is, however, difficult to rigorously test

485 these hypotheses due to a lack of sufficient 3D field exposures or geophysical images that

486 detail dyke-induced fault structure. An important distinction though between the four

487 hypotheses is that, whilst they all produce similar final fault geometries, the different fault

488 growth patterns predict unique fault displacement trends (Fig. 11). Quantitatively analysing

489 displacement across dyke-induced normal faults could thus allow us to evaluate the kinematic

490 relationship between dykes and dyke-induced normal faults and determine how dyke-induced

491 normal faults can be distinguished from rift-related tectonic normal faults (Fig. 11). Instead

492 of dyke-induced normal faulting, some experimental and numerical models have also

493 demonstrated that space for dyke intrusion may be generated by antiformal folding (i.e.

494 uplift) of the overlying rock volume and/or reverse faulting, particularly if the dyke behaves

495 as a viscous indenter (Abdelmalak et al. 2012; Wyrick and Smart 2009; Guldstrand et al.

496 2017).

497 Satellite imagery suggests that inferred dykes on Mars and Venus appear to

498 occasionally be overlain by pit chain craters (e.g. Wyrick et al. 2004; Ferrill et al. 2004;

499 Patterson et al. 2016). Pit chain craters have been observed on Earth, although they are 500 typically smaller than those on other planetary surfaces (e.g. Ferrill et al. 2011). Overall,

501 these elliptical to circular craters commonly occur in linearly aligned arrays, typically have

502 long axes up to 4.5 km, and are several hundred metres deep (e.g. Wyrick et al. 2004; Ferrill

503 et al. 2004). No eruptive material appears to be associated with pit chain craters. Whilst pit

504 chain craters are generally considered to develop through cavity collapse instigated by dyke

505 intrusion, we cannot validate the various mechanisms put forth to explain their formation,

506 including (Fig. 12): (i) tensile fracturing; (ii) dilational faulting; (iii) void collapse following

507 volatile or magma removal (Wyrick et al. 2004); or (iv) pulsed movement of the dyke front

508 during lateral propagation (Patterson et al. 2016). It is, however, plausible that pit chain

509 craters may also develop through processes not associated with dyke intrusion such as

510 tectonic faulting, karst dissolution or collapse of lava tubes (Fig. 12) (Wyrick et al. 2004).

511 Regardless of dyke emplacement mechanics, associated surface deformation

512 structures (e.g. faults and pit chain craters; Figs 11 and 12) represent an important record of

513 ongoing and historical dyke intrusion events. Characterising the surface expression of dyke-

514 induced deformation thus provides crucial insight into the distribution and geometry of

515 subsurface dykes and dyke swarms (e.g. Ernst et al. 2001; Mastin and Pollard 1988).

516 However, to accurately invert ground deformation patterns and recover the properties and

517 dynamics of underlying dykes, it is critical to understand how dyke intrusion translates to

518 elastic and inelastic surface deformation in four-dimensions (e.g. Kavanagh et al. 2018).

519 Whilst dyke-induced ground deformation can be monitored in near real-time at active

520 volcanoes (e.g. Hjartardóttir et al. 2016), we cannot assess the 3D structure of natural systems

521 from surface data alone.

522

523 Imaging and identifying dyke swarms in seismic reflection data 524 One may infer from a global map of known dyke swarms that many should continue offshore

525 onto the continental passive margins (Fig. 1). Furthermore, dyke swarms may be expected to

526 occur along magma-rich margins where continental break-up was potentially driven by

527 intense dyke intrusion, as is currently observed along magmatic segments of the East African

528 Rift (e.g. Kendall et al. 2005; Karson and Brooks 1999; Abdelmalak et al. 2015; Muirhead et

529 al. 2015). However, with the exception of scant evidence from aeromagnetic and limited

530 seismic reflection surveys for the extension of dyke swarms offshore, few have been

531 observed along continental passive margins (cf. Kirton and Donato 1985; Wall et al. 2010;

532 Abdelmalak et al. 2015). This is primarily because the sub-vertical orientation of dykes

533 makes them difficult to image in seismic reflection data, as the vast majority of acoustic

534 energy encountering a dyke contact is reflected downwards and does not return to the surface

535 (Thomson 2007). Where dykes have been interpreted in seismic reflection data, they have

536 been attributed to either: (i) vertical discontinuities in the data, such that the dykes are not

537 imaged per se but rather correspond to zones from which little acoustic energy is returned

538 (e.g. Wall et al. 2010; Bosworth et al. 2015; Ardakani et al. 2017); or (ii) local inclined

539 reflections representing dykes that have been rotated post-emplacement (e.g. Phillips et al.

540 2017; Abdelmalak et al. 2015). Here, we describe how seismic reflection data can be used to

541 constrain the offshore limits of dyke swarms.

542

543 Skagerrak (Jutland) LIP (~300–280 Ma): Farsund Dyke Swarm, offshore southern Norway

544 Phillips et al. (2017) recently recognised and described a dyke swarm (i.e. the Farsund Dyke

545 Swarm) imaged in 2D and 3D seismic reflection data from offshore southern Norway (e.g.

546 Fig. 13A). The dyke swarm is expressed in seismic data as a series of high-amplitude, tuned

547 reflection packages, which dip northwards at ~35–50°N and cross-cut but do not offset gently

548 (~10–20°) S-dipping, Carboniferous-Permian (<320 Ma) strata (Fig. 13A) (Phillips et al. 549 2017). Post-emplacement flexure of the passive margin during Mesozoic extension rotated

550 the dykes to their current inclination, presumably from an initial sub-vertical orientation,

551 facilitating their imaging in seismic reflection data (Fig. 13B) (Phillips et al. 2017). The tuned

552 reflection packages within the inferred dyke swarm (Fig. 13A) display complex wave-trains,

553 suggesting they represent an amalgamation of reflections emanating from two or more

554 intrusion-host rock contacts that cannot be distinguished seismically (Phillips et al. 2017).

555 These tuned reflection packages, thus, do not represent individual dykes but rather likely

556 result from interference between the margins of dykes with thicknesses and/or spacings <75

557 m (Phillips et al. 2017). No borehole data is available to test this hypothesis.

558 The dyke swarm is mapped across multiple seismic datasets within a ~100 km long,

559 ~25 km wide zone that trends WSW-ENE (Fig. 13C) (Phillips et al. 2017). In the centre of

560 the dyke swarm, reflections extend from a depth of ~2.5 s TWT (i.e. ~4 km), below which

561 seismic imaging deteriorates, up to ~0.8 s TWT (i.e. ~1 km) where they are truncated by the

562 Saalian-Altmark (~290–270 Ma) and Base Zechstein (~260 Ma) unconformities (Figs 13A

563 and B) (Phillips et al. 2017). Towards the margins of the dyke swarm, the spacing of the

564 tuned reflection packages increases and their height progressively decreases as the dykes

565 arrest at lower stratigraphic levels (Figs 13A and B) (Phillips et al. 2017). Because the dyke

566 swarm cross-cuts Carboniferous-Permian (<320 Ma) strata and is truncated by the 290–270

567 Ma Saalian-Altmark unconformity, it is likely that emplacement occurred ~320–270 Ma

568 (Phillips et al. 2017). This inferred emplacement age is broadly coincident with the 300–280

569 Ma Skagerrak LIP, also termed the Skagerrak (Jutland) LIP, which is marked by convergent,

570 linear dyke swarms that place the plume centre at the southern end of the Oslo graben (Figs 1

571 and 13C) (e.g. Phillips et al. 2017; Ernst and Buchan 1997a; Torsvik et al. 2008). The

572 Farsund Dyke Swarm occurs proximal to the proposed source of the Skagerrak LIP and its 573 WSW-trend suggests that it links to the Midland Valley Dyke Suite in the UK (Fig. 13C)

574 (Phillips et al. 2017; Torsvik et al. 2008).

575 Variation in dyke height across the swarm, with dyke-related reflections terminating

576 at lower stratigraphic levels towards its margins (Figs 13A and B), implies that the measured

577 width of dyke swarms exposed at Earth’s surface are dependent on the erosion level (Phillips

578 et al. 2017). For example, the width of the Farsund Dyke Swarm is ~12 km at the Saalian-

579 Altmark unconformity, but at deeper levels it increases to ~20 km (Fig. 13A) (Phillips et al.

580 2017). This change in swarm width with height has implications for calculations of intrusion

581 volume and associated extension. For example, assuming the imaged portion of the dyke

582 swarm has a 1:1 dyke to host rock ratio, with dyke heights of 3 km, the variable width of the

583 swarm with depth means that estimates of its volume and the extension it accommodates

584 range from 900–1500 km3 and 6–10 km, respectively (Phillips et al. 2017). These

585 observations and approximate calculations highlight that the exposed width of a dyke swarm

586 cannot necessarily be used to accurately determine melt volumes and extension across ancient

587 dyke swarms.

588

589 North Atlantic Igneous Province (~61–50 Ma): Mull Dyke Swarm, UK

590 The NW-striking Mull Dyke Swarm extends south-westwards from the Mull Central

591 Complex in NW Scotland, which represents the deeply eroded roots of an ancient volcano,

592 for >300 km onshore (Figs 1 and 14A) (e.g. Speight et al. 1982; Macdonald et al. 2010;

593 Ishizuka et al. 2017). Geochemical and rock magnetic data indicate that dykes intruded

594 laterally away from the central complex, with K-Ar dating highlighting emplacement

595 occurred ~58 Ma; i.e. as part of the North Atlantic Igneous Province (e.g. Macdonald et al.

596 2009; MacDonald et al. 2015). Continuation of the Mull Dyke Swarm offshore, into the

597 southern North Sea, has been postulated based on interpretation of NW-striking, linear 598 features identified in both 3D seismic reflection and aeromagnetic data as dykes (Fig. 14A)

599 (Kirton and Donato 1985; Wall et al. 2010). In the seismic reflection data, these inferred

600 dykes correspond to vertical zones of discontinuous host rock reflections, which are 0.5–2 km

601 wide and up to ~40 km long (Fig. 14B) (Wall et al. 2010). Where the seismic reflection data

602 overlaps with high-resolution aeromagnetic data, the inferred dykes correspond to linear,

603 negative magnetic anomalies (Kirton and Donato 1985; Wall et al. 2010). Seismic reflection

604 data further reveal that overlying and extending along the length of the dykes are a series of

605 circular, ellipsoidal, and linear craters (e.g. Fig. 14B) (Wall et al. 2010). At their top, which

606 occurs at the Top Cretaceous horizon, these craters are up to 2 km wide and infilled by

607 overlying Paleogene strata, implying crater formation occurred to the Early Paleogene (Fig.

