The Melville Bugt Dyke Swarm across SE Greenland: A closer link to Mesoproterozoic AMCG-complexes

Accepted Manuscript The Melville Bugt Dyke Swarm across SE Greenland: A closer link to Mesoproterozoic AMCG-complexes M.B. Klausen, M.K.M. Nilsson PII: DOI: Reference:

S0301-9268(17)30551-X https://doi.org/10.1016/j.precamres.2018.06.001 PRECAM 5103

To appear in:

Precambrian Research

Received Date: Revised Date: Accepted Date:

30 September 2017 17 May 2018 2 June 2018

Please cite this article as: M.B. Klausen, M.K.M. Nilsson, The Melville Bugt Dyke Swarm across SE Greenland: A closer link to Mesoproterozoic AMCG-complexes, Precambrian Research (2018), doi: https://doi.org/10.1016/ j.precamres.2018.06.001

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The Melville Bugt Dyke Swarm across SE Greenland: A closer link to Mesoproterozoic AMCG-complexes M.B. Klausen1 and M.K.M. Nilsson2,3,4 1)

Department of Earth Sciences, Stellenbosch University, Private Bag X1, Matieland 7602, RSA 2)

12, 223 62 Lund, Sweden

3)

Department of Geoscience, Swedish Museum of Natural History, Box 50007, SE-10405, Stockholm, Sweden

4)

Department of Earth and Atmospheric Sciences, University of Alberta, Edmonton, Alberta T6G 2E3, Canada

Abstract Following upon previous work on the 1630 Ma Melville Bugt Dyke Swarm (MBDS) along the NW coast of Greenland, this paper confirms its proposed continuation in SE Greenland, into an extraordinary > 2000 km long trans-Greenlandic dyke swarm. This correlation is not only based on the w

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1630 ± 4 Ma U – Pb baddeleyite age

determination and similar bulk rock compositions, as well as distinctive outcrop characteristics. A greater number of thinner dykes across SE Greenland supports a previous notion of MBDS being emplaced laterally from a more southerly located magma chamber; whereas, a doubling of both individual and cumulative dyke thicknesses towards the north may relate to how the swarm’ bladeshaped dyke geometries are exposed in greater abundance farther from its source. Remarkably homogenous yet differentiated, transitional to trachybasaltic compositions indicate a single magma chamber source that was buffered by between 4 – 7 wt% MgO, through magma replenishments and cotectic plagioclase-olivine fractionation at lower crustal depths for >13 ± 6 Myrs (published age range). Geochemical signatures –

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Nb-Ta spikes and

otherwise slightly enriched HFSE patterns – bear a conspicuous resemblance to >100 Myr-older lateto post-Ketilidian appinites (orogenic lamprophyres), which are both thought to have been derived from a common highly metasomatised, young sub-continental lithospheric mantle. The dry mineralogy of the MBDS and a geologically based Nuna supercontinent reconstruction is more consistent with a breakup setting, which together with neighbouring coeval AMCG-complexes may arguably combine into an unusual LIP. However, such a relatively short-lived LIP-setting may be questioned if the plagioclase-phyric MBDS was derived from a similar deep crustal magma chamber source, as proposed for four Mesoproterozoic cycles of AMCG-emplacements during a >500 Myr Nuna – Rodinia supercontinental transition. Consequently, arguing for a more long-lived active continental margin, as suggested by palaeomagnetically based Nuna supercontinent reconstructions. 1. Introduction Precambrian giant mafic dyke swarms are exceptional structures that invites the question of how such voluminous mantle-derived magmas formed and were emplaced across continents. The common

interpretation is that these represent feeder systems below large igneous provinces (LIPs; e.g., Ernst, 2014 and references therein), not only because of the large magma volumes that giant dykes could have funnelled but also because coherent swarms give the impression of having been emplaced within relatively short time periods or in distinct pulses (e.g., < 5 Ma; Bryan and Ernst, 2008). Such rapid generation and fluxing of large volumes of mantle-derived magmas are typically attributed to anomalously hot asthenospheric material in the form of a mantle plume (e.g., Morgan, 1971; White and McKenzie, 1989; Campbell, 2007). This, despite difficulties in finding evidence for such a source below ancient LIPs, where plumes are no longer present, an absence of related oceanic hot spot tracks have been subducted, and that most of any continental flood basalts have been eroded. Various research methods are being used to test or even elucidate the often enigmatic origins of many Precambrian giant mafic dyke swarms. First and foremost, high precision and accuracy geochronology not only allows a swarm to be linked to other fragmented parts, and thereby better constrain the spatial extent of the magmatic event, but also determines the minimum duration of the magmatic event (e.g., Bleeker and Ernst, 2006; Söderlund et al., 2016). Secondly, palaeopoles on magnetite-bearing mafic dykes add further constraints on both palaeo-latitudes and orientations of craton fragments, used for the reconstructions of ancient plate tectonic configurations (e.g., Zhao et al., 2004). Finally, more traditional structural and petrological methods provide independent constraints on the above, as well as elucidate how mantle-derived magmas formed (i.e., petrogenesis), evolved (typically through fractional crystallisation), and became mechanically emplaced. Such a multi-disciplinary approach is often necessary, in order to properly test the tectono-magmatic setting of major magmatic events and result in independent petrogenetic inferences that are more case-specific than model-driven. More specifically, the Mesoproterozoic Nuna – Rodinia supercontinental transition is to large parts dominated by massif anorthosites and equally exotic associated felsic intrusions, including relatively ‘

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rapakivi granites with plagioclase-overgrown alkali

feldspars, collectively referred to as AMCG-complexes (e.g., Emslie et al., 1994). AMCG-complexes are thought to be derived from large and deep crustal mafic magma chambers, which produced both plagioclase cumulates for the massif anorthosites as well as lower crustal melts for the associated felsic intrusions (e.g., Ashwal, 1993; Dempster et al., 1999). The tectonic setting for several, prolonged and apparently cyclic emplacements of AMCG-complexes, during such a prolonged (>500 Myr) Nuna – Rodinia supercontinental transition, however, is still debated. While most researchers seem to favour a poorly constrained setting of repetitive post-orogenic rifting, during this rather unique period of Proterozoic AMCG-magmatism

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(e.g., Emslie et al., 1994), such an

interpretation conflicts with the need for a full Wilson cycle of Nuna breakup and Rodinia amalgamation to have occurred (e.g., Windley, 1993). In this paper, we firstly provide an U-Pb age that links a prominent dyke across NW Greenland (Kalsbeek and Taylor, 1986) to the thickest dyke across SE Greenland, and thereby its associated Melville Bugt Dyke Swarm (MBDS; Nielsen, 1990; into an impressive >2000 km long trans-

Greenlandic structure (Fig. 1a), as proposed by Halls et al., (2011). We then provide further structural, petrographic and bulk rock geochemical data on this SE Greenland segment of the MBDS, arguing for a single more southerly located – yet remarkably long-lived – central magma chamber source near the base of the Ketilidian orogenic crust. Structural evidence further suggest that compositionally homogenous dykes were injected up to 2000 km laterally northwards from such a lower crustal magma chamber and – as suggested by available ages – at a remarkably low frequency. Finally, we discuss model implications of these new results, elaborating on what role the MBDS might have played amongst coeval AMCG-complexes, within an unresolved tectonic setting. 2. Previous work on MBDS across NW Greenland The Geological Survey of Greenland (then GGU, now GEUS) initially mapped and sampled much of the NNW-SSE trending MBDS along the northern west coast of Greenland (NW Greenland in this paper). Kalsbeek and Taylor (1986) focused on a ~100 m thick and >400 km long dyke (cf., Fig. 1a), how its compositional zonation contrasted with its along-strike homogeneity, and how Sr and Pbisotopes recorded remarkably little crustal contamination. Nielsen (1990) followed this up by reviewing the full extent of the NW Greenland MBDS, which was estimated to consist of at least 33 dykes, from Disko Bugt in the south to Ellesmere Island in the north (between 69-80˚N). Th publications, based on field observations and 95 bulk rock geochemical analyses (69 from the same >400 km long dyke!) describe a MBDS with a remarkably restricted compositional range of alkali olivine dolerites, with relatively low Di:Ol-Hy ratios that conform to magmas equilibrating at 9 kbar (Thompson, 1982). This arguably reflects an equally homogenous mantle-derived magma source near the base of the crust, which early authors found difficult to reconcile with an, at the time, >1300 km long dyke swarm. The Mesoproterozoic MBDS – in NW Greenland cutting the youngest Palaeoproterozoic (~1.85 Ga) rocks of the Nagssugtoqidian orogeny and overlain by the oldest Neoproterozoic (~1.2 Ga) Thule Group cover rocks (Nielsen, 1990) – was initially dated by a Rb-Sr whole rock isochron to 1645 ± 35 Ma (Kalsbeek and Taylor, 1986). This age was refined by U-Pb baddeleyite age determinations on four dykes: 1622 ± 3 Ma for a 200 m wide dyke in Olrik fjord, 1629 ± 1 Ma for the >400 km long and 70-100 m wide Torsukattak dyke previously dated by Kalsbeek and Taylor (1986), 1632 ± 1 Ma for a 100 m thick Qeqertaq dyke, and 1635 ± 3 Ma for a 135 m wide dyke in Ryder Isfjord (Halls et al., 2011; Fig. 1a). These results suggested a slightly younger, yet rather prolonged emplacement period of at least 13 ± 6 Myr, which is consistent with both magnetically normal and reversed dykes. Halls et ’ (2011) palaeomagnetic results also orientated the MBDS so that its southern extension, at ~1630 Ma, projected towards a roughly coeval Fennoscandian Rapakivi Granite Province (as shown later in Fig. 13b). Such a link was further strengthened by their discovery (in GoogleTM Earth) of potential MBDS

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Greenland extension and possible southern link to a coeval magmatic province along a globally extensive pre-Grenvillian – Ketilidian – Fennoscandian Belt is what this paper investigates further.

