September 2009 LIP of the Month

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2027-2023 Ma Lac de Gras-Booth River magmatic event in the Slave craton of North America

Kenneth L. Buchan, Anthony N. LeCheminant and Otto van Breemen
Geological Survey of Canada, 601 Booth Street, Ottawa, ON K1S 0E8, Canada;

Extracted and slightly modified from Buchan et al. (2009)

The Slave Province is one of several Archean cratons that amalgamated to form Laurentia (Fig. 1) by the late Paleoproterozoic (Hoffman 1988). The geographic relationship in earlier times among these cratons and with other cratons that now reside in other continents is largely unknown. However, it is the subject of much interest because of its bearing on important issues such as the nature of early supercontinents, the extent of mineral-rich belts within such supercontinents, the supercontinent cycle and the timing of the onset of plate tectonics (e.g., Gurnis 1988; Rogers 1996; Bleeker 2003; Condie and Pease 2008).

Figure 1: Tectonic map of Laurentia with Greenland shown in a pre-Mesozoic reconstruction. The star locates the Lac de Gras sampling area. MCR is Mid-Continent Rift. Figure modified after Hoffman (1988) and Buchan et al. (2009).

Numerous Paleoproterozoic diabase dyke swarms were emplaced into the Slave craton prior to the formation of Laurentia (e.g., Buchan and Ernst 2004; Stubley 2005; Bleeker et al. 2007). In addition to the Lac de Gras dykes discussed herein, several other Paleoproterozoic swarms intruded the southern and central Slave Province (Fig. 2), including the ca. 1884 Ma NE-trending Ghost swarm, ca. 1902 Ma ENE-Hearne swarm, ca. 2108 Ma NW-trending Indin swarm, ca. 2189 Ma NNE-trending Dogrib swarm, ca. 2210 Ma E-trending MacKay (or ‘X’) swarm, and ca. 2230 Ma NE-trending Malley swarm. These swarms are all crosscut by the extensive Mesoproterozoic Mackenzie swarm (1267 ±2 Ma; LeCheminant and Heaman 1989) that forms a giant fan over much of the northern Canadian Shield (Fahrig 1987; Ernst and Baragar 1992). NW-trending ‘305’ dykes, that may be roughly coeval with the Mackenzie swarm on the basis of paleomagnetic evidence (K. Buchan in LeCheminant 1994), also crosscut several of the Paleoproterozoic swarms in the sampling area. Many of the Paleoproterozoic swarms may represent extensional events associated with progressive breakup of a large ancestral craton (LeCheminant et al. 1996; Davis and Bleeker 2007). It is likely that remnants of these swarms exist on other, as yet unidentified, cratonic blocks that were once contiguous with the Slave craton.

Figure 2: Paleoproterozoic dyke swarms of the Slave Province, including ca. 1884 Ma NE-trending Ghost dykes (Atkinson 2004; Davis and Bleeker 2007; Bleeker et al. 2008a), ca. 1902 Ma ENE-trending Hearne dykes (Pietrzak 2003; Bleeker et al. 2008b), ca. 2027-2023 Ma N-trending Lac de Gras dykes (Buchan et al. 2009), ca. 2108 Ma NW-trending Indin dykes (McGlynn and Irving 1975; Davis and Bleeker 2007), ca. 2189 Ma NNE-trending Dogrib dykes (McGlynn and Irving 1975; LeCheminant et al. 1997), ca. 2210 Ma E-trending MacKay (or ‘X’) dykes (McGlynn and Irving 1975; Fahrig et al. 1984; LeCheminant et al. 1995), and ca. 2230 Ma NE-trending Malley dykes (Frith 1987; LeCheminant et al. 1995). BRIC is Booth River igneous complex; KB is Kilohigok Basin. Figure modified after Buchan and Ernst (2004) and Buchan et al. (2009).

During ca. 1.9-1.8 Ga amalgamation of Laurentia, the Slave craton was affected by collisional events along both its eastern and western margins. As a result, the Paleoproterozoic dyke swarms are more metamorphosed towards the Thelon Tectonic Zone on the east and Wopmay Orogen on the west (e.g., Henderson et al. 1990; Frith 1993).