608 14B) (Wall et al. 2010). The craters typically have inward-dipping walls, which extend down

609 to the top of the inferred dykes (Fig. 14B) (Wall et al. 2010). These craters are interpreted to

610 be pit chain craters formed in response to explosive volatile release as the propagating dyke

611 approached the contemporaneous seabed (Wall et al. 2010). The age of the pit chain craters,

612 which formed during dyke emplacement, and the trend of the imaged dykes suggests that the

613 Mull Dyke Swarm extends for >600 km (Kirton and Donato 1985; Wall et al. 2010). Where

614 the dykes and pit chain craters are observed in a single 2D seismic section, their appearance

615 is reminiscent of some hydrothermal vents developed above the lateral tips of sills, which can

616 correspond to a roughly cylindrical, vertical zone of dim or chaotic reflections overlain by a

617 crater (Fig. 7B) (e.g. Jamtveit et al. 2004; Hansen 2006); to differentiate between

618 hydrothermal vents and pit chain craters above a dyke, it is necessary to evaluate the along-

619 strike extent of the structure using 3D seismic reflection data.

620

621 Dyke and dyke-induced faults imaged in Egypt 622 Subtle vertical disruptions to host rock reflections within 3D seismic data from the Abu

623 Gharadig Basin, onshore Egypt, have been inferred to correspond to dykes (Fig. 15)

624 (Bosworth et al. 2015). Similar to the pit chain craters observed above the dykes in the North

625 Sea, these Egyptian dykes are also overlain by crater-like features up to several hundred

626 meters in diameter (Fig. 15) (Bosworth et al. 2015). However, the dykes in the Abu Gharadig

627 Basin are also associated with overlying, dyke-parallel normal faults that dip towards and

628 extend upwards from the dyke tip for ~0.3 s TWT (Fig. 15) (Bosworth et al. 2015). These

629 normal faults bound graben and extend for >100 km and have offsets of only a few tens of

630 meters maximum (Bosworth et al. 2015). The association with underlying dykes suggests that

631 these graben-bounding faults nucleated and grew in response to dyke emplacement, i.e. they

632 are dyke-induced normal faults (Bosworth et al. 2015). Such dyke-induced normal faults have

633 not been recognised in other seismic datasets, but these data from Egypt indicate that

634 (Bosworth et al. 2015): (i) dyke-induced normal faults can be more easily recognised in

635 seismic reflection data than dykes (Fig. 15); and (ii) seismic reflection data can allow

636 displacement across dyke-induced normal faults to be quantitatively assessed in 3D. This

637 latter observation is important because mapping displacement patterns across dyke-induced

638 normal faults could allow their kinematic history to be resolved, providing a means of testing

639 existing hypotheses concerning dyke-induced normal fault growth; i.e. where dyke-induced

640 normal faults nucleate, they are likely to be active for longer and therefore accrue more

641 displacement, producing diagnostic displacement-depth trends.

642

643 Field observations of intrusive complexes

644 Intrusive complexes, also referred to as central complexes, represent the deeply eroded roots

645 of major volcanic centres within LIPs (e.g. Emeleus and Bell 2005; Jerram and Bryan 2015;

646 Ewart et al. 2002). These complexes are likely vertically extensive and comprise a dense 647 network of closely spaced magma conduits (e.g. dykes, sills, cone sheets, and ring dykes) and

648 major plutons (e.g. Emeleus and Bell 2005; Jerram and Bryan 2015; Ewart et al. 2002). Such

649 intrusive complexes are perhaps most well-known from onshore the British and Irish

650 Palaeogene Igneous Province, which forms part of the North Atlantic Igneous Province,

651 where they are typically circular in plan-view and up to ~15 km in diameter (Emeleus and

652 Bell 2005). The structure of these intrusive complexes and the geometry of individual

653 components is partly controlled by their composition, with many comprising a combination

654 of ultramafic-mafic layered intrusions and granitoid plutons (e.g. laccoliths) (e.g. Emeleus

655 and Bell 2005; Jerram and Bryan 2015; Bryan et al. 2010; Ewart et al. 2002). Whilst intrusive

656 complexes provide a site for assimilation-fractional crystallisation processes, they can also

657 feed dyke swarms and sill-complexes (e.g. Emeleus and Bell 2005; MacDonald et al. 2015;

658 Preston et al. 1998; Jolly and Sanderson 1995; Bryan et al. 2010). For example, the Mull

659 Intrusive Complex on the Isle of Mull, Scotland fed a laterally propagating dyke swarm (i.e.

660 the Mull dyke swarm), which extends >600 km to the SW (Fig. 14), and the Loch Scridain

661 sill-complex (e.g. Emeleus and Bell 2005; Ishizuka et al. 2017; MacDonald et al. 2015;

662 Preston et al. 1998; Jolly and Sanderson 1995; Speight et al. 1982).

663

664 Discussion

665 Dyke swarms and sill-complexes facilitate the transport of magma through Earth’s crust to

666 the surface during the formation of LIPs, controlling the distribution of eruption sites and

667 associated economic ore deposits. Whilst field observations coupled with petrological,

668 geochemical, geochronological, structural, and geophysical analyses have advanced our

669 understanding of how magma is emplaced and evolves within these systems, the 3D structure

670 of dyke swarms and sill-complexes has largely remained enigmatic. We have demonstrated

671 that, over the last two decades, seismic reflection data has provided unprecedented insights 672 into the 3D structure of LIP plumbing systems. For example, reflection seismology has

673 revealed: (i) sill-complexes commonly transport magma over significant lateral (>10’s km)

674 and vertical (up to 12 km) distances (e.g. Cartwright and Hansen 2006; Schofield et al. 2017;

675 Magee et al. 2016); (ii) sill-complexes can feed extensive networks of hydrothermal vents,

676 which may contribute to and/or drive climate change (e.g. Svensen et al. 2007; Svensen et al.

677 2004); (iii) dykes and dyke swarms can be identified and mapped across continental margins

678 (e.g. Phillips et al. 2017; Abdelmalak et al. 2015; Wall et al. 2010); and (iv) the width of dyke

679 swarms may decrease with height, implying erosion level may control measured widths, and

680 thus volume estimates, of dyke swarms observed at Earth’s surface (Phillips et al. 2017).

681 Despite the advances reflection seismology has offered, here we select and discuss four

682 outstanding questions and/or problems.

683

684 Plumbing system volumes

685 Many sheet intrusions imaged in seismic reflection data correspond to tuned reflection

686 packages; i.e. reflections from the top and base contact of the intrusion interfere on their way

687 back to the surface and cannot be distinguished (e.g. Smallwood and Maresh 2002; Thomson

688 2005; Peron-Pinvidic et al. 2010; Magee et al. 2015; Eide et al. 2017; Rabbel et al. 2018;

689 Phillips et al. 2017). This tuning effect is dependent on the thickness of the body relative to

690 the limit of separability (i.e. vertical resolution) of the seismic reflection data, which itself is a

691 function of data frequency and the seismic velocity of the intrusion (Brown 2004). Intrusions

692 with thicknesses below the limit of separability but above the limit of visibility are imaged as

693 tuned reflection packages. The thickness of intrusions expressed as tuned reflection packages

694 is thus below the limit of separability but above the limit of visibility and cannot be

695 constrained further (Brown 2004). Without being able to accurately assess intrusion

696 thickness, magma volume estimates of LIP plumbing systems imaged in seismic reflection 697 data are compromised. Whilst higher resolution seismic reflection data will help to further

698 constrain intrusive volumes within LIPs, the following would also be useful: (i) a better

699 understanding of how intrusions are expressed in seismic reflection data, which can be

700 achieved through synthetic seismic forward modelling (e.g. Magee et al. 2015; Peron-

701 Pinvidic et al. 2010; Eide et al. 2017; Rabbel et al. 2018); (ii) constraints on the seismic

702 velocity of sills, perhaps through analysis of borehole data or geophysical techniques such as

703 full-waveform inversion (e.g. Morgan et al. 2013); and (iii) a comprehensive analyses of

704 intrusion scaling relationships within exposed and seismically imaged sill-complexes (cf.

705 Cruden et al. 2017 and references therein).

706

707 Lateral magma flow in sill-complexes

708 Seismic reflection data, coupled with field observations, clearly show connections between

709 individual intrusions within sill-complexes can define tortuous magma flow pathways

710 (Magee et al. 2016). Although it is difficult to accurately constrain the duration of sill-

711 complex formation, seismic-based studies suggest that emplacement may occur incrementally

712 over millions of years (e.g. Reeves et al. 2018; Magee et al. 2014). However, most intrusions

713 within sill-complexes are emplaced at shallow-levels (>3 km depths) within relatively cold

714 sedimentary strata and are expressed as tuned reflection packages, implying their thickness is

715 below the limit of separability for any given seismic volume (i.e. typically <100 m) (see

716 Magee et al. 2016 and references therein). It may therefore be expected that individual

717 intrusions will solidify rapidly, raising the question as to how magma transits through long-

718 lived sill-complexes. Several studies have demonstrated that sheet accretion or magma

719 channelization within discrete intrusions can facilitate sill-complex emplacement (e.g. Annen

720 et al. 2015; Holness and Humphreys 2003; Magee et al. 2016), but whether these processes

721 can maintain elevated host rock temperatures and thereby inhibit solidification remains 722 unknown. To understand how magma moves through these intrusion networks it is critical

723 integrate rock magnetic, petrological, geochemical, and high-resolution geochronological

724 data to track flow pathways, residence times, and magma evolution across exposed sill-

725 complexes.

726

727 Sill-complex and dyke swarm connectivity

728 Understanding the 4D connectivity and evolution of intrusive components in LIP plumbing

729 systems is integral to evaluating controls on the accumulation and location of volumetrically

730 large magma bodies, the eruption of which (e.g. to form flood basalts) can trigger

731 catastrophic climatic and environmental changes. Magma transport within LIPs is dominated

732 by sill-complexes and/or dyke swarms, yet it remains unclear how these two intrusion

733 networks relate to each other. For example, it is traditionally assumed that dykes bring

734 magma from mantle source areas vertically into the lower crust, and upward from the

735 magmatic underplate into the upper brittle crust (see Ernst 2014 and references therein).

736 Given that sill-complexes are predominantly observed within sedimentary basins and appear

737 to have been emplaced at relatively shallow-levels (>5 km depth), it may be expected that

738 they are fed by dyke swarms. This potential relationship LIP plumbing system components is

739 supported by the apparent feeding of the 2.21 Ga Nipissing sill-complex by the

740 contemporaneous Senneterre dyke swarm (e.g. Ernst and Buchan 1997a; Buchan et al. 1993).