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Figure 1: (a) Geology of Greenland (modified from Dawes, 2009), onto which simplified traces of coast-parallel Tertiary dyke swarms and igneous centres (Klausen and Larsen, 2002; Weatherly et al., 2016) and individual MBDS-dykes across Greenland. Note that (a) depicts a more conservative correlation of dyke segments than Nielsen (1990) into no more than 26 coherent dykes across NW Greenland. Selected sample localities in NW Greenland indicated by red (dated1 samples by Halls et al., 2011), yellow and white dots (dyke centres2 and margins3, respectively, by Kalsbeek and Taylor, 1986). T = Tasiilaq, where there are two N-S trending and MBDS-like dykes. Likewise, Chadwick (1969) mapped some WNW-ESE trending dykes across SW Greenland (C) with Rb-Sr ages of 1.2-1.6 Ga. (b) Geological map of Archaean North Atlantic Craton, as exposed along the SE coast of Greenland, between margins of the Nagssugtoqidian and Ketilidian Deformation Fronts to the far north and south, respectively, of this map. Sample localities indicated by orientated orange diamonds are only labelled outside more detailed map areas, shown in Figures 2-4. North is roughly up in (a) and more precisely indicated by longitudinal grid in (b).

3. Field relationships across SE Greenland In 2012, the authors of this paper mapped, studied and sampled potential MBDS dykes across the Archaean North Atlantic Craton, where these are exposed along the south-eastern coast of Greenland (e.g., Kolb et al., 2013 and references therein; Fig. 1b). Despite logistical challenges, almost complete mapping and sampling traverses across this NNW-SSE trending dyke swarm were accomplished across three areas, referred to as Umivik, Skjoldungen and Timmiarmiit (see Fig. 1b for locations). Despite not mapping a large area between Skjoldungen and Timmiarmiit, it will be shown that our field work covered much of what appears to be a well constrained ~250 km wide dyke swarm. More detailed maps of each of the three areas in Figures 2-4, respectively, illustrate field relationships, c

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important outcrop features. 3.1. Umivik As shown in Figure 1, the northernmost Umivik area covers the easternmost margin of the studied dyke swarm. There are a few compositionally similar, yet peripheral, dykes ~250 km farther northh T

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(“T”

Fig. 1a), but these are not included in this study. Thus, the most

easterly located dyke, represented by samples 656321,-33 and -34, roughly follows the coast and partly overlaps with the southernmost known extent of more coast-parallel dykes that presumably were emplaced during the Tertiary opening of the Atlantic (cf., Fig. 1b; Weatherly et al., 2016). These ‘T

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from other dykes on the basis of their unique OIB-like

signatures, which are tentatively related to a more northerly located proto-Icelandic hot spot source (Klausen et al., 2018), and do otherwise not appear to interfere with the studied dyke swarm that appears to extend more obliquely NNW

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. Figure 2

summarises some interesting field relationships within this most northerly and eastern part of the potential MBDS segment in SE Greenland, where at least ten individual dykes are exposed along the Umivik traverse (more easterly located dykes are indicated on Fig. 1b). In Figure 2(a), a 500 m tall cliff surface shows how NNW-SSE trending and more brown-weathered mafic dykes typically cross-cut not only the pale Archaean orthogneissic basement, but also E-W trending and more amphibolitised, black pre-Nagssugtoqidian and Palaeoproterozoic meta-dykes. These are first-order field relationships that help recognise potential MBDS dykes from a distance. At Umivik, a rare local reddish colouring of the pale host orthogneiss was observed, which only formed along the outer sides of a sub-parallel dykelet pair, and resembles how feldspars in such granitoid host rocks alter within brittle shear zones (possibly through hydrothermal oxidation of tiny magnetite inclusions). Therefore, this discoloration may somehow be related to a nearby WSW-ENE trending transverse fault behind this exposure, which displaces these dykes dextrally by 71 m (cf., Fig. 2d). Such apparent dextral transform faults represent the only obvious deformation of an otherwise well preserved dyke swarm, observed along the entire extent of the south east coast of Greenland, and are

tentatively related to a northward propagating proto-Atlantic opening (e.g., Srivastava and Tapscott, 1986)

Figure 2: The Umivik area. (a) ~500 m-tall cliff surface exposes a sub-vertical, sinistrally offset and brownweathered dyke that cuts obliquely across Precambrian TTG-gneiss basement structures and fabrics, as well as more E-W trending and darker meta-dykes. (b) A close pair of sub-parallel and -vertical dykes, which never seem to merge or taper off like overlapping segments along a single offset dyke. Across the upper part of this ~300 m-tall cliff exposure, pale TTG-gneiss host rocks are locally red-discoloured on both outer sides of this dyke pair, typical for how feldspars hydrothermally alter within brittle shear zones. (c) A sub-vertical dyke displays a rare along- k ‘b ’ c b q E-W trending and cross-cut meta-dyke, where it locally becomes at least 5-6 times wider than this otherwise 8 m-thick dyke. (d) GoogleTM Earth map view of the inland Umivik area, showing at least 10 parallel, relatively young and thereby likely MBDS dykes across a completely exposed ~25.5 km-wide swarm transect. (e) Detail of dyke (sample 527191), where it locally hosts the greatest concentration of some of the largest plagioclase phenocrysts observed in the field (lens cap is ~5 cm wide).

A conspicuous ‘b

’(

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h‘ w

’) along the dyke in Figure 2(c) is another unusual

feature of this swarm, which may be related to either localized thermal erosion of the host rock or, in this case, anomalously large dilation near a cross cut Palaeoproterozoic dyke. Such bulging dyke geometries are rare and rarely mentioned in the literature – let alone explained – and we had no opportunity to study the feature any closer. The same dyke was sampled farther down the cliff, along the coast, where it also exhibited the highest recorded concentration of up to 6 cm large plagioclase

phenocrysts (Fig. 2e). The presence of such phenocrysts is perhaps the single most diagnostic feature of dykes belonging to this swarm, although not mentioned for the MBDS in NW Greenland. More commonly, however, such phenocrysts are very scattered and also somewhat smaller (<3 cm), but nevertheless typically found within any outcrop of a NNW-SSE trending fresh dolerite dyke, and thereby successfully used to identify dykes belonging to this swarm segment across SE Greenland. 3.2. Skjoldungen While a few potential MBDS dykes were sampled by the first author from more inland parts of the Skjoldungen area in 2011, most dykes were in 2012 more systematically mapped and sampled along a Graah Fjord traverse. As this area is covered by low-resolution GoogleTM Earth images, it is not possible to trace dykes within the ~10 km wide section shown in Figure 3(a), where four NNW-SSE trending and sub-parallel dykes are offset ~27 m by another WSW-ENE trending dextral transverse fault. Presuming that three oblique sets of nearby smaller faults, which offset a WSW-ENE trending meta-dyke, are all related as secondary riedel shears (e.g., Dresen, 1991) to the mentioned dextral fault that runs just south of these, it is found that their sense of shear along only one NW-SE set are consistent with overall dextral displacement; whereas, the other two shear sets are more consistent with sinistral displacement (cf., inserted strain ellipses in Fig. 3a, as well as its figure caption, for details). These inconsistent riedel shear indicators may reflect that – although dykes are commonly displaced dextrally along such faults – some of these dextral displacements may have reactivated pre’ (2015) F

existing sinistral faults, as shown for the same Skjoldungen fault by

3.

A 27 m thick N-S trending dyke in Fig. 3(b) displayed some regular dyke-perpendicular aplitic veins that never continued into its host rock and resemble those reported from the 100 m thick and >400 km long MBDS dyke in NW Greenland (Kalsbeek and Taylor, 1986). Kalsbeek and Taylor (1986) offer three interpretations

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“(1) products of melting of the country rock, (2) late stage

products of the solidification of the basic magma, (3) material deposited from hydrothermal fluids circulating through joints - or products formed by a combination of these processes”. We think that these represent very late stage veining by contact-melted host rocks along h

k ’ cooling joints,

which also altered the surrounding dyke rock. A large magma throughput probably provided the excess heat to melt the Archaean host rock, but was never contaminated behind firstly an insulating chilled margin and then as the crustal melt back-veined into an already crystallized and jointed dyke interior. Regardless of their speculative genesis, however, such aplitic veins appear to be another particular feature of the MBDS, which was noted for 8 out of the 36 potential MBDS dykes mapped in SE Greenland; yet, never observed in any x

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’ other dyke swarms. These eight dyke (≥26

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consistent with thicker dykes being able to partially melt its host rock more easily than thinner dykes.

Figure 3: The Skjoldungen area. (a) Only a selected small map area (in poor GoogleTM Earth resolution) is shown, where four individual NNW-SSE trending dykes were sampled. A distinct WSW-ENE trending lineament is on the basis of a pair of ~27 m offset dykes interpreted as another coast-perpendicular dextral fault. Some secondary riedel shears are also mapped and related to inserted strain ellipses, as discussed in the text. The most westerly located MBDS dyke was sampled and studied in more detail, as exemplified by the following three field photos. (b) Looking north, along the sampled 27 m thick dyke, next to which there is a 3 m thick dyke (white arrow) that was not sampled. Person in green jacket provides a scale. (c) 1-2 cm thick aplitic vein that runs perpendicular across the dyke, and around which the host rock is discoloured. (d) Rare fine grained, c c c h k ’ c c . Th c w gether w h h k ’ c .P ~13 c .

Finally, a ~0.5 m wide, irregularly rounded, fine grained and melanocratic enclave was discovered near the centre of the 27 m thick dyke, which was sampled together with samples of h

k ’ c

and margin. As will be shown later, this was fortuitous since there is a compositional relationship between these three samples that provides further insight into magma differentiation processes that may have operated during dyke emplacement.