The Lac de Gras dyke swarm (Fig. 2) is distinctive in trend and geochemistry. It consists of an ~100 km-wide array of ~010º-trending, steep to vertical, diabase dykes (LeCheminant 1994). It extends north more than 300 km from Lac de Gras to the Booth River igneous complex (Roscoe et al. 1987; Hulbert 2005), a large layered mafic-ultramafic-felsic intrusion overlain by Goulburn Group sedimentary rocks of the Kilohigok Basin. Dykes of similar trend to the north of the Booth River intrusion (Fahrig 1987; Buchan and Ernst 2004) may also be part of the Lac de Gras swarm (Fig. 2). South of Lac de Gras, the swarm can be traced to the vicinity of Snap Lake at latitude 63.5ºN (Stubley 2005).

Geochemical characteristics of Lac de Gras dykes are distinct from those of other Proterozoic diabase dykes in the southern and central Slave Province (Wilkinson et al. 2001; Ernst and Hulbert 2003). Lac de Gras dykes plot in the alkaline field of a total alkali-silica diagram, whereas other Proterozoic swarms near Lac de Gras are subalkaline in composition (Wilkinson et al. 2001). Lac de Gras dykes have higher light rare-earth element content and steeper rare-earth element patterns than those of other swarms. Their chemistry is consistent with alkalic intraplate basaltic magmatism, whereas older Paleoproterozoic Malley and Mackay dykes and younger Mesoproterozoic Mackenzie and ‘305’ dykes have signatures typical of continental tholeiitic magmatism (Wilkinson et al. 2001). 

Two Lac de Gras dykes have been dated (Buchan et al. 2009) at 2023 ±2 Ma and 2027 ±4 Ma (U-Pb baddeleyite). These ages are similar to U-Pb zircon ages of 2023 +4/-2 Ma for the felsic phase (Roscoe et al. 1987) and 2026 ±1 and 2025 ±1 Ma for the mafic phase (Davis et al. 2004) of the Booth River igneous complex. Therefore, the Lac de Gras dyke swarm and Booth River intrusion are coeval and part of a magmatic event that occurred shortly before collision(s) along the eastern margin of the Slave Province that resulted in formation of the 2.0-1.9 Ga Thelon Tectonic Zone (Henderson et al. 1990). An ENE-trending dyke in the Great Slave Lake Shear Zone (Fig. 2) along the southern margin of the Slave Province has been dated at 2038 ±3 Ma (Pehrsson et al. 1993), only a few million years older than the Lac de Gras dykes and Booth River intrusion.

Several magmatic units on other cratonic blocks around the world are of similar age to Lac de Gras-Booth River magmatism: ca. 2008 Ma diabase dykes and ca. 2031 Ma Cheela Springs volcanics of the Pilbara craton (Müller et al. 2005), the 2011 ±1 Ma Kennedy dyke swarm in the Wyoming Province of North America (Cox et al. 2000), 2038 +4/-2 Ma Korak sills of the Cape Smith belt of North America (Machado et al. 1993), the 2040 Ma Kangǎmiut dyke swarm of western Greenland (Nutman et al. 1999), 2040 ±6 Ma diabase dykes in the West African craton (Walsh et al. 2002), and 2045 ±3 Ma Iglusuataliksuak dykes of the Nain Province of eastern North America (Hamilton et al. 1998). Some of these units may be related to Lac de Gras-Booth River magmatism, having rifted from the Slave craton during continental breakup, and hence comprise parts of a larger mafic igneous province. Alternatively, they may represent near-coeval but distinct magmatic events (e.g., Ernst and Buchan 2002). Although a primary paleomagnetic pole has been reported for the Lac de Gras dykes (Buchan et al. 2009), primary paleopoles have not been determined for any of the near-coeval units on other cratons. Therefore, reconstructions cannot yet be made to test whether any of these blocks were close to the Slave craton when the Lac de Gras dykes and Booth River intrusion were emplaced.


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