741 However, the mechanics and implications of the transition from dyke swarm to sill-complex

742 have not been assessed. Furthermore, field- and seismic-based observations demonstrate sill-

743 complexes can transport magma up to 12 km vertically, from lower crustal areas, and develop

744 within crystalline rocks (Cartwright and Hansen 2006; Ivanic et al. 2013a; McBride et al.

745 2018). Whilst these observations of basement-hosted sill-complexes questions the volumetric

746 importance of dykes at depth, it also raises the possibility that sill-complexes could feed dyke 747 swarms. In particular, the repeated and rapid input of magma into a sill-complex could

748 promote the development of a large intrusive centre (e.g. Cawthorn 2012; Ernst and Buchan

749 1997a; Annen et al. 2015; Annen 2011), potentially capable of feeding dyke swarms (e.g. the

750 Mull Intrusive Complex; Emeleus and Bell 2005; Ishizuka et al. 2017; MacDonald et al.

751 2015; Preston et al. 1998; Jolly and Sanderson 1995; Speight et al. 1982).

752 Our poor of understanding of sill-complex and dyke swarm connectivity is partly due

753 to: (i) a lack of exposed dyke and sill feeding relationships within LIPs; and (ii) difficulties in

754 imaging dykes in seismic reflection. Resolving possible feeder dyke relationships within sill-

755 complexes using seismic reflection data is particularly challenging because their imaging

756 results in a significant reduction in data quality directly beneath the complex as acoustic

757 energy is reflected back or scattered by the sills (e.g. Planke et al. 2005; Eide et al. 2017).

758 From a reflection seismology perspective, one way to potentially circumvent this problem is

759 to map magma flow indicators across sill-complexes to determine the location and geometry

760 of magma input zones (e.g. Schofield et al. 2017; Magee et al. 2016; Schofield et al. 2012a;

761 Thomson and Hutton 2004), which may provide insight into whether sills are fed by dykes

762 (e.g. Goulty and Schofield 2008). Mapping flow patterns and identifying feeder locations

763 within seismically imaged flood basalts spatially associated with sill-complexes may also

764 provide a method for distinguishing the role of dykes and dyke swarms within LIPs.

765

766 Conclusions

767 Large Igneous Provinces (LIPs) have become an important focus for research in recent years

768 due to their use in paleocontinental reconstructions, in exploration targeting, as planetary

769 analogues, and as a result of their links to dramatic climate change. A key frontier within the

770 study of LIPs concerns deciphering the 3D geometry of their plumbing systems and, in

771 particular, identifying how dyke swarms and sill-complexes grow and interact. Seismic 772 reflection data uniquely reveal the 3D geometry of LIP magma plumbing systems,

773 highlighting that dyke swarms and sill-complexes dictate melt distribution. Here, we review

774 several sill-complexes and dyke swarms either exposed at Earth’s surface or analysed using

775 seismic reflection data. In particular, we highlight that: (i) mapping sill-complexes in 3D and

776 recognising magma flow indicators can be used to unravel their source regions, sill

777 connectivity, and overall growth history; (ii) sill-complexes can facilitate significant vertical

778 and lateral magma transport both within sedimentary basins and crystalline crust, potentially

779 instigating widespread hydrothermal venting; and (iii) dyke swarm width varies with dyke

780 height, implying that the observed width of dyke swarms exposed at Earth’s surface is a

781 function of erosion level and may not represent the true swarm size. Furthermore, we also

782 show that seismic reflection data allows the visualisation and quantification of host rock

783 deformation, including deformation of the contemporaneous surface, induced by dyke swarm

784 and sill-complex emplacement. By understanding how dyke swarms and sill-complexes are

785 expressed at the surface we can estimate properties of LIP magma plumbing systems that are

786 unexposed on Earth (e.g. East African Rift System) or other planetary bodies (e.g. Mars).

787 Overall, building our understanding of LIP magma plumbing system geometries,

788 emplacement mechanics, and relative timings through analysis of seismic reflection data is

789 critical to evaluating associated volcanic hazards, either in active systems or where

790 magmatism (both extrusive and intrusive) may have driven mass extinctions, and

791 accumulation of economic resources.

792

793 Acknowledgements

794 CM is funded by a Junior Research Fellowship at Imperial College London. REE was

795 partially supported from Mega-Grant 14.Y26.31.0012 of the government of the Russian

796 Federation. JDM is supported by National Science Foundation grant EAR-1654518. We 797 would like to thank Rajesh Srivastava for editorial handling and an anonymous reviewer and

798 Dougal Jerram for their comments.

799

800 Figure Captions

801 Figure 1: Map showing selected plume centres and associated dyke swarms and sill-

802 complexes, with each representing part of the plumbing system of a Large Igneous Province

803 (LIP) (modified from Ernst, 2014). Dyke swarms: A = 1140 Ma Abitibi swarm, precursor to

804 the 1115–1085 Ma Keweenawan LIP; AA = 30–0 Ma Afar-Arabian swarms; C = 17–0 Ma

805 Columbia River swarms; CAMP = 201 Ma Central Atlantic Magmatic Province swarm; D =

806 66 Ma Deccan swarm; F = 725–716 Ma Franklin swarm; G = 779 Ma Gunbarrel swarms; Ga

807 = 799 Ma Gannakouriep swarm; H = 130–90 Ma High Arctic LIP (HALIP) swarm; J = 301

808 Ma Skagerrak (Jutland) swarms; K = 183 Ma Karoo swarms; M = 2510 Ma Mistassini

809 swarm; Mac = 1267 Ma Mackenzie swarm; Md = 89 Ma Madagascar swarm; Mt = 2480–

810 2450 Ma Matachewan swarm; NAIP = 62–55 Ma North Atlantic Igneous Province swarms

811 (e.g. Mu is the Mull dyke swarm); P-E = ~135–128 Ma Paraná-Etendeka dyke swarms; S =

812 251 Ma Siberian Traps swarm; U = 2217-2210 Ma Ungava swarm; Y = 370 Ma Yakutsk-

813 Vilyui swarm. Also show is the 1385 Ma Victoria dyke swarm (V) (Mäkitie et al. 2014).

814 Selected sill-complexes: E = ~150 Ma Exmouth Sub-basin sills; F = ~55 Ma Faroe-Shetland

815 Basin sill-complex; IR = ~65–50 Ma Irish Rockall Basin sill-complex; Ka = 183 Ma Karoo

816 sill-complex; Si = 251 Ma Siberian Traps sill-complex; VM = Vøring-Møre basins sill-

817 complex; W = 1070 Ma Warakurna sill-complex. The ~183 Ma Ferrar sill-complex in

818 Antarctica is shown in Figure 2A.

819

820 Figure 2: (A) Paleogeographic reconstruction showing the distribution of intrusive elements

821 within the Karoo and Ferrar LIPs (modified from Leat, 2008). (B) Interpreted photograph of 822 interconnected sills and inclined sheets, part of the Ferrar sill-complex, intruded into the

823 Beacon Supergroup sedimentary rocks (image courtesy of James White). (C) Interpreted

824 Google Earth image of the Golden Valley Sill and surrounding intrusions, located in the

825 Karoo Basin, South Africa. GVS = Golden Valley Sill; MVS = MV Sill; GS = Glen Sill; HS

826 = Harmony Sill; MSS = Morning Sun Sill; GVD = Golden Valley Dyke.

827

828 Figure 3: (A and B) Plan-view and 3D view (vertical exaggeration = 2) of the Alu dome and

829 surrounding lava flows. (C) Sill contraction area and magnitude inverted from InSAR data

830 collected during a basalt eruption in 2008 (Pagli et al. 2012). Note that the area of sill

831 contraction underlies the Alu dome. (D) Schematic of interpreted plumbing system beneath

832 Alu. Images modified from Magee et al. (2017).

833

834 Figure 4: (A-C) Examples of strata-concordant (strat-con.) and saucer-shaped sills, as well as

835 inclined sheets, observed in seismic reflection data (i.e. in cross-section and 3D views) and

836 field outcrop. Field photographs courtesy of Donny Hutton (A), Google Earth (B), and James

837 White (C). Figure modified from Magee et al. (2016). (D) Intrusive step and broken bridge

838 traces (red lines), which allow magma flow patterns to be inferred (dashed white arrows)

839 observed on a plan-view variance map of a sill located in the Irish Rockall Basin (see Magee

840 et al. 2014). Cross-sections through the sill, depicting the different magma flow indicators,

841 are also shown.

842

843 Figure 5: (A) Seismic section from the Irish Rockall Basin showing an interconnected

844 network of high-amplitude reflections, interpreted to be sills and inclined sheets, overlain by

845 domed strata (Magee et al. 2014). Seismic-stratigraphic onlap, downlap, and truncation

846 patterns occur throughout the folded strata. (B) Map of the 82 seismically resolved sills 847 interpreted by Magee et al. (2014), with colours depicting how the sills are vertically stacked.

848 (C) Schematic of the evolution of sill emplacement and forced fold growth, over a protracted

849 period of time, in the Irish Rockall Basin (Magee et al. 2014).

850

851 Figure 6: (A) Map of seismically resolved sills interpreted across the Faroe-Shetland Basin,

852 showing magma flow patterns and interpreted zones of magma input into the basin (Schofield

853 et al. 2017). (B) Uninterpreted and interpreted seismic section through the Faroe-Shetland

854 Basin, highlighting the distributions, structure, and connectivity of sills within the Cretaceous

855 strata (Schofield et al. 2017). See Figure 10A for location.

856

857 Figure 7: (A) Map showing the areal coverage of the sill-complex and associated

858 hydrothermal vent complexes (filled circles) spanning the Vøring and Møre basins, offshore

859 Norway (modified from Svensen et al. 2004). (B) Schemtic showing the different types of

860 hydrothermal vent complex, which are classified based on the morphology of their upper

861 section; i.e. dome-, crater-, or eye-shaped.

862

863 Figure 8: Two seismic lines from the Yilgarn Craton, Western Australia, showing a network

864 of sills intruded into crystalline basement rocks (modified from Ivanic et al. 2013).

865

866 Figure 9: (A and B) Uninterpreted and interpreted total magnetic intensity anomaly map

867 acquired during the Tellus airborne geophysical survey of Northern Ireland (modified from

868 Cooper et al. 2012). Linear magnetic anomalies are attributed to different dyke swarms, with

869 larger anomalies corresponding to the Antrim Plateau (AP) basalt lavas and the Slieve

870 Guillian (SG) and Mourne Mountains (MM) intrusive complexes. Offsets in the dyke-related

871 linear anomalies mark the position of strike-slip faults: CF = Carnlough Fault; CLF = 872 Camlough Fault; CVF = Clogher Valley Fault; LF = Loughguile Fault; NF = Newry Fault;

873 OF = Omagh Fault; SF = Sixmilewater Fault; TSF = Tempo-Sixmilecross Fault; TVF =Tow

874 Valley Fault.

875

876 Figure 10: Schematics showing differences in the geometry of linear, radial, and

877 circumferential dyke swarms. See text for discussion.