3.3. Timmiarmiit Field evidence of differentiation due to fractional crystallisation during dyke injection is exposed as trough-like modal layering from the ENE-margin of the thickest (62 m) recorded NNW-SSE trending k ’ trough-like

dyke in SE Greenland (Fig. 4a), dated and named here as the Snehatten dyke. Th

geometries resemble trough layering within the Skaergaard intrusions (Irvine, 1987) and may be interpreted as reflecting accumulation during channelled magma flowage, just as cross bedded contacts between different troughs also reflect some crystal mush erosion. In a neighbouring dyke (Fig. 4b), there is a different kind of margin-parallel grain size layering, which formed as symmetrical b h

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margins within many other swarms and are not believed to relate to any significant fractionation process. From both dyke margins inward, each symmetrical set of sharply bound zones appear to start with relatively coarser grain sizes that decrease exponentially inwards across a few centimetres. Apart from reflecting fluctuating changes in nucleation versus crystallization rates, it is uncertain what kind of variable conditions gave rise to this type of rhythmic zonation, but might record successive magma pulses and/or fluctuating PH2O during dyke dilation. Combined, the symmetrically zoned margins and interior trough-like rhythmic layering also bear some resemblance to some giant (<800 m thick) alkaline dykes in the Gardar Province (Upton, 1987). Farther east from the Snehatten dyke, there is a dyke with several felsic enclaves (Fig. 4c), which appear to have the same felsic compositions as aplitic veins that also cut this dyke (as illustrated in Fig. 2d), as well as alter the surrounding dyke rock in a similar fashion. In this case, however, we do not think that these enclaves were melt globules but more likely partly dissolved host rock xenoliths. In other cases, this distinction becomes more difficult and we cannot rule out that some contact melted host rock could have been incorporated as earlier melt globules, rather than just being injected later as aplitic back veins. It may be noted here that – although both felsic enclaves and aplitic veins are relatively common throughout thicker dykes, across both SE and NW Greenland – Kalsbeek and T

’ (1986) isotope results on the 100 m thick and >400 km long MBDS dyke did not record much

crustal assimilation. Thus, we have reasons to suspect that most examples of either partly dissolved xenoliths or melt globules rarely mixed but rather mingled into the dyke magma. The locations of the above three outcrops are indicated on the GoogleTM Earth map in Figure 4(d), and illustrate how much the interiors of potential MBDS dykes vary within a relatively restricted area. Like within the two previous areas (Figs. 2-3), another major WSW-ENE trending dextral transverse fault also displaces potential MBDS dykes across the southernmost Timmiarmiit area. In this case by as much as ~400 m; yet, possibly also displacing an E-W trending younger Gardar dyke (Bartels et al., 2016) by an apparent smaller amount (Fig. 4d) and thereby again suggesting that some faults may have been reactivated. F

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-25’N) c

h h-resolution GoogleTM

Earth area shown in Halls et al. (2011), we have identified another seven similar dextral transverse

faults (two of which are shown later in Fig. 11a), which in close succession displace the most southerly located potential MBDS dykes laterally by an accumulated amount of ~865 m.

Figure 4: The Timmiarmiut area. (a) Rhythmic modal trough layering on the eastern side of the 62 m thick Snehatten dyke, which was sampled for both geochemistry (541360 and 541361) and geochronology (527259). End of sledge hammer shaft in lower left corner provides a scale. (b) Rhythmic contact-parallel grain size zoning along east margin of a 21.5 m thick dyke, which is mirrored along its west margin (i.e., zoned symmetrically inwards). Hammer head is ~15 cm long. (c) Aggregates of felsic pods inside a 15.5 m thick dyke, around which the host rock is discoloured in a similar fashion as in Figure 3(c). Regular aplitic back veins also cut dykes at (a) and (c). (d) GoogleTM Earth map showing six N-S trending dykes, which cut E-W trending Palaeoproterozoic dykes as well as a late-Ketilidian appinite sill, but are cut by an E-W trending Gardar dyke (Bartels et al., 2016). A WSW-ENE trending dextral fault displaces dykes by ~400 m. Locations of (a-c) are indicated on this map.

The N-S trending dykes across the Timmiarmiit area consistently cut E-W trending Palaeoproterozoic meta-dykes (Klausen et al., 2016a), as well as appinitic sheets across the Timmiarmiit area (Klausen et al., 2016b), including a layered sill in Figure 4(d), which were emplaced before the 1.90-1.78 Ma Ketilidian orogeny (Garde et al., 2002) and during its latter stages, respectively. Whereas, older E-W trending dykes may constitute several more or less meta-doleritic to amphibolitic swarms extending across the North Atlantic craton and into the Nagssugtoqidian Orogen (e.g., Klausen et al., 2016b; Nilsson et al., this volume), the appinites are restricted to the Ketilidian Orogen (cf., Fig. 1) and are regarded as classical late-orogenic injections of hydrous magmas with highly variable and particularly Ba- and Sr rich compositions (Murphy, 2013; Klausen et al., 2016b).

4. Geochronology A coarse grained geochronological sample (527259) was collected from the so-called Snehatten dyke

(see Fig. 4d for location), together with a geochemical sample (541360) and a modally layered hand specimen (541361) from other parts of the same dyke (Fig. 4a). Separation of baddeleyite from sample 527259 was done using a Wilfley (water-shaking) table at the Department of Geology, Lund University, Sweden, largely following the procedure described in detail by Söderlund and Johansson (2002). The recovered baddeleyite consist of medium brown to honey-coloured euhedral grains and fragments, typically 50–100 μ

h c-axis. Grains of visually best quality (i.e. no

visible inclusions or zircon overgrowths) were handpicked in optical microscope and transferred to Teflon dissolution bombs, where they were washed repeatedly with HNO 3 and H2O. After spiking with a 236-233U-205Pb tracer solution, the grains were dissolved in a 10:1 HF:HNO3 mixture over 72 h at 210º C. The bombs with the dissolved grains were then put on a hot plate and evaporated, and then redissolved in 10 drops of 3.1 N HCl. A single drop of H2PO4 was then added, and the fractions were let to evaporate on a hot plate. Before loading fractions on outgassed Re-filaments, 2 μl of Si-gel was added to each fraction. U-Pb analyses were done on a Finnigan Triton TIMS at Vegacenter, Swedish Museum of Natural History, Stockholm, Sweden. Further details on sample preparation, ID-TIMS analysis, data reduction and plotting are identical to what is described in Nilsson et al. (this volume).

Figure 5: U-Pb concordia diagram of Snehatten dyke, Timmiatmiit, South-East Greenland. 207Pb/206Pb weighted mean diagram for baddeleyite fractions a,b and c is shown as an inset. Age uncertainties, error ellipses and bars 2σ.

Three near-concordant fractions gives a 207Pb/206Pb weighted average of 1629.7 ± 3.9 Ma, which we interpret as the best estimate of the emplacement age for the Snehatten dyke. Free regression yields an upper intercept of 1646 ± 41 Ma and a lower intercept of 990 ± 980 Ma. Concordia diagram is shown in Figure 5 and the U-Pb data is presented as Table 1. Thus, the age of the Snehatten dyke lies squarely within the 1622 to 1635 Ma range derived from four MB dykes, which were sampled between 900 and 1850 km farther north, along the NW Greenland segment of the MBDS (Halls et al., 2011). There is even an indistinguishable match with the 1629.4 ± 0.8 Ma age of Halls et al. (2011) derived from the ~100 m thick and >400 km long dyke, initially studied by Kalsbeek and Taylor (1986). This age correlation is the most conclusive evidence, linking the studied dyke swarm in SE Greenland to the MBDS across NW Greenland and thereby allow us now to refer to sub-parallel and similar looking dykes across SE Greenland as certain (not just potential) MBDS dykes. 5. Thickness distributions A

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not often

used to its full potential in the scientific literature. Dyke thicknesses for most MBDS dykes were routinely measured in the field and/or from high resolution GoogleTM Earth images. In Figure 6, the thicknesses of 26 different SE Greenland dykes are plotted in a logarithmic cumulative frequency versus thickness diagram, where any negative exponential distribution can be quantified by a linear best fit (e.g., Jolly and Sanderson, 1995) and the inverse exponential coefficient (IEC) of its slope defines an average thickness for such an ideal distribution (Klausen, 2006). The IEC-thickness is a b

‘

’

, not only because

w

’ dyke thicknesses are not normally distributed

around an arithmetic mean, but also because a more typical negative exponential distribution by any given swarm is often biased by an anomalous absence of thinner dykes (and thereby produce more log-normal or Weibull distributions; e.g., Krumbholz et al., 2014). This bias is simply because thinner dykes freeze quicker, before being able to propagate as far from a common source as thicker dykes; yet, would have conformed to the negative exponential distribution determined by thicker dykes if this thermal restriction had not played a role. Based on the above reasoning, Figure 6 shows that the MBDS dykes across SE Greenland define an IEC-thickness of 13 m, which might be typical for other Precambrian giant mafic dyke swarms but is more than twice that of dyke and sheet swarms measured, for example, within the Tertiary North Atlantic Igneous Province, including Iceland (cf., Klausen, 2006). Furthermore, this distribution appears to have a marked decline in the number of dykes thinner than a cut-off value of ~12.5 m, which according to the above reasoning suggests that these dykes were emplaced correspondingly far from their source. In comparison, four dyke traverses – located at progressively greater distances along a ~250 km long Tertiary coast-parallel dyke swarm segment and south of its magmatic centre above the proto-Icelandic hot spot track – display a systematic ~2 to 5 m decrease in cut off thicknesses (the three dyke traverses closest to the magmatic centre are presented in Klausen and Larsen, 2002). Compared to the MBDS across NW Greenland, however, a thickness data set on only 13 MBDS

dykes, mainly recorded by Halls et al. (2011) and double checked on high resolution GoogleTM Earth imagery, provides an IEC-thickness of 43.5 m and a cut off thickness at 45 m, below which only two thinner dykes have been measured. Even if there may be other thinner MBDS dykes across NW Greenland, which have not been recorded in the field or cannot be resolved on GoogleTM Earth, we do not think there are enough misses to minimize the large cut-off thickness and thereby devalue its later use to argue for northward lateral dyke propagation of the MBDS across Greenland

Figure 6: Cumulative frequency versus thickness distribution of MBDS dykes across SE (orange triangles) and NW (yellow triangles) Greenland. Fitted straight segments reflect negative exponential thickness variations, the slope of which (i.e., inverse exponential coefficient) corresponds to an IEC-thickness that represents the average of that distribution. A cut off thickness, below which there is an under-representation of thinner dykes, can be tentatively related to the distance from a central magma source where only thicker dykes are able to propagate before freezing. Conversely, some anomalously thick dykes deviate from the negative exponential distribution and could represent multiple, or in any other way unusually thick, dykes. Note that Halls et al (2011) describe MB2 and OF as a variably 70-100 m and 200 m thick dyke, respectively, where we use different thickness TM estimates derived from Google Earth. See Klausen (2006) for details on this type of plot, which clearly shows that MBDS dykes across SE Greenland (SEG) are thinner and have a smaller cut off thickness than MBDS dykes across NW Greenland (NWG).

Finally, it is striking how much thicker the MBDS dykes are in NW Greenland, compared to the thickest dyke measured in SE Greenland. Judging from how E

’ 62 m thick Snehatten

dyke and three >100 m thick dykes in NW Greenland do not conform to downward extrapolations of their respective negative exponential distribution (Fig. 6), there is reason to suspect that these are all anomalously thick. Without supporting field evidence, however, one can only speculate on whether such anomalous thicknesses are the results of (1) multiple intrusions, erroneously measured as single dykes, (2) dyke width irregularities, such as keel-shaped giant dykes or local ‘b

s’ (e.g., Fig. 2b),

(3) dyke dilations under different conditions (e.g., from a different source and/or inside host rocks with different elastic properties), or (4) measuring different parts of laterally propagating blade-shaped dykes with ellipsoidal cross-sections, as will be discussed later.

6. Petrography All MBDS dykes across SE Greenland are more or less olivine-bearing dolerites to gabbros, similar to those described for dykes across NW Greenland, and consist of mainly elongated euhedral plagioclase, eu-subhedral olivine, interstitial beige augite as well as sub-anhedral opaque minerals and accessory apatite. The rocks are generally fresh but may locally – as observed along thin alteration veins – have serpentinized olivines, augites replaced by hornblende and more skeletal opaques. We did not note much alkali feldspar or quartz in our thin sections, as was recorded in more evolved samples by Kalsbeek and Taylor (1986). As most of the above petrography is repetitive of what has been reported on the NW Greenland MBDS, we focus on

h

w

’ ch

c

c

c

phenocrysts,

which may be more common in SE Greenland, but has also been reported for NE Greenland (Nilsson and Hamilton, 2017). Kalsbeek and Taylor (1986) only describe <1 mm-large micophenocrysts within chilled margins and <1 cm-large crystals within thick dyke centres that are not specified as phenocrysts.

Figure 7: Petrography of sample 541367. (a) Scanned thin section showing an overabundance of elongated euhedral plagioclases (including a larger phenocryst) and roughly equal proportions of olivine (fractured and with a higher relief) and interstitial beige augites, together with some opaque oxides. (b) Photo micrographic image of thin section area outlined by red rectangle in (a), where its left and right halves were photographed under plane and crossed polarized light, respectively. Typical augites (au) and olivines (ol) are indicated, amongst plagioclase laths and opaque oxides. A rounded contact between the plagioclase phenocryst core and its rim overgrowth is traced by a dashed red line. Note how the olivine above the green arrow has grown around smaller plagioclase laths; yet, is itself hosted within the plagioclase rim, suggesting early cotectic growth of these two minerals. According to CIPW normative calculations, this sample contains ~61 vol % plagioclase, ~16 vol % olivine ~8 vol % augite and ~3 vol % oxides, together with almost 10 vol % orthoclase.

One thin section exhibits a relatively euhedral and lath-shaped phenocryst (Fig. 7a), which under the microscope reveals an older rounded core with albite twin extinction angles of ~40˚ (Fig. 7b). This reflects a phenocryst core composition of ~An70, which is slightly more calcic than coexisting matrix plagioclases of
’ C

A

c F c

, as specified in greater detail by Klausen et al.

(2017). Thus, samples were first cleaned and fresh blocks cut for steel jaw crushing. A small handful of quarter-and-cone split crushed sample material was milled in a carbon-steel swing mill, intermittent with clean quartz. Powders were fused into La-free glass beads for both (1) major element X-Ray F

c c (XRF)

Ph

’ PW1404w

(Ax

PAN

c w h

2.4 kWatt Rh X-ray Tube), and (2) trace element analysis, using an Agilent 7500ce ICP-MS coupled with a Nd-YAG 223 nm New Wave LASER ablation (LA) system operating at a 12 Hz frequency with a mixed He-Ar carrier gas (following method by Eggins, 2003). A wide range of international (NIST®) and national (SARM®) standards (Pearce et al., 1996) were analysed concurrently with the samples. All 34 sample analyses are presented as supplementary Table 2. Following Kalsbeek and Taylor (1986), a similar normative ne-ol-di-hy-q diagram is shown in Figure 8(a), where the MBDS dyke ’

low di conform to a ~9 kbar curve (Thompson, 1982). We

agree that this argues for these magmas to have equilibrated, during differentiation, at depths corresponding to the lowermost part, or even the base, of a nearly 30 km thick Greenlandic continental crust. Otherwise – as far as a limited number of elements on samples from NW Greenland allows – normative mineral compositions and other geochemical similarities further link the two swarm segments into a coherent trans-Greenlandic dyke swarm. Thus, like NW Greenland (Kalsbeek and Taylor, 1986; Nielsen, 1990), samples from SE Greenland are equally transitional to weakly trachytic basalts (Fig. 8b) and most plot above M

h

’ (1978) curved alkaline-subalkaline divide. Only

three more andesitic and two silica-poorer trachy-basalts plot outside an otherwise tight cluster, which defines a weak positive trend between 48-51 wt% SiO2.

Figure 8: (a) Composite ternary diagrams displaying weight percentages of CIPW normative mineral proportions, simply calculated according to Irvine and Baragar (1971). One atmosphere and ~9 kbar curves are copied from Thompson (1982). (b) Total alkali versus silica (TAS) classification diagram, where all samples have been normalised to 100% volatile-free major element proportions and the grey field outlines transitional basalts (according to Le Maitre 2002). Mi h ’ (1978) division between more alkali and sub-alkali basalts is added as a red curve. White and yellow diamonds represent chilled margin and centres, respectively, of Kalsbeek T ’ (1986) sampled MBDS dykes from NW Greenland (NWG), except for the one sample from Artíng (2004) emphasised with an asterisk. Other symbols are all results from this paper.

Selected major and trace elements, plotted against a slightly wider 4-8 wt % MgO range (Fig. 9) also, expectedly, follow trends defined by NW Greenland dykes (Kalsbeek and Taylor, 1986; Nielsen, 1990). The four most compatible elements define more or less obvious positive trends for Al 2O3, CaO, Sr and Ni (Fig. 9a-d ), when plotted against MgO. Such trends are for the first three elements all consistent with plagioclase fractionation; whereas, the distinct decline of Ni against declining MgO are both consistent with additional olivine fractionation. Reversed extrapolations of trends in Figure 10(ab), provide some simple olivine:plagioclase estimates that conform to Kalsbeek and T

’ (1986)

mentioned 1:3 phenocryst proportions, preserved inside chilled margins. Thus, there is a perfect agreement between field observations, petrography and bulk rock geochemical variations, suggesting that such cotectic proportions of ~75% plagioclase and ~25% olivine fractionation (± some in situ cc

)

h MB

’

c

.

Figure (9e-h) displays negative trends for selected incompatible elements, where TiO2 reflects little to no (Fe,Ti)-oxides fractionation, and thereby crystallization under relatively low P H2O conditions. K2O, a large ionic lithophile element (LILE), displays a notably distinct negative trend, showing that these fresh rocks experienced very little alteration. Its trend is even more distinct than those for less mobile high field strength elements (HFSEs), such as Zr and Nb (Fig. 9g-h), the relative enrichments of which c

h <50%

h

w

’

c

h

c

c c

c

+

olivine assemblage, in order to produce its most evolved magmas. The published data for NW Greenland samples (Kalsbeek and Taylor, 1986; Nielsen, 1990)

’ XRF b

y,

show some systematically higher Nb-values, which most likely relates to some systematic analytical

error. Apart from Nb, however, all available elements from both SE Greenland and NW Greenland consistently match each other.

Figure 9: Eight selected variation diagrams with MgO as a common differentiation index for mafic igneous rocks, from right to left along the x-axis. Left and right columns show more compatible and incompatible elements, respectively. Same samples and symbols as in Figure 8. Manually fitted trend lines are indicated, as well as a pair of rough estimates of relative fractions (F) of enrichment.

The ICPMS-data of this paper presents the first full compositional spectrum of incompatible elements for a large proportion of the MBDS (at least 36 dykes as mapped in Fig. 1b). The apparent geochemical homogeneity amongst MBDS-samples across both NW- and SE Greenland is perhaps best illustrated by multi-element patterns in so-called spider diagrams, where we find very consistent overlaps between 27 out of 31 SE Greenland samples (Fig 10a). Their HFSE-patterns are slightly more enriched than a present day E-MORB, yet less enriched than an OIB, and are characterized by very distinctly positive Pb and negative Nb-Ta spikes. Together with higher LILE/HFSE ratios, these are c ‘arc’ (

‘c

h

h

’) signatures, where the MBDS – as noted by Nielsen (1990) –

also stands out by having particularly high Ba (typically between 600-1000 ppm). These patterns, furthermore, resemble seven patterns of ~100 Myr older late-Ketilidian spessartites (i.e., appinite subgroup) from the Timmiarmiit area, which just have overall slightly lower concentrations, as well as (1) higher Sr due to the fractionation of hornblende rather than plagioclase, and (2) distinct negative Tianomalies reflecting an earlier onset of oxide fractionation from these more hydrous magmas (Klausen et al., 2016b). Fortunately, one sample from Artíng (2014) – collected from the same dyke and place as H

’

(2011) MB2 sample of the ~100 m-thick and 400 km-long MBDS dyke in W Greenland – is analysed for the same suite of incompatible elements shown in Figure 10, and is near-perfectly matched with samples 541360 and 541361 from the thickest Snehatten MBDS-dyke across SE Greenland (cf., Fig. 4a,d for location). This further strengthens a direct correlation between these two dyke segments b

h

’

c c

wh ch

h

ch

c

b

indistinguishable ages of 1629.4 ± 0.8 Ma (Halls et al., 2011) and 1629.7 ± 3.9 Ma (this study). It also installs greater confidence in the precisions and accuracies of both the geochemical data and geochronological results presented in this paper. The two most primitive SE Greenland MB-samples, collected from the amygdaloidal centre (541333c) and margin (541334m) of an only 1.2 m-thick and N-S trending dyke, exhibit more subdued patterns Figure 10(b) that conform quite well to the bulk of the MBDS (e.g., almost equally high Ba) except for much lower concentrations of Cs, Rb, Th and U. Following upon Kalsbeek and T

’ (1986)

display of greater compositional variation across than along a ~100 m-thick and 400 km-long MBDS dyke segment in NW Greenland, it could be argued that this anomalous pair represents a less differentiated parental magma, which was injected and crystallised ahead of the bulk of the more differentiated magma, plotted in Figure 10(a). However, this would also suggest that the bulk of the MBDS was selectively enriched in only Cs, Rb, Th and U, which seems unrealistic. Alternatively, a more conservative view would be to regard this lone dyke, with unusual amygdales and no diagnostic plagioclase phenocrysts, as belonging to another swarm. Because this thin Skjoldungen dyke is relatively fresh and located most proximal to the Atlantic Ocean (cf., Fig. 1b), it is tempting to classify it as Tertiary, even if it is not as coast-parallel and does not share the same OIB-signatures of more

northerly located

T

k

(c . ‘

k’ w

/

k

Figs. 1b and 2d, respectively;

Weatherly et al., 2016).