878

879 Figure 11: Sketches depicting how different inferred models of dyke-induced fault growth

880 produce graben observed above dykes on Earth and other planetary bodies: (i) faults nucleate

881 at the surface and propagate downwards (Acocella et al. 2003); (ii) faults nucleate at the dyke

882 tip and propagate upwards (Grant and Kattenhorn 2004); (iii) faults nucleate at the surface

883 and dyke tip, propagating towards each other and becoming dip-linked (Tentler 2005;

884 Rowland et al. 2007); and (iv) dykes nucleate between the dyke tip and surface, propagating

885 towards each (Mastin and Pollard 1988). Also shown are the predicted displacement-depth

886 plots for the different fault models.

887

888 Figure 12: Photograph of a linear array of pit chain craters in the Natron Basin, Tanzania that

889 occur above a dilated normal fault and formed in relation to a dyke intrusion event in 2007

890 (see Muirhead et al. 2015 for further detail). Also shown are schematics of different models

891 for pit chain crater formation (modified from Wyrick et al. 2004).

892

893 Figure 13: (A) Seismic section from the Farsund Basin, offshore southern Norway that

894 images part of a dyke-swarm (modified from Phillips et al. 2017). Note that the dyke swarm

895 cross-cuts Carboniferous-Permian strata, but is truncated by the Saalian-Altmark and Base

896 Zechstein unconformities (Phillips et al. 2017). (B) Schematic detailing how basin flexure 897 rotated the originally sub-vertical dykes, allowing them to be imaged in seismic reflection

898 data (modified from Phillips et al. 2017). (C) Map showing the modern day distribution of

899 dyke swarms and volcanic rocks related to the Skagerrak (Jutland) LIP (Phillips et al. 2017).

900

901 Figure 14: (A) Map of the Mull dyke swarm, which emanates from the Mull Central Complex

902 (based on Wall et al. 2010; Macdonald et al. 2015; Ishizuka et al. 2017). (B and C) Seismic

903 sections from Wall et al. (2011) showing the dykes, which correspond to vertical zones of

904 low-amplitude, disturbed reflections, and the pit chain craters. See Figure 14A for location.

905

906 Figure 15: Seismic section and map showing a dyke, dyke-induced faults (and graben), and

907 pit chain craters imaged in data from Egypt (Bosworth et al. 2015). Figure 15B shows the

908 location of the seismic section in (A).