Figure 10: Incompatible element diagrams for rocks normalised according to Sun and McDonough (1989) to a typical E-MORB. (a) 32 MBDS dyke centre samples from across SE Greenland, which all display remarkably parallel patterns that only exhibit a narrow concentration range that is tentatively attributed to relative enrichments from <50% differentiation. This range is used as an orange background reference field to all subsequent plots, below. Range of seven late-Ketilidian spessartites are shown for comparison as a green field. (b) Solid black lines are two samples from the MBDS in W Greenland (Artíng, 2004). Dashed black lines are two samples from the amygdaloidal centre and margin of a 1.2 m thick dyke, which are the most primitive parents in the data set. See text for discussion. (c) Three samples from the centre, margin and a fine grained mafic enclave inside the same dyke (cf., Figure 3 for details), exhibiting progressively lower (more primitive) element patters.

If the above potential parental pair from the thin amygdaloidal Skjoldungen dyke is not part of the MBDS, then the more geochemically similar fine-grained enclave (541317), hosted inside a 27 mthick and plagioclase phenocrystric dyke (Fig. 3), represent our most primitive parental magma for the MBDS. In Figure 11(c), this enclave is plotted together with samples from h h

k ’

(541318) and centre (541319), where one should note that all three samples exhibit roughly parallel patterns; except for Sr that is most likely buffered by fractionating plagioclase. The geochemical signature of this dyke centre is indistinguishable from the consistent range of the main MBDS (orange background), also primarily based on sampled dyke centres. Thus, the progressively higher incompatible element patterns

h

k ’ fine-grained enclave, margin and centre, respectively, are

consistent with successive crystallisation of more differentiated magmas during dyke emplacement. Especially, if the irregular fine-grained enclave (Fig. 3e) is a chilled margin autolith that crystallised before the sampled dyke margin. In other words, these three samples are consistent with a normal compositional section across a dyke, which contradict many of K

b k

T

’ (1986) less

systematic compositional sections across their >400 km long and >100 m thick MBDS dyke in NW Greenland. Following upon interpretations by Chistyakova and Latypov, (2012), a normal zonation reflects progressively more differentiated magmas feeding into the dyke and/or the magma progressively becoming more differentiated due to in situ conduit fractional crystallization during dyke emplacement. 8. Model implications This section discusses implications of evidence presented by the MBDS – through results published by Kalsbeek and Taylor, (1986), Nielsen (1990), Halls et al. (2011) and now this paper – by first reviewing evidence for its correlation into a >2000 km long dyke swarm that cuts across the entire large island of Greenland. We thereafter examine h

w

’ infrastructure, with special emphasis on

dyke thickness variations, before arguing for lateral dyke propagation from a lower crustal magma source within the Ketilidian orogen. This magma source must have been a single relatively large and homogenised chamber, where magmas – over a remarkably long time period – were buffered within a narrow compositional range during cotectic fractionation of plagioclase and olivine. We then speculate on how h ‘ c’-like parental magmas for this chamber were generated within the post-Ketilidian orogenic belt, either through wholesale crustal assimilation of more asthenospheric melts and/or through direct partial melting of a sub-continental lithospheric mantle (SCLM). Two opposing endmember configurations of the Nuna supercontinent are reviewed, which provide distinctly different tectonic settings for the swarm and its associated magma chamber. Regardless of its uncertain tectonic setting, we proceed with speculations on what role such a magma chamber (and its associated MBDS) played within a unique Mesoproterozoic AMCG-province, made up of coeval massif anorthosites and exotic felsic intrusions, including rapakivi granites.

Figure 11 (previous page): Infrastructure of the MBDS. (a) Detail of most southerly located MBDS dykes, remotely identified as #16 and #17 in H ’ (2011) Figure 27.11, cut by a likely Tertiary dyke (aka #28) and all displaced by two SSW-NNE trending dextral faults. (b) High resolution GoogleTM Earth imagery of the irregular northernmost extent of the >400 km long and >100 m thick MB dyke in NW Greenland (studied by Kalsbeek and Taylor, 1986), as well as a ~45 m thick dyke (MB11) east of it. See Figure 1(a) and (c) for location and note that exposed segments of the irregular north-end are sinistrally offset. (c) Updated map view of MBDS outcrops, modified from Figure 1, Nielsen (1990), Halls et al. (2011), as well as further GoogleTM Earth surveys. Note perfect dashed grey correlation between similar aged dyke segments, beneath the inland ice sheet. (d) Zoomed in cross section view of the southernmost ~550 km of the MBDS, depicting (1) a hypothetical lower crustal magma chamber (grey) below a currently ~40 km thick continental crust; (2) three overlapping schematic examples of laterally propagating blade-shaped dykes (orange); as well as (3) how massif anorthosites and rapakivi granites may have formed above the mafic magma chamber, over a period of at least 13 ±6 myrs.

8.1. Correlation across Greenland Based on GoogleTM Earth mapping, Halls et al. (2011) initially proposed an extrapolation of the MBDS – as exposed along the NW coast of Greenland – beneath an inland ice sheet, and >2000 km across to the SE coast of Greenland. This correlation (illustrated in Fig. 11) is now confirmed by similar ages, as well as both petrographic and geochemical similarities between dykes across both NW and SE Greenland, presented above. As (1) our 1629.7 ± 3.9 Ma baddeleyite U-Pb age(Fig. 5) is indistinguishable from H

.’ (2011) 1629.4 ± 0.8 Ma age, (2) both ages were derived from the

thickest dyke on either side of the Greenlandic ice sheet, with (3) coinciding trends and similar locations ~50-75 km to the east of the inferred western margin of the MBDS (cf., Fig. 11c), and (4) bulk rock geochemical analyses from both dyke segments provide a perfectly match (Fig. 10b; Artíng, 2004), our dated Snehatten dyke in SE Greenland is likely part of same up to >100 m-thick and now >1300 km long dyke, initially studied by Kalsbeek and Taylor (1986). Such a near perfect geochronological, structural and geochemical correlation not only adds further credence to the precision and accuracy of these methods, but more importantly accentuates the remarkable compositional homogeneity along individual MBDS dykes. As shown by Figure 1(a), the MBDS cuts diagonally across the entire extent of the Greenlandic island, and for most part at a relatively high angle across primarily more E-W trending older orogens and subparallel internal fabrics, as well as pre-existing and pre-orogenic Palaeoproterozoic meta-dyke swarms (not shown; e.g., Nilsson et al., 2013; this volume). Thus, there are no field evidence of the swarm following any pre-exiting lithological fabrics or tectonic structures, just as no associated, coinciding faulted rift structure has ever been recognised. The continuous N-S to NNW-SSE trending pattern of the MBDS, furthermore, suggest that it is largely unaffected by any subsequent deformations. It may be dextrally offset by as much as 200 km along a Nares Strait transverse fault zone (e.g., Denyszyn et al., 2009), but that would according to Halls et al. (2011) only have affected the northernmost tapering end of the swarm. Otherwise, our mapping of WSW-ENE trending transverse faults – apparently cutting across SE Greenland at a regular spacing of ~40-60 km (Figs. 2-4 and more southern locations in Fig. 11a) and most likely formed by a northward propagating proto-Atlantic rift (e.g., Srivastava and Tapscott, 1986) – only constitutes a cumulative ~1363 m of dextral displacement of the MBDS across SE Greenland.

8.2. Lateral dyke propagation from a southerly located magma source We agree that it is difficult to envisage giant mafic dykes propagating laterally for >1000 km through only ~40 km thin continental plates (e.g., Nielsen, 1990) and understand why Kalsbeek and Taylor (1986) consequently opted for a more homogenous asthenospheric, rather than the underlying more heterogeneous lithospheric, mantle source to explain the compositional homogeneity along most of the MBDS. However, lateral dyke propagation has since been advocated for many other giant dyke swarms, starting with Ernst and Baragar’ (1992) anisotropy of magnetic susceptibility (AMS) constraints on the 1500 km long and slightly radiating Mackenzie dyke swarm. As supported by Halls et al.’ (2011) palaeomagnetically constrained plate reconstruction (shown later in Fig. 13b) a >2000 km long trans-Greenlandic MBDS may well extrapolate to a magmatic centre within a coeval Fennoscandian rapakivi province, from which dykes must then have been laterally emplaced. As explained in Section 5, a larger cut-off thickness (i.e., the thickness below which there are anomalously few thinner dykes; Fig, 8) for the NW Greenland segment of the MBDS, compared to SE Greenland , is consistent with thicker dykes propagating farther northwards than thinner dykes before freezing. We think these cut-off thicknesses are representative for SE Greenland, where large parts of the MBDS were measured in the field, as well as NW Greenland, where only 13 individual dykes, out of N

’ (1990) initial 33 dyke segments (28 in Fig. 1a), are verified by both H

and subsequent GoogleTM Earth mapping (cf., Fig. 11c). Some sub-

‘

c

’ (2011) ’

h

western Disko Bay margin of the MBDS, could host up to a few meter-thick unconfirmed dykes, but are more likely part of an overlapping and more coast parallel swarm (or tectonic structures) related to Tertiary rifting along the offshore volcanic rifted margin of the Labrador Sea (e.g., Larsen et al., 2009). Thus, we are confident that the two cut-off thicknesses in Figure 6 are representative and the product of a northward propagating dyke swarm. However, such infrastructure (e.g., Gudmundsson, 1995) would have a greater accumulated dyke thickness closer to its source, made up of an additional number of thinner dykes that did not propagate as far as thicker dykes, as opposed to the accumulated dyke thicknesses across NW Greenland (>831 m) being double that of SE Greenland (>424 m). We do not think the MBDS was dextrally offset along its NW Greenland segment (cf., subdued displacement arrows in Fig. 11c) – even if that would reduce our summation of dyke thicknesses to >236 m and >595 m across the southern and northern segment of NW Greenland, respectively – because the northernmost irregular part of the >400 km long and >100 m thick dyke in Figure 11(b) indicates an opposite (sinistral) sense of displacement. On the basis of the southward decrease in accumulated dyke thicknesses, one could argue for southward dyke propagation from a northern source, but then the swarm must be made up of individual dykes that bifurcate towards the south, in order to explain the greater number of thinner dykes across SE Greenland. Such branching is mechanically infeasible, however, because a relatively limited magma overpressure inside a single ellipsoidal dyke needs to overcome the tensile strength along all h