909

910 References 911 912 Abdelmalak MM, Andersen TB, Planke S, Faleide JI, Corfu F, Tegner C, Shephard GE, 913 Zastrozhnov D, Myklebust R (2015) The ocean-continent transition in the mid- 914 Norwegian margin: Insight from seismic data and an onshore Caledonian field 915 analogue. Geology 43 (11):1011-1014 916 Abdelmalak MM, Mourgues R, Galland O, Bureau D (2012) Fracture mode analysis and 917 related surface deformation during dyke intrusion: Results from 2D experimental 918 modelling. Earth and Planetary Science Letters 359:93-105 919 Acocella V, Korme T, Salvini F (2003) Formation of normal faults along the axial zone of the 920 Ethiopian Rift. Journal of Structural Geology 25 (4):503-513 921 Agirrezabala LM (2015) Syndepositional forced folding and related fluid plumbing above a 922 magmatic laccolith: Insights from outcrop (Lower Cretaceous, Basque-Cantabrian 923 Basin, western Pyrenees). Geological Society of America Bulletin:B31192. 31191 924 Airoldi G, Muirhead JD, White JD, Rowland J (2011) Emplacement of magma at shallow 925 depth: insights from field relationships at Allan Hills, south Victoria Land, East 926 Antarctica. Antarctic Science-Institutional Subscription 23 (3):281 927 Airoldi G, Muirhead JD, Zanella E, White JD (2012) Emplacement process of Ferrar Dolerite 928 sheets at Allan Hills (South Victoria Land, Antarctica) inferred from magnetic fabric. 929 Geophysical Journal International 188 (3):1046-1060 930 Airoldi GM, Muirhead JD, Long SM, Zanella E, White JD (2016) Flow dynamics in mid- 931 Jurassic dikes and sills of the Ferrar large igneous province and implications for long- 932 distance magma transport. Tectonophysics 683:182-199 933 Annen C (2011) Implications of incremental emplacement of magma bodies for magma 934 differentiation, thermal aureole dimensions and plutonism–volcanism relationships. 935 Tectonophysics 500 (1):3-10 936 Annen C, Blundy JD, Leuthold J, Sparks RSJ (2015) Construction and evolution of igneous 937 bodies: Towards an integrated perspective of crustal magmatism. Lithos 230:206-221. 938 Archer SG, Bergman SC, Iliffe J, Murphy CM, Thornton M (2005) Palaeogene igneous rocks 939 reveal new insights into the geodynamic evolution and petroleum potential of the 940 Rockall Trough, NE Atlantic Margin. Basin Research 17 (1):171-201. 941 Ardakani EP, Schmitt DR, Currie CA (2017) Geophysical evidence for an igneous dike 942 swarm, Buffalo Creek, Northeast Alberta. Geological Society of America Bulletin 943 Aspler LB, Cousens BL, Chiarenzelli JR (2002) Griffin gabbro sills (2.11 Ga), Hurwitz 944 Basin, Nunavut, Canada: long-distance lateral transport of magmas in western 945 Churchill Province crust. Precambrian Research 117 (3–4):269-294. 946 Baer G (1991) Mechanisms of dike propagation in layered rocks and in massive, porous 947 sedimentary rocks. Journal of Geophysical Research: Solid Earth 96 (B7):11911- 948 11929 949 Baragar W, Ernst R, Hulbert L, Peterson T (1996) Longitudinal petrochemical variation in 950 the Mackenzie dyke swarm, northwestern Canadian Shield. Journal of Petrology 37 951 (2):317-359 952 Bleeker W, Ernst R (2006) Short-lived mantle generated magmatic events and their dyke 953 swarms: the key unlocking Earth’s paleogeographic record back to 2.6 Ga. Dyke 954 swarms—time markers of crustal evolution:3-26 955 Bosworth W, Stockli DF, Helgeson DE (2015) Integrated outcrop, 3D seismic, and 956 geochronologic interpretation of Red Sea dike-related deformation in the Western 957 Desert, Egypt–The role of the 23Ma Cairo “mini-plume”. Journal of African Earth 958 Sciences 109:107-119 959 Brown AR (2004) Interpretation of three-dimensional seismic data, vol 42. AAPG Memoir 960 42, SEG Investigations in Geophysics No. 9, 6th edn. AAPG and SEG, Oklahoma, 961 USA 962 Bryan SE, Ernst RE (2008) Revised definition of large igneous provinces (LIPs). Earth- 963 Science Reviews 86 (1-4):175-202 964 Bryan SE, Ferrari L (2013) Large igneous provinces and silicic large igneous provinces: 965 Progress in our understanding over the last 25 years. GSA Bulletin 125 (7-8):1053- 966 1078 967 Bryan SE, Peate IU, Peate DW, Self S, Jerram DA, Mawby MR, Marsh JG, Miller JA (2010) 968 The largest volcanic eruptions on Earth. Earth-Science Reviews 102 (3-4):207-229 969 Buchan K, Ernst R (2013) Diabase dyke swarms of Nunavut, Northwest Territories and 970 Yukon, Canada. Geological Survey of Canada, Open File 7464 (10.4095):293149 971 Buchan K, Mortensen J, Card K (1993) Northeast-trending Early Proterozoic dykes of 972 southern Superior Province: multiple episodes of emplacement recognized from 973 integrated paleomagnetism and U–Pb geochronology. Canadian Journal of Earth 974 Sciences 30 (6):1286-1296 975 Buchan KL, Ernst R (2004) Diabase Dyke Swarms and Related Unites in Canada and 976 Adjacent Regions. Geological Survey of Canada, 977 Bunger AP, Menand T, Cruden A, Zhang X, Halls H (2013) Analytical predictions for a 978 natural spacing within dyke swarms. Earth and Planetary Science Letters 375:270-279 979 Burgess S, Bowring S, Fleming T, Elliot D (2015) High-precision geochronology links the 980 Ferrar large igneous province with early-Jurassic ocean anoxia and biotic crisis. Earth 981 and Planetary Science Letters 415:90-99 982 Burgess S, Muirhead J, Bowring S (2017) Initial pulse of Siberian Traps sills as the trigger of 983 the end-Permian mass extinction. Nature communications 8 (1):164 984 Burke K, Dewey J (1973) Plume-generated triple junctions: key indicators in applying plate 985 tectonics to old rocks. The Journal of Geology 81 (4):406-433 986 Cartwright J, Hansen DM (2006) Magma transport through the crust via interconnected sill 987 complexes. Geology 34 (11):929-932. 988 Cawthorn R (2012) Multiple sills or a layered intrusion? Time to decide. South African 989 Journal of Geology 115 (3):283-290 990 Chevallier L, Woodford A (1999) Morpho-tectonics and mechanism of emplacement of the 991 dolerite rings and sills of the western Karoo, South Africa. South African Journal of 992 Geology 102 (1):43-54 993 Coffin MF, Eldholm O (1992) Volcanism and continental break-up: a global compilation of 994 large igneous provinces. In: Storey BC, Alabaster T, Pankhurst RJ (eds) Magmatism 995 and the Causes of Continental Break-up, vol 68. vol 1. Geological Society of Special 996 Publications, pp 17-30 997 Coffin MF, Eldholm O (1994) Large igneous provinces: crustal structure, dimensions, and 998 external consequences. Reviews of Geophysics 32 (1):1-36 999 Coffin MF, Eldholm O (2005) Large igneous provinces. Encyclopedia of Geology:315-323 1000 Cooper M, Anderson H, Walsh J, Van Dam C, Young M, Earls G, Walker A (2012) 1001 Palaeogene Alpine tectonics and Icelandic plume-related magmatism and deformation 1002 in Northern Ireland. Journal of the Geological Society 169 (1):29-36 1003 Cruden A, McCaffrey K, Bunger A (2017) Geometric scaling of tabular igneous intrusions: 1004 Implications for emplacement and growth. 1005 Donnadieu F, Merle O (1998) Experiments on the indentation process during cryptodome 1006 intrusions: new insights into Mount St. Helens deformation. Geology 26 (1):79-82 1007 Dragoni M, Lanza R, Tallarico A (1997) Magnetic anisotropy produced by magma flow: 1008 theoretical model and experimental data from Ferrar dolerite sills (Antarctica). 1009 Geophysical Journal International 128 (1):230-240. 1010 du Toit AL (1920) The Karroo dolerites of South Africa: a study in hypabyssal injection. 1011 South African Journal of Geology 23 (01):1-42 1012 Dumont S, Klinger Y, Socquet A, Doubre C, Jacques E (2017) Magma influence on 1013 propagation of normal faults: Evidence from cumulative slip profiles along Dabbahu- 1014 Manda-Hararo rift segment (Afar, Ethiopia). Journal of Structural Geology 95:48-59 1015 Dumont S, Socquet A, Grandin R, Doubre C, Klinger Y (2016) Surface displacements on 1016 faults triggered by slow magma transfers between dyke injections in the 2005–2010 1017 rifting episode at Dabbahu–Manda–Hararo rift (Afar, Ethiopia). Geophysical Journal 1018 International 204 (1):399-417 1019 Ebinger CJ, Sleep N (1998) Cenozoic magmatism throughout east Africa resulting from 1020 impact of a single plume. Nature 395 (6704):788 1021 Eide CH, Schofield N, Lecomte I, Buckley SJ, Howell JA (2017) Seismic Interpretation of 1022 Sill-complexes in Sedimentary Basins: The ‘sub-sill Imaging Problem’. Journal of the 1023 Geological Society, 175(2), pp.193-209. 1024 Elliot DH (1992) Jurassic magmatism and tectonism associated with Gondwanaland break- 1025 up: an Antarctic perspective. Geological Society, London, Special Publications 68 1026 (1):165-184. 1027 Elliot DH, Fleming TH (2000) Weddell triple junction: the principal focus of Ferrar and 1028 Karoo magmatism during initial breakup of Gondwana. Geology 28 (6):539-542. 1029 Elliot DH, Fleming TH (2004) Occurrence and dispersal of magmas in the Jurassic Ferrar 1030 large igneous province, Antarctica. Gondwana Research 7 (1):223-237. 1031 Elliot DH, Fleming TH (2017) The Ferrar Large Igneous Province: field and geochemical 1032 constraints on supra-crustal (high-level) emplacement of the magmatic system. 1033 Geological Society, London, Special Publications 463:SP463. 1034 Elliot DH, Fleming TH, Kyle PR, Foland KA (1999) Long-distance transport of magmas in 1035 the Jurassic Ferrar large igneous province, Antarctica. Earth and Planetary Science 1036 Letters 167 (1):89-104. 1037 Emeleus CH, Bell B (2005) The Palaeogene volcanic districts of Scotland, vol 3. British 1038 Geological Survey. 1039 Encarnación J, Fleming TH, Elliot DH, Eales HV (1996) Synchronous emplacement of Ferrar 1040 and Karoo dolerites and the early breakup of Gondwana. Geology 24 (6):535-538 1041 Ernst RE, Baragar WRA (1992) Evidence from magnetic fabric for the flow pattern of 1042 magma in the Mackenzie giant radiating dyke swarm. Nature 356 (6369):511-513 1043 Ernst RE, Head J, Parfitt E, Grosfils E, Wilson L (1995) Giant radiating dyke swarms on 1044 Earth and Venus. Earth-Science Reviews 39 (1):1-58 1045 Ernst RE, Buchan KL (1997a) Giant radiating dyke swarms: their use in identifying pre‐ 1046 Mesozoic large igneous provinces and mantle plumes. Large igneous provinces: 1047 continental, oceanic, and planetary flood volcanism:297-333 1048 Ernst RE, Buchan K (1997b) Layered mafic intrusions: a model for their feeder systems and 1049 relationship with giant dyke swarms and mantle plume centres. South African Journal 1050 of Geology 100 (4):319-334 1051 Ernst RE, Grosfils E, Mege D (2001) Giant dike swarms: Earth, venus, and mars. Annual 1052 Review of Earth and Planetary Sciences 29 (1):489-534 1053 Ernst RE, Bleeker W, Söderlund U, Kerr AC (2013) Large Igneous Provinces and 1054 supercontinents: Toward completing the plate tectonic revolution. Lithos 174:1-14 1055 Ernst RE, Buchan KL, Campbell IH (2005) Frontiers in large igneous province research. 1056 Lithos 79 (3):271-297 1057 Ernst RE, Bleeker W (2010) Large igneous provinces (LIPs), giant dyke swarms, and mantle 1058 plumes: significance for breakup events within Canada and adjacent regions from 2.5 1059 Ga to the Present This article is one of a selection of papers published in this Special 1060 Issue on the the theme Lithoprobe—parameters, processes, and the evolution of a 1061 continent. Lithoprobe Contribution 1482. Geological Survey of Canada Contribution 1062 20100072. Canadian Journal of Earth Sciences 47 (5):695-739 1063 Ernst RE, Jowitt SM (2013) Large igneous provinces (LIPs) and metallogeny. Tectonics, 1064 Metallogeny, and Discovery: the North American Cordillera and Similar Accretionary 1065 Settings Edited by M Colpron, T Bissig, BG Rusk, and JFH Thompson Society of 1066 Economic Geologists Special Publication 17:17-51 1067 Ernst RE (2014) Large igneous provinces. Cambridge University Press. 1068 Ernst RE, Youbi N (2017) How Large Igneous Provinces affect global climate, sometimes 1069 cause mass extinctions, and represent natural markers in the geological record. 1070 Palaeogeography, Palaeoclimatology, Palaeoecology 478:30-52 1071 Ewart A, Milner SC, Duncan AR, Bailey M (2002) The Cretaceous Messum igneous 1072 complex, S.W. Etendeka, Namibia: reinterpretation in terms of a downsag-cauldron 1073 subsidence model. Journal of Volcanology and Geothermal Research 114 (3–4):251- 1074 273. 1075 Ferrill DA, Wyrick DY, Morris AP, Sims DW, Franklin NM (2004) Dilational fault slip and 1076 pit chain formation on Mars. GSA Today 14 (10):4. 