k ’ leading edges (e.g., Rubin, 1995; Rivalta et al., 2015); i.e., any bifurcation would

double the energy required for the magma to fracture its host rock to propagate. Thus, any observed bifurcation typically relates to local magmatic offsets, where one of the ‘b

ch ’

s as a

relatively short apophyse. It might be mechanically feasible for two separate, yet coeval, dykes to merge in an opposite propagating direction into a single thicker dyke, but – while this could be reconciled with an observed greater number of thinner dykes closer to a source (i.e., SE Greenland) and fewer yet significantly thicker merged dykes farther from the source (i.e., NW Greenland) – it does not explain a northward doubling of accumulated dyke thicknesses. This dyke merging hypothesis is also untested and requires an unrealistic concurrent injections of several dykes from a magma source. Thus, our preferred explanation

h MB

’

c

relates to a dyke’

typically ellipsoidal cross sectional geometry (Rubin, 1995; Rivalta et al., 2015), where erosional levels across SE Greenland exposes thinner dyke edges and NW Greenland exposes thicker dyke interiors. This can be reconciled with northward propagating giant bladed dykes, as schematically illustrated in Figure 11(c), where upper edges of bladed dykes are shown to rise higher into the crust at increasing distance from their sub-crustal magma chamber source. Consequently, the current erosional level tend to expose thinner dyke edges closer to the source and thicker dyke interiors farther from the source. An illustrative example is provided by the measured ~62 m thick ‘edge’(i.e., the Snehatten dyke) and the >100 m-thick ‘interior’ (i.e., the dyke investigated by Kaalsbek and Taylor, 1986) of the same >1300 km-long dyke across SE and NW Greenland, respectively, as correlated in Section 8.1. In conclusion, an overall northward dyke propagation from a southern magma source is supported by (1) Halls et al.’ (2011) observations of dykes thinning (i.e., tapering)

w

h

w

’

northernmost outcrops; (2) a greater number of thinner dykes in SE Greenland, as opposed to fewer thicker dykes in NW Greenland (Fig, 6); with (3) correspondingly smaller and larger cut-off thickness, respectively; but where (4) erosional levels in SE Greenland intersect fewer and/or thinner upper parts of bladed dykes with ellipsoidal cross sections, as tentatively illustrated in Figure 11(d). Such a dyke swarm infrastructure requires the existence of a more southerly located magma reservoir, discussed next. 8.3. A single long-lived yet compositionally homogeneous lower crustal magma chamber source As noted by Kalsbeek and Taylor (1986), the relatively evolved (MgO between 7-4 wt %) compositional homogeneity of the MBDS requires a single differentiated magma chamber source, which on the basis of Figure 8(a) was located at lower crustal depths. This compositional homogeneity is supported by remarkably parallel spider patterns in Figure 10(a), which is only shifted on the basis of <50% differentiation due to a bulk fractionation of 75% plagioclase and 25% olivine (Fig. 9). In conclusion, all of the dykes within the MBDS were injected from the same, more or less differentiated, magma chamber, which must have been voluminous enough to sustain the emplacement of individual ellipsoidal dykes that could be up to 2000 km long, presumably no higher than the ~30 km-thick crust, yet up to >100 m-thick ( . . V = 4/3π × 1000 × 15 × 0.05 = 3142 k

3

). The magma must also have

highly homogenised throughout most of the chamber (e.g., turbulently mixed) as indicated by

compositional homogeneity along the largest >1300 km long and >100 m thick dyke, correlated across both SE Greenland and NW Greenland (Figs. 8-10; Artíng, 2014; Kalsbeek and Taylor, 1986). More extraordinarily, the compositional similarity between most dykes within the MBDS suggests that the same, single magma chamber must have remained molten, compositionally homogenous, and just slightly more or less differentiated for a period of >13 ± 6 Myr, as constrained by H

.’ (2011)

ages on just four of its dykes. This may be compared to the<1 Myr duration it took the >65,000 km3large but shallow crustal Bushveld igneous complex to crystallize (Zeh et al., 2015). Even if a greater longevity is more sustainable near the hot base of a continental crust (e.g., Teng and Santosh, 2015), it is still remarkable how such a single magma chamber could have been replenished by compositionally similar parental magmas and remained on its plagioclase-olivine cotectic during such an extended period of time. One explanation would be that the magma chamber was buffered by potential temperatures near that of a dry and cotectic plagioclase-olivine mineral assemblage. However, this would also be well above even a dry solidus of the surrounding felsic host rock, where Kalsbeek and Taylor’ (1986) relatively unradiogenic initial (87Sr/86Sr)1630 ratios of 0.70144-0.70475, both along and across its >400 km long, >100 m thick and 1629.4 ± 0.8 Ma dyke (compared to 0.71565 for a cross cutting aplitic back-vein), argue against any significant assimilation of older crust; at least after half of the 1635 ± 3 to 1622 ± 3 Ma MBDS had been emplaced. As geochemical similarities (Fig. 10a), furthermore, argue for all other dykes to be equally unradiogenic (although that still remains to be properly tested), it seems unlikely that the magma chamber assimilated much old continental rocks. However, it could still have assimilated relatively young crust, if the deep magma chamber was location within the 1.90-1.78 Ga Ketilidian orogeny, south of the Archaean North Atlantic Craton. 8.4. Primary melts from a metasomatised sub-continental mantle source Incompatible element patterns and/or ratios are often used to discriminate between different tectonic settings (e.g., Pearce, 2008, and references therein), as well as the kind(s) of mantle source(s) primary magmas were generated from, their degree and depth of partial melting, as well as subsequent differentiation through typically fractional crystallization ± crustal assimilation. As described in Section 6 and shown in Figure 10, HFSEs are relatively enriched, as opposed to a more depleted mantle signature indicated by K

b k

T

’ (1986) Pb-isotopes (206Pb/204Pb, 207Pb/204Pb and

208

Pb/204Pb of 17.483-18.524, 15.461-15.525 and 36.584-38.161, respectively). However, these ratios

have yet to be properly age corrected due to the lack of corresponding 238U/204Pb, 235U/204Pb and 232

Th/204Pb analyses. ‘ c’

More significantly, a c

b

h MB

’

superimposed on the HFSEs, as most prominently h h LILEs and particularly low Nb (cf., Fig. 10). This is perhaps

best illustrated in Figure 12(a-b), where Nb is normalized by Yb and 1/La, respectively, to reduce the effect of crystal fractionation. In Figure 12(a), the MBDS plots relatively far above that mantle array, with a slightly more elevated ‘arc’

h

h

’

P aeoproterozoic meta-dyke

swarms across the North Atlantic Craton and partly overlapping the compositionally similar late- to

post-Ketilidian spessartites. Although it can be difficult, within continental settings, to determine if ch ‘ c’-like igneous rocks were derived directly from an active mantle wedge, a metasomatised SCLM, or through the crustal assimilation of asthenospheric melts, the following discussion provides some clues for the MBDS.

Figure 12: Incompatible trace element ratio plots. (b) Nb/Yb versus Th/Yb diagram (Pearce, 2008), with an mantle array, its various asthenospheric members ranging from a normal mid-oceanic ridge basalt (N-MORB), primordial mantle (PM), more enriched E-MORB, to the most enriched oceanic island basalt (OIB), and how compositions of mantle-derived melts typically change through combined crustal assimilation and fractional crystallisation (AFC). (a) La/Nb versus La/Ba plot (Jourdan et al., 2007) with fields for various asthenospheric mantle components, including a high U/Pb (HIMU) basalt. MBDS plots with a strong lithospheric signature, where older data from NW Greenland may deviate because of systematically higher Nb due to an analytical error (cf., Fig. 9h). Data ranges for (1) the Soumenniemi dyke swarm (Rämö, 1991); (2) late- to post-Ketilidian spessartites (Klausen et al., 2016b); (3) ~2.2 Ga meta dykes across SE Greenland (Nilsson et al., this volume); (4) the ~2.0 Ga Kangâmiut dyke swarm (Mayborn and Lesher, 2006); (5) older dated meta-dykes across SW Greenland (Nilsson et al., 2013); and (6) single analysis from the ~1.60-1.54 Ga Breven-Hällefors dyke swarm (Risku-Norja, 1992), shown for comparisons. SEG = SE Greenland; NWG = NW Greenland.

Firstly, the ‘

’

c

c

Figure 7

‘ c’-setting. Secondly, the

MBDS formed ~150 Myr after the ~1.90-1.78 Ga Ketlidian orogeny (Garde et al., 2002) and did therefore not source a mantle wedge above a dehydrating and subducting oceanic slab, preceding that orogeny. So, how did the MBDS acquire its exceptionally Ba-rich signatures (already noted by Nielsen, 1990) and thereby stand out

‘

h

h

c’ and so much

‘ h

h

c’ in

Figure 12(b) than any other Palaeoproterozoic dyke swarms across the North Atlantic Craton? We think the answer lies in the MBDS’ compositional matching with the >100 Myr older spessartites (Fig. 10), which could both have been derived from the same relatively unradiogenic source area within the relatively young Ketilidian orogen. Such a common mantle source is not only supported by their coinciding source locations but also by equally unradiogenic mafic and felsic rocks within the bimodal Ketilidian rapakivi granite suite, with (87Sr/86Sr)i and ɛNd values, ranging from CHUR and towards a depleted mantle (Brown et al., 2003). It is deemed more unlikely for such similar geochemical signatures to have been derived through a more complex crustal assimilation during fractional crystallisation (i.e., AFC) process during two different magmatic events. Instead, their incompatible element signatures were more likely acquired during partial melting of a similarly metasomatised Ketilidian lithospheric mantle. This, despite the fact that h

’

b q tous hornblendes

(Murphy, 2013),

h MB

’

b

(Fig. 7), implies that this SCLM

source must have dehydrated and subsequently become heated to a much greater degree in order to also generate the larger volumes of magma for the MBDS. In conclusion, we think that the deep crustal magma chamber source for the MBDS was replenished by primary melts from a SCLM source that had been extensively metasomatized (and maybe even formed) during the Ketilidian orogeny. We cannot exclude the possibility of wholesale assimilation of surrounding Ketilidian-aged crust but deem this less likely to have sustained a similar compositional homogeneity as recorded by the MBDS over a period of >13 ± 6 Myr. 8.5. Tectonic setting within an enigmatic Nuna supercontinent The Mesoproterozoic age of the MBDS suggests that it was emplaced within a Nuna (Columbia) supercontinent, which is generally accepted to have amalgamated around 1.75 Ga, even though Pisarevsky et al. (2014) also argue for a final merger of an East and West Nuna as late as ~1.65 Ga. For the sake of our focus on West Nuna, the earlier merger is in line with many coeval orogenies across the Laurentia, Baltic and Amazonian blocks, including the Nagssugtoqidian and Ketilidan orogenies of Greenland (e.g., Zhao et al., 2002). There are more uncertainties, however, concerning the Mesozoic configuration of these craton blocks (e.g., Pisarevsky et al., 2014, and reviewed reconstructions therein), where Johansson (2014) emphasizes two main options that are either geologically- or palaeomagnetically-based. A geologically based reconstruction of eastern Laurentia (including Greenland) and Baltica is shown in Figure 13(a), which brings the southern extension of the MBDS ont

c

c

120˚ b

along the joined margins of these two craton blocks. If so, the MBDS could have formed along a failed rift (e.g., Burke and Dewey, 1973) emanating from a LIP-centre that arguably formed above a mantle plume (e.g., Ernst, 2014). In fact, roughly coeval mafic dyke swarms across the Baltic Shield’ Fennoscandian province – where the Soumenniemi dyke swarm (Rämö, 1991) is coeval with the MBDS – also conform to such a radiating pattern around a common centre shown by the star in Figure 13(a). In such a plate configuration, however, you would also expect another continental plate to have occupied the conjugate vacancy, where we have arbitrarily positioned Amazonia in subdued colours (Fig. 13a) and without much supporting evidence. Regardless, the model is consistent with Halls et .’ (2011) palaeomagnetically-based ~1.6 Ga reconstruction in Figure 13(b), which could roughly represent a drift-stage, subsequent to breakup in Figure 13(a). Such a breakup model for the Nuna supercontinent is adequately explained by Windley (1993), and where the plume helps explain any anomalously high mafic magma production, as well as c

‘

’c

b

surrounding mangerites, charnockites and rapakivi granites, as well as the MBDS. Most of the plutons that are coeval with the MBDS abound within the Fennoscandian province (e.g., Rämö and Haapala, 1995), however, and it is to those that H

.’ (2011) also linked the MBDS in Figure 13(b). On

the other hand, the coeval Mealy Mountains AMCG-complex (Ashwal, 2010) widens that particular early magmatic stage in Figure 13(a), across the MBDS and into Laurentia.

Figure 13 (previous page): Different tectonic settings within the Nuna supercontinent. (a) A geologically-based reconstruction, partly following Johansson (2014). Amazonia is arbitrarily used as a conjugate block within a hypothetical triple rift scenario, centred on the red star and where hypothetical successful rifts formed along the thick dashed lines. (b) The palaeomagnetically-constrained reconstruction by Halls et al (2011) at 1.6 Ga, which may tentatively be viewed as a drifting stage following upon (a). (c) Another, more widely accepted, palaeomagnetically-constrained reconstruction, mainly following Evans and Mitchell (2011) at ~1.35 Ga. (d) Age compilation of orogens (e.g., Zhao et al., 2002), post-orogenic high-K suites (Andersson et al., 2006; 2007b; Rutanen et al., 2011), mafic dyke swarms (Ernst et al., 2008; Söderlund et al. 2005; Bartels et al., 2016) and ACMG-suites (Emslie and Hunt, 1990; Windsley, 1993; Hamilton et al., 2004; Andersson et al., 2007a; Ashwal, 2010; McLelland et al., 2010; Teixeira et al., 2013).

Alternatively, many researchers on Proterozoic rapakivi granites and associated massif anorthosites also speculate on what appears to be more like a (post-)orogenic setting (e.g., Fig. 13c), which – compared to a short-lived LIP (e.g., Bryan and Ernst, 2008) – is more consistent with the prolonged and repetitive emplacement of AMCG-complexes over a >500 Myr period. This would be in line with Ashwal and Bybee’ (2017) preferred massif anorthosite model, where plagioclases crystallised and accumulated slowly from basaltic magma chambers that formed along subduction zones, and where crystal mushes were subsequently emplaced during intra-arc rift stages. In case of the four major AMCG-events identified under East Laurentia in Figure 13(d) (three events in Emslie and Hunt, 1990), this is more consistent with an equally long-lived (>500 Myr) active margin along that part of a (post-)Nuna supercontinent, as opposed to a passive margin according to Windley (1993), x

c

‘arc-

’ cycles, possibly during periods of normal subduction and slab roll-

back, respectively. Regardless of whether or not the Nuna supercontinent transitioned into the Rodinia supercontinent as a Mesoproterozoic passive or active margin, it eventually evolved into a continuous Grenvillian to Sveconorwegian orogen (e.g., Slagstad et al., 2017), which for most parts seem to have terminated the AMCG-era (excluding the post-orogenic 0.96-0.93 Ga Rogaland anorthosite complex). Within this uncertain plate tectonic evolution, it should be noted (Fig. 13d) that margin-perpendicular mafic dyke swarms, like the MBDS, mainly formed during the initial stages of the >500 Myr-long cyclic emplacement of the Mesoproterozoic AMCG-province. From 1.30 to 1.25 Ga, many more margin-parallel mafic dyke swarms were emplaced, along what may have been continental back-arc rifts (Bartels et al., 2016). 8.6. The MB

’ potential role within an AMCG-province

As shown by Figure 13(d), the ~1.63 Ga MBDS (and the Soumenniemi dyke swarm in Finland) are coeval with early rapakivi granites, mostly across Fennoscandia (Rämö and Haapala, 1995), as well as the ~1.65-1.62 Ga Mealy Mountain AMCG-complex in Newfoundland (Emslie and Hunt, 1990; Bybee et al., 2015). It is tempting to combine all of these coeval igneous intrusions into a LIP (e.g., Ernst, 2014), especially considering the possible radiating mafic dyke swarm pattern depicted in Figure 13(a). However, the MBDS and the Soumenniemi dyke swarm do not share the same geochemical signatures in Figure 12(a-b) and do therefore not share the same petrogenesis, as might be expected within a radiating swarm. Thus, we think each coeval swarm formed separately, yet both

were preceded by a major Palaeoproterozoic orogeny with late- to post-orogenic shoshonites, appinites and lamprophyres (Andersson et al., 2006; 2007b; Rutanen et al., 2011; Klausen et al., 2016b). Even if there are no coeval AMCG-complex preserved directly above where a Ketilidian magma chamber source is inferred by the MBDS (nearby rapakivi granites are 1.75-1.73 Ga; Groscott et al., 1999; Garde et al., 2002), such evidence could have been removed during subsequent breakup events and erosion, and we think that the exiting intimate temporal and spatial relationship between the MBDS and AMCG-complexes is sufficient to validate a discussion on whether these are more closely linked and even petrogenetically related. More direct interactions between both mafic and felsic magmas abound within the AMCG-province, as evidences by (1) m

’

h

(including Mealy Mountains) typically

being surrounded by charnockitic and mangaritic intrusions (e.g., Fig. 13c), (2) most rapakivi granites forming bimodal suites together with mafic igneous rocks (e.g., Rämö, 1991; Brown et al., 2003), and (3) in-mixing of more mafic magmas is commonly used to explain the reversely plagioclaseovergrown alkali feldspars crystals that are so diagnostic for rapakivi granites (Hibbard, 1981; Wark and Stimac, 1992). Specifically pertaining to this paper, an additional link to massif anorthosites is provided by the MBDS’ b q

scattered plagioclase phenocrysts (at least across SE Greenland),

which are locally very large and accumulated (Fig. 2f). Many other plagioclase-phyric dykes, including much younger (1.25-1.15 Ga) ‘B F

k ’

h Gardar alkaline province

(Bridgwater, 1967; Bridgwater and Harry, 1968), have for similar reasons been linked to coeval massif anorthosites (Wiebe, 1985). Finally, the most accepted model for the formation of massif anorthosites – and absence of ultramafic cumulates – is that these formed within large, long-lived mafic magma chambers along the base of the continental crust (Ashwal, 1993; Dempster et al., 1999), much like what has been inferred for the source of the MBDS in Section 8.3 and by Figure 8(a). Adopting Ashwal’ (1993) and Dempster et al.’ (1999) model for Proterozoic massif anorthosites, it seem logical to propose that the plagioclase-phyric MBDS was injected from a similar deep crustal magma chamber over a period of >13 ±6 Myr, which could potentially also have generated large volumes of plagioclase cumulates, as well as surrounding crustal melts, as tentatively illustrated in Figure 11(d). If so, the data presented on the MBDS provide valuable constraints on the formation of massif anorthosites, where, for example, its dry mineralogy suggests that magmas were not as hydrous as predicted by Ashwal and Bybee’ (2017) deep magma chamber, having been formed within an arc setting. In addition, the remarkably homogenised and buffered composition of such a long-lived magma chamber indicate that large volumes of plagioclases may simply have formed through >13 ±6 Myrs of dry, high-P and relatively low nucleation versus growth crystallisation and buoyant accumulation into a roof zone. Separate ultramafic cumulates probably also formed during cotectic crystallization of denser olivines, likely sinking back into the mantle as proposed Ashwal (1993) and Dempster et al. (1999).