1077 Ferrill DA, Wyrick DY, Smart KJ (2011) Coseismic, dilational-fault and extension-fracture 1078 related pit chain formation in Iceland: Analog for pit chains on Mars. Lithosphere 3 1079 (2):133-142 1080 Galerne CY, Galland O, Neumann E-R, Planke S (2011) 3D relationships between sills and 1081 their feeders: evidence from the Golden Valley Sill Complex (Karoo Basin) and 1082 experimental modelling. Journal of Volcanology and Geothermal Research 202 (3- 1083 4):189-199. 1084 Galerne CY, Neumann E-R, Planke S (2008) Emplacement mechanisms of sill complexes: 1085 Information from the geochemical architecture of the Golden Valley Sill Complex, 1086 South Africa. Journal of Volcanology and Geothermal Research 177 (2):425-440. 1087 Galland O (2012) Experimental modelling of ground deformation associated with shallow 1088 magma intrusions. Earth and Planetary Science Letters 317:145-156 1089 Glazner AF, Bartley JM, Carl BS (1999) Oblique opening and noncoaxial emplacement of 1090 the Jurassic Independence dike swarm, California. Journal of Structural Geology 21 1091 (10):1275-1283 1092 Goulty NR, Schofield N (2008) Implications of simple flexure theory for the formation of 1093 saucer-shaped sills. Journal of Structural Geology 30 (7):812-817. 1094 Grant JV, Kattenhorn SA (2004) Evolution of vertical faults at an extensional plate boundary, 1095 southwest Iceland. Journal of Structural Geology 26 (3):537-557 1096 Grapes R, Reid D, McPherson J (1974) Shallow dolerite intrusion and phreatic eruption in the 1097 Allan Hills region, Antarctica. New Zealand Journal of Geology and Geophysics 17 1098 (3):563-577 1099 Gudmundsson A (2003) Surface stresses associated with arrested dykes in rift zones. Bull 1100 Volcanol 65 (8):606-619. 1101 Gudmundsson A, Brenner L (2004) Local stresses, dyke arrest and surface deformation in 1102 volcanic edificesand rift zones. Annals of Geophysics 47 (4) 1103 Guldstrand F, Burchardt S, Hallot E, Galland O (2017) Dynamics of surface deformation 1104 induced by dikes and cone sheets in a cohesive Coulomb brittle crust. Journal of 1105 Geophysical Research: Solid Earth 122 (10):8511-8524 1106 Gunn B, Warren G (1962) Geology of Victoria Land between the Mawson and Mulock 1107 Glaciers: New Zealand Geological Survey. Bulletin 71:133-135 1108 Halls H (1982) The importance and potential of mafic dyke swarms in studies of geodynamic 1109 processes. Geoscience Canada 9 (3):145-154 1110 Hansen DM (2006) The morphology of intrusion-related vent structures and their 1111 implications for constraining the timing of intrusive events along the NE Atlantic 1112 margin. Journal of the Geological Society 163:789-800 1113 Hansen DM, Cartwright J (2006a) Saucer-shaped sill with lobate morphology revealed by 3D 1114 seismic data: implications for resolving a shallow-level sill emplacement mechanism. 1115 Journal of the Geological Society 163 (3):509-523 1116 Hansen DM, Cartwright J (2006b) The three-dimensional geometry and growth of forced 1117 folds above saucer-shaped igneous sills. Journal of Structural Geology 28 (8):1520- 1118 1535. 1119 Hansen DM, Cartwright JA, Thomas D (2004) 3D seismic analysis of the geometry of 1120 igneous sills and sill junction relationships. In: Davies RJ, Cartwright J, Stewart SA, 1121 Lappin M, Underhill JR (eds) 3D seismic technology: Application to the Exploration 1122 of Sedimentary Basins, vol 29. vol 1. Geological Society, London, Special 1123 Publications, pp 199-208 1124 Hastie WW, Watkeys MK, Aubourg C (2014) Magma flow in dyke swarms of the Karoo 1125 LIP: Implications for the mantle plume hypothesis. Gondwana Research 25 (2):736- 1126 755 1127 Head JW, Coffin MF (1997) Large igneous provinces: a planetary perspective. Large Igneous 1128 Provinces: Continental, Oceanic, and Planetary Flood Volcanism:411-438 1129 Hjartardóttir ÁR, Einarsson P, Gudmundsson MT, Högnadóttir T (2016) Fracture movements 1130 and graben subsidence during the 2014 Bárðarbunga dike intrusion in Iceland. Journal 1131 of Volcanology and Geothermal Research 310:242-252 1132 Hofmann B (2013) How do faults grow in magmatic rifts? LiDAR and InSAR observations 1133 of the Dabbahu rift segment, Afar, Ethiopia. University of Leeds, 1134 Hofmann C, Courtillot V, Feraud G, Rochette P, Yirgu G, Ketefo E, Pik R (1997) Timing of 1135 the Ethiopian flood basalt event and implications for plume birth and global change. 1136 Nature 389 (6653):838 1137 Holness M, Humphreys M (2003) The Traigh Bhàn na Sgùrra sill, Isle of Mull: Flow 1138 localization in a major magma conduit. Journal of Petrology 44 (11):1961-1976 1139 Hutton DHW (2009) Insights into magmatism in volcanic margins: bridge structures and a 1140 new mechanism of basic sill emplacement - Theron Mountains, Antarctica. Petroleum 1141 Geoscience 15 (3):269-278. 1142 Ishizuka O, Taylor RN, Geshi N, Mochizuki N (2017) Large‐volume lateral magma transport 1143 from the Mull volcano: An insight to magma chamber processes. Geochemistry, 1144 Geophysics, Geosystems 18 (4):1618-1640 1145 Ivanic T, Korsch R, Wyche S, Jones L, Zibra I, Blewett R, Jones T, Milligan P, Costelloe R, 1146 Van Kranendonk M Preliminary interpretation of the 2010 Youanmi deep seismic 1147 reflection lines and magnetotelluric data for the Windimurra Igneous Complex. In: 1148 Youanmi and Southern Carnarvon seismic and magnetotelluric (MT) workshop, 1149 Geological Survey of Wetern Australia, Record, 2013a. 1150 Ivanic T, Zibra I, Doublier M, Wyche S (2013e) Geological Interpretation of the Youanmi 1151 and Southern Carnarvon seismic lines 10GA-YU1, 10GA-YU2, 10GA-YU3, and 1152 11GA-SC1. Youanmi and Southern Carnarvon seismic and magnetotelluric (MT) 1153 workshop 2013. Geological Survey of Western Australia 1154 Jackson CA-L, Schofield N, Golenkov B (2013) Geometry and controls on the development 1155 of igneous sill–related forced folds: A 2-D seismic reflection case study from offshore 1156 southern Australia. Geological Society of America Bulletin 125 (11-12):1874-1890. 1157 Jackson MD, Pollard DD (1988) The laccolith-stock controversy: New results from the 1158 southern Henry Mountains, Utah. Geological Society of America Bulletin 100 1159 (1):117-139. 1160 Jamtveit B, Svensen H, Podladchikov YY, Planke S (2004) Hydrothermal vent complexes 1161 associated with sill intrusions in sedimentary basins. Physical geology of high-level 1162 magmatic systems 234:233-241 1163 Jerram DA, Bryan SE (2015) Plumbing systems of shallow level intrusive complexes. In 1164 Physical Geology of Shallow Magmatic Systems (pp. 39-60). Springer, Cham. 1165 Johnson AM, Pollard DD (1973) Mechanics of growth of some laccolithic intrusions in the 1166 Henry mountains, Utah, I: field observations, Gilbert's model, physical properties and 1167 flow of the magma. Tectonophysics 18 (3):261-309 1168 Jolly R, Sanderson DJ (1995) Variation in the form and distribution of dykes in the Mull 1169 swarm, Scotland. Journal of Structural Geology 17 (11):1543-1557 1170 Jowitt SM, Williamson M-C, Ernst RE (2014) Geochemistry of the 130 to 80 Ma Canadian 1171 High Arctic large igneous province (HALIP) event and implications for Ni-Cu-PGE 1172 prospectivity. Economic Geology 109 (2):281-307 1173 Karson JA, Brooks CK (1999) Structural and magmatic segmentation of the Tertiary East 1174 Greenland volcanic rifted margin. Geological Society, London, Special Publications 1175 164 (1):313-338 1176 Kavanagh JL, Boutelier D, Cruden A (2015) The mechanics of sill inception, propagation and 1177 growth: Experimental evidence for rapid reduction in magmatic overpressure. Earth 1178 and Planetary Science Letters 421:117-128 1179 Kavanagh JL, Burns AJ, Hazim SH, Wood E, Martin SA, Hignett S, Dennis DJ (2018) 1180 Challenging dyke ascent models using novel laboratory experiments: Implications for 1181 reinterpreting evidence of magma ascent and volcanism. Journal of Volcanology and 1182 Geothermal Research 1183 Kendall J-M, Stuart G, Ebinger C, Bastow I, Keir D (2005) Magma-assisted rifting in 1184 Ethiopia. Nature 433 (7022):146-148 1185 Kirton S, Donato J (1985) Some buried Tertiary dykes of Britain and surrounding waters 1186 deduced by magnetic modelling and seismic reflection methods. Journal of the 1187 Geological Society 142 (6):1047-1057 1188 Koch F, Johnson A, Pollard D (1981) Monoclinal bending of strata over laccolithic 1189 intrusions. Tectonophysics 74 (3):T21-T31 1190 Le Gall B, Daoud MA, Maury RC, Rolet J, Guillou H, Sue C (2010) Magma-driven antiform 1191 structures in the Afar rift: The Ali Sabieh range, Djibouti. Journal of Structural 1192 Geology 32 (6):843-854. 1193 Leat PT (2008) On the long-distance transport of Ferrar magmas. Geological Society, 1194 London, Special Publications 302 (1):45-61. 1195 Lister JR, Kerr RC (1991) Fluid‐mechanical models of crack propagation and their 1196 application to magma transport in dykes. Journal of Geophysical Research: Solid 1197 Earth 96 (B6):10049-10077 1198 Macdonald R, Bagiński B, Upton BGJ, Dzierżanowski P, Marshall-Roberts W (2009) The 1199 Palaeogene Eskdalemuir dyke, Scotland: long-distance lateral transport of rhyolitic 1200 magma in a mixed-magma intrusion. Mineralogical Magazine 73 (3):285-300. 1201 Macdonald R, Bagiński B, Upton BGJ, Pinkerton H, MacInnes DA, MacGillivray JC (2010) 1202 The Mull Palaeogene dyke swarm: insights into the evolution of the Mull igneous 1203 centre and dyke-emplacement mechanisms. Mineralogical Magazine 74 (4):601-622. 1204 MacDonald R, Fettes D, Bagiński B (2015) The Mull Paleocene dykes: some insights into the 1205 nature of major dyke swarms. Scottish Journal of Geology 51 (2):116-124 1206 Magee C, Bastow ID, de Vries BvW, Jackson CA-L, Hetherington R, Hagos M, Hoggett M 1207 (2017) Structure and dynamics of surface uplift induced by incremental sill 1208 emplacement. Geology 45 (5):431-434 1209 Magee C, Briggs F, Jackson CA-L (2013) Lithological controls on igneous intrusion-induced 1210 ground deformation. Journal of the Geological Society 170 (6):853-856. 1211 Magee C, Jackson CL, Schofield N (2014) Diachronous sub‐volcanic intrusion along deep‐ 1212 water margins: insights from the Irish Rockall Basin. Basin Research 26 (1):85-105 1213 Magee C, Maharaj SM, Wrona T, Jackson CA-L (2015) Controls on the expression of 1214 igneous intrusions in seismic reflection data. Geosphere:GES01150. 01151 1215 Magee C, Muirhead JD, Karvelas A, Holford SP, Jackson CA, Bastow ID, Schofield N, 1216 Stevenson CT, McLean C, McCarthy W (2016) Lateral magma flow in mafic sill 1217 complexes. Geosphere:GES01256. 01251 1218 Mäkitie H, Data G, Isabirye E, Mänttäri I, Huhma H, Klausen MB, Pakkanen L, Virransalo P 1219 (2014) Petrology, geochronology and emplacement model of the giant 1.37 Ga 1220 arcuate Lake Victoria Dyke Swarm on the margin of a large igneous province in 1221 eastern Africa. Journal of African Earth Sciences 97:273-296 1222 Malehmir A, Bellefleur G (2010) Reflection seismic imaging and physical properties of base- 1223 metal and associated iron deposits in the Bathurst Mining Camp, New Brunswick, 1224 Canada. Ore Geology Reviews 38 (4):319-333 1225 Malthe-Sørenssen A, Planke S, Svensen H, Jamtveit B (2004) Formation of saucer-shaped 1226 sills. In: Breitkreuz C, Petford N (eds) Physical geology of high-level magmatic 1227 systems. Geological Society, London, Special Publications, vol 234. Geological 1228 Society, London, Special Publications, pp 215-227 1229 Manga M, Michaut C (2017) Formation of lenticulae on Europa by saucer-shaped sills. Icarus 1230 286:261-269 1231 Mark N, Schofield N, Pugliese S, Watson D, Holford S, Muirhead D, Brown R, Healy D 1232 (2017) Igneous intrusions in the Faroe Shetland basin and their implications for 1233 hydrocarbon exploration; new insights from well and seismic data. Marine and 1234 Petroleum Geology 1235 Marsh B (2004) A magmatic mush column rosetta stone: the McMurdo Dry Valleys of 1236 Antarctica. Eos, Transactions American Geophysical Union 85 (47):497-502 1237 Mastin LG, Pollard DD (1988) Surface deformation and shallow dike intrusion processes at 1238 Inyo Craters, Long Valley, California. Journal of Geophysical Research: Solid Earth 1239 93 (B11):13221-13235 1240 McBride JH, William Keach R, Leetaru HE, Smith KM (2018) Visualizing Precambrian 1241 basement tectonics beneath a carbon capture and storage site, Illinois Basin. 1242 Interpretation 6 (2):T257-T270 1243 Mège D, Cook AC, Garel E, Lagabrielle Y, Cormier MH (2003) Volcanic rifting at Martian 1244 grabens. Journal of Geophysical Research: Planets 108 (E5) 1245 Menand T (2008) The mechanics and dynamics of sills in layered elastic rocks and their 1246 implications for the growth of laccoliths and other igneous complexes. Earth and 1247 Planetary Science Letters 267 (1-2):93-99. doi:10.1016/j.epsl.2007.11.043 1248 Morgan J, Warner M, Bell R, Ashley J, Barnes D, Little R, Roele K, Jones C (2013) Next- 1249 generation seismic experiments: wide-angle, multi-azimuth, three-dimensional, full- 1250 waveform inversion. Geophysical Journal International 195 (3):1657-1678. 