If the magma chamber was buffered at a dry plagioclase-olivine cotectic temperature over such a long time period, the surrounding dry lower crustal host rocks must invariably also have become extensively melted and thereby produced coeval, felsic and characteristically dry intrusions. There is little evidence of bimodal mafic and felsic magma components mixing into more intermediate magmas, but at most mingling into net-veined complexes (e.g., Rämö, 1991; Brown et al., 2003). In most cases, we agree with Emslie et al. (1994) that more buoyant felsic magmas led the way up into shallower parts of the crust, insulating any coexisting mafic magmas and/or anorthositic mushes that followed along the same pathways (Figure 11d). Such upward propagation probably compare to how other granitic magmas typically are emplaced successively as dyke-sill intrusions (e.g., Grocott et al., 1999), where prolonged emplacement of plagioclase mushes are evidenced by a wide age-span of most massif anorthosites (Ashwal and Bybee, 2017, and references therein). Intuitively, it seems easier for such emplacements to have occurred mainly within an extensional setting, where the four main episodes of AMCG-complex formation in Figure 13(d) are more consistent with intermittent extensional events along a discussed >500 Myr active margin. If the MBDS is related to AMCG-complexes as outlined above, then it remains to be resolved how such long mafic dykes were injected at a relatively high angle from a continental margin, as an intracontinental swarm, and why such margin-perpendicular dyke swarms seem to be restricted to the onset of the Mesoproterozoic AMCG-province. Whereas, the MBDS is more easily explained as failed rift during a ~1.63 Ga breakup of the Nuna supercontinent, it is difficult to envisage how continued AMCG magmatism could have persisted along a resulting passive margin, unless it transformed into an active margin before ~1.46 Ga (cf., Fig. 13d), yet produced similar AMCG-complexes in both instances, as well as repeatedly during at least another two cycles. Thus, more research is needed to resolve this enigmatic tectono-magmatic setting, its repeated generation of AMCG-complexes and how these may relate to associated mafic dyke swarms such as, possibly, the MBDS. 9. Conclusion Together with similar and often diagnostic outcrop features and coinciding structural trends, compositional matching and identical U-Pb ages now confirm that the MBDS across NW Greenland combines – beneath an inland ice cap – into a > 2000 km long and nearly 250 km wide swarm that extends across SE Greenland. The swarm is made up of transitional to trachybasaltic dolerites, with remarkably limited compositional variation (7-4 wt % MgO), explained by cotectic bulk fractionation of ~75% plagioclase and ~25% olivine at ~9 kbar. Along-strike variations in dyke intensities and thickness distributions, are more consistent with a laterally northward propagating MBDS from a single more centralized magma chamber source, most likely located beneath the Ketilidian orogeny. The compositionally homogenous MBDS is best explained by a magma source that was buffered by compositionally similar replenishments, sustaining dry plagioclase-olivine cotectic temperatures for >13 ±6 Myr. Even if wholesale assimilation of more asthenospheric melts by Ketilidian-aged crust cannot be excluded, similar incompatible element patterns as late- to post-orogenioc spessartites are

more readily explained by replenishing melts being generated mainly from a common SCLM, which had been strongly metasomatised above the Palaeoproterozoic subduction zone. This SCLM-source must just have been dehydrated and partially (re)melted at higher temperatures, in order to produce MBDS-magmas that were drier than the spessartites. ~1.63 Ga breakup of the Nuna supercontinent offers the easiest tectono-magmatic explanation for the relatively high temperatures and dry MBDS-magmas to have been emplaced along a failed rift, during the onset of a >500 Myr-long period of Mesoproterozoic AMCG-magmatism. We propose that the above results on the MBDS provide further analogue constraints on existing petrogenetic models for Proterozoic massif anorthosites. Thus, a relatively dry and compositionally buffered and homogenised magma chamber crystallised buoyant plagioclase and denser olivines over an extended period of >13 ±6 Myr. Sustained high temperatures invariably melted the surrounding dry lower crust and thereby produced equally dry and coeval mangeritic, charnockitic and rapakivi granite intrusions. Massive volumes of plagioclase-rich mushes probably utilized conduits instigated by the less dense crustal melts, and thereby formed composite AMCG-complexes during up to four repetitive cycles. We are just uncertain about how such rhythmic AMCG-magmatism could have been sustained during the extended Mesoproterozoic plate tectonic transition from Nuna supercontinent breakup to a reamalgamation into Rodinia, and whether or not the AMCG-province formed entirely along an enduring active continental margin. Acknowledgements All research was generously supported by the Geological Survey of Denmark and Greenland (GEUS), as part of their SEGMENT project. A grant to M. Nilsson from the Royal Physiographic Society, Lund helped pay for TIMS analyses and fieldwork. Amongst the many team members in the SEGMENT 2012 field expedition, we would like to emphasize its main leader Bo Møller Stensgaard, departmental head Karen Hanghøj and field colleagues Thomas Kokfelt, Troels Nielsen and Sebastian Tappe. The paper was greatly improved by constructive reviews by Kristoffer Szilas and Steven Denyszyn. References Andersson, U.B., Eklund, O., Frödjö, S. And Konopelko, D., 2006. 1.8 Ga magmatis in the Fennoscandian Shield; lateral variations in subcontinental mantle enrichment. Lithos 86, 110-136. Andersson, U.B., Neymark, L.A. and Billström, K., 2007a. Petrogenesis of Mesoproterozoic (Subjotnian) rapakivi complexes of central Sweden: Implications from U-Pb zircon ages, Nd, Sr and Pb isotopes. Earth and Environmental Science Transactions of the Royal Society of Edinburgh 92, 201-228. Andersson, U.B., Rutanen, H., Johansson, Å., Manfeld, J. And Rimsa, A., 2007b. Characterization of the Paleoproterozoic Mantle beneath the Fennoscandian Shield: Geochemistry and Isotope Geology (Nd, Sr) of ~1.8 Ga Mafic Plutonic Rocks from the Transscandinavian Igneous Belt in Southeast Sweden. International Geology Review 49, 587-625. Ashwal, L.D., 1993. Anorthosites. Springer Verlag Berlin-Heidelberg, 422 p.

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Analysis no. (number of grains)

U/ Th

Pbc/

206

Pb/

Pbtot1)

204

Pb

207

Pb/ 235

U

± 2s % err

206

Pb/ 238

U

± 2s % err

207

Pb/ 235

U [age, Ma] 1624.1 1624.3 1619.6

206

Pb/ 238

U

raw2) [corr]3) 527259 a (3) 1.6 0.016 3521.5 3.9505 0.42 0.28535 0.27 1618.3 527259 b (4) 1.7 0.032 1733.6 3.9515 0.38 0.28535 0.35 1618.3 527259 c (3) 1.6 0.012 4406.8 3.9286 0.26 0.28417 0.25 1612.4 1) Pbc = common Pb; Pbtot = total Pb (radiogenic + blank + initial). 2) measured ratio, corrected for fractionation and spike. 3) isotopic ratios corrected for fractionation (0.1% per amu for Pb), spike contribution, blank (1 pg Pb and <1 Pb. Initial common Pb corrected with isotopic compositions from the model of Stacey and Kramers (1975) a

565334

North 565333

13.0 5.0 65.10772 65.09406 40.83187 40.79173

565321

527187

527186

0.5 1.7 14.0 64.93662 64.32395 64.32365 40.74964 41.26974 41.27458

527190

Umivik 527191

1.8 8.5 64.30476 64.30437 41.12760 41.12666

527193

527197

527195

26.0 1.5 36.0 64.30352 64.28293 64.28273 41.10951 40.98955 40.98547

48.4

48.4

48.2

49.1

49.9

48.8

49.3

49.3

49.4

49.7

1.97

2.15

2.09

2.31

1.79

2.13

2.05

1.84

2.29

1.97

15.38 7.84

14.82 7.59

14.89 7.55

14.49 7.10

15.51 6.41

15.02 7.51

15.13 7.64

16.30 7.77

14.58 7.42

15.60 7.66

15.2

15.4

15.3

15.8

12.9

15.4

15.2

13.6

16.0

15.1

1.79 5.38 0.20

2.04 4.65 0.20

1.94 4.96 0.21

2.19 4.57 0.21

2.05 5.25 0.18

2.10 4.98 0.21

1.98 4.98 0.20

1.78 5.06 0.18

2.10 4.61 0.21

1.96 5.06 0.19

3.02

2.94

2.92

3.05

2.90

2.97

2.98

3.13

3.10

3.24

0.43 -0.36 99.3

0.48 0.59 99.2

0.46 0.53 99.0

0.55 0.30 99.7

0.36 2.51 99.8

0.45 0.53 100.1

0.43 0.22 100.1

0.42 0.78 100.2

0.49 0.39 100.7

0.39 -0.57 100.4

31.5 269 67 77 63 98 147 29 385 31 135 5.59 0.99 0.56 865 22.5 49.1 6.63 26.9 6.09 2.14 5.93 1.01 5.51 1.22 3.22 0.47 2.99

34.0 289 63 80 43 57 139 36 375 34 153 6.33 1.07 1.12 965 25.5 56.6 7.55 30.8 7.02 2.39 6.55 1.13 6.14 1.36 3.48 0.51 3.28

32.4 280 62 123 51 55 144 33 373 32 147 6.16 0.92 0.91 942 24.6 53.7 7.10 28.6 6.51 2.23 6.33 1.05 5.76 1.24 3.22 0.48 2.86

28.2 225 44 49 37 59 143 40 304 30 145 6.54 0.95 2.01 762 23.5 51.6 6.77 30.2 6.55 1.90 6.16 0.88 5.47 1.13 3.48 0.45 2.93

22.9 192 32 47 64 76 157 67 293 22 102 4.48 0.91 2.28 681 16.6 36.0 4.73 20.1 4.46 1.28 4.37 0.65 4.04 0.81 2.37 0.33 2.12

28.4 224 62 50 44 58 129 47 311 26 122 5.25 0.83 2.41 706 18.6 41.2 5.41 24.5 5.58 1.86 5.42 0.79 5.07 1.05 2.81 0.41 2.78

27.9 216 61 58 51 55 130 34 333 26 123 5.05 1.06 1.09 748 19.4 42.6 5.50 24.3 5.42 1.90 5.57 0.76 5.05 0.96 2.78 0.39 2.70

24.6 196 72 56 66 58 126 28 381 24 112 4.50 0.86 1.50 706 17.2 38.1 5.12 22.7 5.19 1.71 5.02 0.73 4.61 0.91 2.57 0.36 2.47

29.3 247 66 60 37 67 152 38 303 30 137 5.82 1.04 1.38 763 21.3 46.2 6.06 27.2 6.06 1.84 5.94 0.88 5.50 1.17 3.19 0.45 2.95

26.7 222 80 65 44 73 132 34 301 26 120 5.14 0.77 0.82 654 18.6 40.2 5.22 22.8 5.30 1.66 5.17 0.77 4.76 0.99 2.69 0.39 2.52

0.48 3.53 0.28 7.66 2.19 0.84

0.51 4.18 0.34 8.01 2.22 0.79

0.50 3.91 0.35 7.71 2.11 0.74

0.45 3.75 0.36 12.18 2.19 0.84

0.34 2.70 0.26 35.83 2.24 0.81

0.38 3.28 0.25 8.79 2.06 0.79

0.38 3.29 0.29 9.77 2.26 0.86

0.36 2.93 0.24 7.39 1.79 0.74

0.43 3.65 0.31 23.66 2.53 1.10

0.38 3.37 0.28 10.02 2.62 1.17

Confirmed southward continuation of the Mesoproterozoic Mellville Bugt Dyke Swarm Thickness distributions consistent with lateral dyke propagation Presence of plagioclase megacrysts link dykes to coeval massif anorthosites Geochemical signatures consistent with a sub-continental lithospheric mantle source