1251 Morgan S, Stanik A, Horsman E, Tikoff B, de Saint Blanquat M, Habert G (2008) 1252 Emplacement of multiple magma sheets and wall rock deformation: Trachyte Mesa 1253 intrusion, Henry Mountains, Utah. Journal of Structural Geology 30 (4):491-512. 1254 Muirhead J, Kattenhorn S, Lee H, Mana S, Turrin B, Fischer T, Kianji G, Dindi E, Stamps D 1255 (2016) Evolution of upper crustal faulting assisted by magmatic volatile release 1256 during early-stage continental rift development in the East African Rift. Geosphere 12 1257 (6):1670-1700 1258 Muirhead JD, Airoldi G, Rowland JV, White JD (2012) Interconnected sills and inclined 1259 sheet intrusions control shallow magma transport in the Ferrar large igneous province, 1260 Antarctica. Geological Society of America Bulletin 124 (1-2):162-180 1261 Muirhead JD, Airoldi G, White JD, Rowland JV (2014) Cracking the lid: Sill-fed dikes are 1262 the likely feeders of flood basalt eruptions. Earth and Planetary Science Letters 1263 406:187-197 1264 Muirhead JD, Kattenhorn SA, Le Corvec N (2015) Varying styles of magmatic strain 1265 accommodation across the East African Rift. Geochemistry, Geophysics, Geosystems 1266 16(8), pp.2775-2795. 1267 Neumann, E.R., Svensen, H., Galerne, C.Y. and Planke, S., 2011. Multistage evolution of 1268 dolerites in the Karoo large igneous province, Central South Africa. Journal of 1269 Petrology, 52(5), pp.959-984. 1270 Passey S, Hitchen K (2011) Cenozoic (igneous). Geology of the Faroe-Shetland Basin and 1271 adjacent areas. British Geological Report Survey Research Report, RR/11/01; 1272 Jarofeingi Research Report RR/11/01 1273 Patterson CW, Ernst RE, Samson C (2016) Pit Chains Belonging to Radiating Graben‐Fissure 1274 Systems on Venus: Model for Formation during Lateral Dyke Injection. Acta 1275 Geologica Sinica (English Edition) 90 (s1):143-144 1276 Peron-Pinvidic G, Shillington DJ, Tucholke BE (2010) Characterization of sills associated 1277 with the U reflection on the Newfoundland margin: evidence for widespread early 1278 post-rift magmatism on a magma-poor rifted margin. Geophysical Journal 1279 International 182 (1):113-136 1280 Phillips TB, Magee C, Jackson CA-L, Bell RE (2017) Determining the three-dimensional 1281 geometry of a dike swarm and its impact on later rift geometry using seismic 1282 reflection data. Geology 46 (2):119-122 1283 Pirajno F, Hoatson DM (2012) A review of Australia's Large Igneous Provinces and 1284 associated mineral systems: implications for mantle dynamics through geological 1285 time. Ore Geology Reviews 48:2-54 1286 Planke S, Rasmussen T, Rey SS, Myklebust R (2005) Seismic characteristics and distribution 1287 of volcanic intrusions and hydrothermal vent complexes in the Vøring and Møre 1288 basins. In: Doré AG (ed) Petroleum Geology: North-West Europe and Global 1289 Perspectives - Proceddings of the 6th Petroleum Geology Conference. Geological 1290 Society, London, pp 833-844. 1291 Planke S, Svensen H, Myklebust R, Bannister S, Manton B, Lorenz L (2015) Geophysics and 1292 Remote Sensing. In. Springer Berlin Heidelberg, Berlin, Heidelberg, pp 1-16. 1293 Pollard DD, Delaney PT, Duffield WA, Endo ET, Okamura AT (1983) Surface deformation 1294 in volcanic rift zones. Tectonophysics 94 (1-4):541-584 1295 Pollard DD, Johnson AM (1973) Mechanics of growth of some laccolithic intrusions in the 1296 Henry Mountains, Utah, II: bending and failure of overburden layers and sill 1297 formation. Tectonophysics 18 (3):311-354 1298 Polteau S, Ferré EC, Planke S, Neumann ER, Chevallier L (2008a) How are saucer-shaped 1299 sills emplaced? Constraints from the Golden Valley Sill, South Africa. Journal of 1300 Geophysical Research 113 (B12). 1301 Polteau S, Mazzini A, Galland O, Planke S, Malthe-Sørenssen A (2008b) Saucer-shaped 1302 intrusions: occurrences, emplacement and implications. Earth and Planetary Science 1303 Letters 266 (1):195-204 1304 Preston RJ, Bell BR, Rogers G (1998) The Loch Scridain xenolithic sill complex, Isle of 1305 Mull, Scotland: fractional crystallization, assimilation, magma-mixing and crustal 1306 anatexis in subvolcanic conduits. Journal of Petrology 39 (3):519-550 1307 Rabbel O, Galland O, Mair K, Lecomte I, Senger K, Spacapan JB, Manceda R (2018) From 1308 field analogues to realistic seismic modelling: a case study of an oil-producing 1309 andesitic sill complex in the Neuquén Basin, Argentina. Journal of the Geological 1310 Society:jgs2017-2116 1311 Reeves C (2000) The geophysical mapping of Mesozoic dyke swarms in southern Africa and 1312 their origin in the disruption of Gondwana. Journal of African earth sciences 30 1313 (3):499-513 1314 Reeves J, Magee C, Jackson CAL (2018) Unravelling intrusion-induced forced fold 1315 kinematics and ground deformation using 3D seismic reflection data. Volcanica 1 1316 (1):17 1317 Reynolds P, Holford S, Schofield N, Ross A (2017) The shallow depth emplacement of mafic 1318 intrusions on a magma-poor rifted margin: An example from the Bight Basin, 1319 Southern Australia. Marine and Petroleum Geology 88:605-616 1320 Rivalta E, Taisne B, Bunger A, Katz R (2015) A review of mechanical models of dike 1321 propagation: Schools of thought, results and future directions. Tectonophysics 638:1- 1322 42 1323 Rowland J, Baker E, Ebinger C, Keir D, Kidane T, Biggs J, Hayward N, Wright T (2007) 1324 Fault growth at a nascent slow‐spreading ridge: 2005 Dabbahu rifting episode, Afar. 1325 Geophysical Journal International 171 (3):1226-1246 1326 Rubin AM (1992) Dike‐induced faulting and graben subsidence in volcanic rift zones. 1327 Journal of Geophysical Research: Solid Earth 97 (B2):1839-1858 1328 Rubin AM (1995) Propogation of magma-filled cracks. Annual Review of Earth and 1329 Planetary Sciences 23:49 1330 Schmiedel T, Kjoberg S, Planke S, Magee C, Galland O, Schofield N, Jackson CA-L, Jerram 1331 DA (2017) Mechanisms of overburden deformation associated with the emplacement 1332 of the Tulipan sill, mid-Norwegian margin. Interpretation 5 (3):SK23-SK38 1333 Schofield N, Alsop I, Warren J, Underhill JR, Lehné R, Beer W, Lukas V (2014) Mobilizing 1334 salt: Magma-salt interactions. Geology:G35406. 35401 1335 Schofield N, Heaton L, Holford SP, Archer SG, Jackson CA-L, Jolley DW (2012a) Seismic 1336 imaging of ‘broken bridges’: linking seismic to outcrop-scale investigations of 1337 intrusive magma lobes. Journal of the Geological Society 169 (4):421-426 1338 Schofield N, Holford S, Millett J, Brown D, Jolley D, Passey SR, Muirhead D, Grove C, 1339 Magee C, Murray J, Hole M, Jackson CAL, Stevenson C (2017) Regional magma 1340 plumbing and emplacement mechanisms of the Faroe-Shetland Sill Complex: 1341 implications for magma transport and petroleum systems within sedimentary basins. 1342 Basin Research 29 (1):41-63. 1343 Schofield NJ, Brown DJ, Magee C, Stevenson CT (2012c) Sill morphology and comparison 1344 of brittle and non-brittle emplacement mechanisms. Journal of the Geological Society 1345 169 (2):127-141. 1346 Skogseid J, Pedersen T, Eldholm O, Larsen BT (1992) Tectonism and magmatism during NE 1347 Atlantic continental break-up: the Voring Margin. Geological Society, London, 1348 Special Publications 68 (1):305-320. 1349 Smallwood JR, Maresh J (2002) The properties, morphology and distribution of igneous sills: 1350 modelling, borehole data and 3D seismic from the Faroe-Shetland area. In: Jolley 1351 DW, Bell BR (eds) The North Atlantic Igneous Province: Stratigraphy, tectonic, 1352 Volcanic and Magmatic Processes, vol 197. vol 1. Geological Society, London, 1353 Special Publications, pp 271-306. 1354 Spacapan JB, Galland O, Leanza HA, Planke S (2017) Igneous sill and finger emplacement 1355 mechanism in shale-dominated formations: a field study at Cuesta del Chihuido, 1356 Neuquén Basin, Argentina. Journal of the Geological Society 174 (3):422-433 1357 Speight J, Skelhorn R, Sloan T, Knaap R (1982) The dyke swarms of Scotland. Igneous rocks 1358 of the British Isles:449-459 1359 Storey M, Duncan RA, Tegner C (2007) Timing and duration of volcanism in the North 1360 Atlantic Igneous Province: Implications for geodynamics and links to the Iceland 1361 hotspot. Chemical Geology 241 (3-4):264-281. 1362 Svensen H, Corfu F, Polteau S, Hammer Ø, Planke S (2012) Rapid magma emplacement in 1363 the Karoo Large Igneous Province. Earth and Planetary Science Letters 325-326:1-9. 1364 Svensen H, Jamtveit B, Planke S, Chevallier L (2006) Structure and evolution of 1365 hydrothermal vent complexes in the Karoo Basin, South Africa. Journal of the 1366 Geological Society 163:11 1367 Svensen H, Planke S, Chevallier L, Malthe-Sørenssen A, Corfu F, Jamtveit B (2007) 1368 Hydrothermal venting of greenhouse gases triggering Early Jurassic global warming. 1369 Earth and Planetary Science Letters 256 (3-4):554-566 1370 Svensen H, Planke S, Malthe-Sorenssen A, Jamtveit B, Myklebust R, Rasmussen Eidem T, 1371 Rey SS (2004) Release of methane from a volcanic basin as a mechanism for initial 1372 Eocene global warming. Nature 429 (6991):542-545 1373 Svensen HH, Polteau S, Cawthorn G, Planke S (2015) Sub-volcanic intrusions in the Karoo 1374 basin, South Africa. 1375 Tentler T (2005) Propagation of brittle failure triggered by magma in Iceland. 1376 Tectonophysics 406 (1):17-38 1377 Thomson K (2005) Volcanic features of the North Rockall Trough: application of 1378 visualisation techniques on 3D seismic reflection data. Bull Volcanol 67 (2):116-128. 1379 Thomson K (2007) Determining magma flow in sills, dykes and laccoliths and their 1380 implications for sill emplacement mechanisms. Bull Volcanol 70 (2):183-201. 1381 Thomson K, Hutton D (2004) Geometry and growth of sill complexes: insights using 3D 1382 seismic from the North Rockall Trough. Bull Volcanol 66 (4):364-375. 1383 Torsvik TH, Smethurst MA, Burke K, Steinberger B (2008) Long term stability in deep 1384 mantle structure: Evidence from the~ 300 Ma Skagerrak-Centered Large Igneous 1385 Province (the SCLIP). Earth and Planetary Science Letters 267 (3):444-452 1386 Townsend MR, Pollard DD, Smith RP (2017) Mechanical models for dikes: a third school of 1387 thought. Tectonophysics 703:98-118 1388 Trude J, Cartwright J, Davies RJ, Smallwood JR (2003) New technique for dating igneous 1389 sills. Geology 31:4 1390 van Wyk de Vries B, Márquez A, Herrera R, Bruña JG, Llanes P, Delcamp A (2014) Craters 1391 of elevation revisited: forced-folds, bulging and uplift of volcanoes. Bull Volcanol 76 1392 (11):1-20 1393 Wall M, Cartwright J, Davies R, McGrandle A (2010) 3D seismic imaging of a Tertiary Dyke 1394 Swarm in the Southern North Sea, UK. Basin Research 22 (2):181-194. 1395 Wilson L, Head JW (2002) Tharsis‐radial graben systems as the surface manifestation of 1396 plume‐related dike intrusion complexes: Models and implications. Journal of 1397 Geophysical Research: Planets 107 (E8) 1398 Wilson PI, McCaffrey KJ, Wilson RW, Jarvis I, Holdsworth RE (2016) Deformation 1399 structures associated with the Trachyte Mesa intrusion, Henry Mountains, Utah: 1400 Implications for sill and laccolith emplacement mechanisms. Journal of Structural 1401 Geology 87:30-46 1402 Wingate MT, Pirajno F, Morris PA (2004) Warakurna large igneous province: a new 1403 Mesoproterozoic large igneous province in west-central Australia. Geology 32 1404 (2):105-108 1405 Wyche S, Pawley M, Chen S, Ivanic T, Zibra I, Van Kranendonk M, Spaggiari C, Wingate M 1406 Geology of the northern Yilgarn Craton. In: Youanmi and Southern Carnarvon 1407 Seismic and Magnetotelluric (MT) Workshop 2013: Extended Abstracts, Compiled by 1408 S Wyche, TJ Ivanic and I Zibra. Geological Survey of Western Australia. Record, 1409 2013. 1410 Wyrick D, Ferrill DA, Morris AP, Colton SL, Sims DW (2004) Distribution, morphology, 1411 and origins of Martian pit crater chains. Journal of geophysical research 109 1412 (E6):E06005 1413 Wyrick DY, Smart KJ (2009) Dike-induced deformation and Martian graben systems. 1414 Journal of Volcanology and Geothermal Research 185 (1-2):1-11. 1415 Zaleski E, Eaton DW, Milkereit B, Roberts B, Salisbury M, Petrie L (1997) Seismic 1416 reflections from subvertical diabase dikes in an Archean terrane. Geology 25 (8):707- 1417 710 1418 F H S 180o 0o Mac Y VM FS 60oN U NAIP J Si M IR Mu G A C Mt 30oN CAMP

AA D

V

E K o W P-E Ga 30 S Ka Md

Selected plume centres, dyke swarms, and sill-complexes Plume centres Dyke swarms Sill-complexes Cenozoic Late Proterozoic Linear Sill-complex Mesozoic Middle Proterozoic Radiating Paleozoic Early Proterozoic Circumferential 30 m (A) Karoo Africa ~1000 km (B) Basin Sills M Mad. Inclined sheets South India America

Beacon Supergroup

Antarctica (C) N GS

HS GONDWANA MVS Tasmania ~180 Ma GVS MSS GVD New Zealand Australia GoogleTM earth Distribution of sill-complexes and Ferrar LIP Image Landsat/Copernicus Beacon Supergroup (+ correlatives) Karoo LIP Image © 2017 DigitalGlobe Proposed plume head positions Dyke swarms Image © 2017 CNES/Astrium 5 km (A) (C)

N Fig. 3B viewing direction GoogleTM earth Image © DigitalGlobe Fig. 3C Image © CNES/Astrium 5 km

TM (B) N Google earth N Sill contraction (m) Image © DigitalGlobe 2 km ~0.5 km Image © CNES/Astrium 0 -2 -4 -6 (D) Alu Alu

Host rock Normal fault Magma pulse 1-4 Onlap Not to scale Seismic reflection data Field examples 3D/Map view Cross-section SW NE (A) -2.2 Broken bridges (s) -2.4 Section TWT

Sill N Sill

VE ≈ 5 Strat-con. sill VE ≈ 5 ~1.5 km 0.5 s (TWT) 2 km ~0.1 km (B) TWT (s) W E - + Section

-2

-3 Sill

TWT (s) TWT Sill N N

Saucer-shaped sill VE ≈ 5 VE ≈ 2

~0.5 km 0.2 s (TWT) -2.73 1 km ~2 km (C) NW SE Inclined sheet -4.8

-5.0 Sill 0.2 s (TWT)

TWT (s) TWT -5.2 N VE ≈ 2 Inclined sheet Section ~1 km VE ≈ 5 1 km ~100 m (D) A A’ 0.5 km VE ≈ 2 S1 0.1 s (TWT)

B B’ 0.5 km A’ VE ≈ 2 B’ B2 C’ B1 0.1 s (TWT)

A B N C C’ C 0.5 km VE ≈ 2 S2 1 1 km B3 S

S Step 0.1 s (TWT) B Bridge structure (A) S Onlap Truncation Downlap N

Top Lower Eocene

Top Paleocene

Top Cretaceous

Sill

VE ≈ 5

0.25 s (TWT) 2 km (B) 6204000 6212000 6220000 6228000 6236000 N 44600 5 km

Fig. 5A

45400 Number of intrusions stacked vertically 1 2 3 4 5 6

(C) t0 - Initial intrusion and instantaneous fold growth Key New magma pulse volume

t0 magma pulse volume

t1 magma pulse volume Pre-intrusion t1 - Continued intrusion and onset of fold coalescence strata

t1 syn-folding strata and onlap

t2 syn-folding strata and onlap Individual forced fold base datum t2 - Increased fold growth accommodating later magma pulses Compound forced fold base datum (1) Compound forced fold base datum (2) NB: Intrusion volume is to highlight different magma pulses, not the distribution of magma within single sills from each pulse. (A) Faroe Islands 4°W 2°W Fig. 6B 62°N

Fig. 6A SBF SB Shetland Islands

Scotland FLB

FBF

61°N

FBF

100 km Shetland Structural high Lineament or Faroe-Shetland Sill transfer zone Basin limit Stratigraphic height Basin - + Normal Fault Magma flow Magma input zone

(B) NW 214/4-1 214/9-1 208/21-1 SE 1 s (TWT)

20 km NW 214/4-1 214/9-1 208/21-1 SE 1 s (TWT)

20 km Paleogene Lavas and hyaloclastites Magma flow Normal fault Cretaceous Sills Magma input zone (A) 0° 3°E 6°E 9°E (B) Vøring Marginal Dome- Crater- Eye- 67°N High shaped shaped shaped

Vøring Basin 65°N Sill-complex Møre extent Marginal High

63°N Møre Basin Onlap Downlap Truncation 15 10 TWT (s) 15 10 TWT (s) 5 0 5 0 CDP CDP 10GA-YU1 10GA-YU2 10GA-YU2 3000 5000 4000 6000 5000 Fault Form line Sills 7000 6000 8000 Base maficrocks Base felsicvolcanicrocks Granite Base Cenozoic 7000 9000 8000 10000 9000 Moho Moho 1 1000 Base ultramaficzone Base lowerzone Base roofzone 10000 12000 1 Upper mantle 1000 Upper mantle 13000

12000

Moho T Base middlecrust Base uppercrust op Mohotransitionzone 14000

13000

15000

14000

16000 15000 17000 25 km 25 km 16000 (A) 200000 250000 300000 350000 20 km 5103 184 54 450000 8 -18 -37 -54 Northern Ireland -69 nT -85 -104 AP -125 -154 -261 -4977 400000

Lough Belfast Neagh

350000 N

20 km SG MM

©Crown Copyright 2011 (B) 200000 250000 300000 350000 Dyke

450000 swarms

Oligocene clays Sills TVF LF Mourne Mountains Slieve Guillion Upper Basalt Formation CF Lower Basalt Formation Newry Igneous SF

400000 Complex

CSF OF

350000 N CVF

CLF NF 20 km

©Crown Copyright 2011 (A) Linear dyke swarms

σ σ3 3

σ3 Dykes Rift zone Plume centre Radiating dyke swarms (B) Sub-swarms

σ3

Dykes rotate as far-field stress dominates (C) Circumferential dyke swarms

Plume head limit Initial Final Predicted Initial Final Predicted growth geometry displacement growth geometry displacement (i) Displacement Graben (iii) Displacement Depth Depth

(ii) Displacement (iv) Displacement Depth Depth Karst dissolution Dyke

Empty lava tube

Dyke

Volume reduction of host rock / dyke

0.5 m Dyke -700 (A) N S Upper tip -800 Lower line Bedding-related -900 Cretaceous reflectivity 1.0 -1000 Carboniferous - Permian Jurassic -1100

Dyke swarm -1200

Triassic -1300

-1400 Base Zechstein Unconformity -1500 Base Carboniferous- Saalian Unconformity TWT (s) TWT -1600

Permian -1700

1.8 -1800

-1900 Dyke-related reflections -2000

Upper -2100 Top dyke swarm Permian 2.2 1 km -2200

-2300 N Dyke upper tip line S 300–292 Ma (B) Erosion surface (C) Dykes Oslo Graben Volcanics NOR Skagerrak Dyke swarm 296 Ma Graben emplacement SWE Farsund ScLIP 10 km Dyke Swarm Permian 297 Ma Midland Valley Dyke Suite 302–292 Ma Cretaceous Jurassic DK

Triassic 300, 287 Ma Clausen et al 299 Ma (2016)

Dyke swarm Early UK 302–297 Ma rotation Cretaceous 250 km GER (A) (B) Top Cretaceous Fig. 14C

UK Crater

57 ° N Mull Intrusive Complex 0.5 s TWTT 0.5

Dyke(s) Dykes

2.5 km 55 ° N Fig. 14B

100 km 4°W 0° (C) 1 km 1 km Top Cretaceous

0.1 s TWTT 0.1 s TWTT Crater

Zone of disturbance Dyke Onlap (A) 0.4 s TWTT fault induced Dyke- 2 km Dyke crater Pit chain (B) fault Dyke-induced Fig. 15A 2 km crater Pit chain