Geochemistry of global ca. 1880 Ma LIP magmatism: is there evidence for a comagmatic origin and connections in supercontinent reconstructions?
Matthew Minifie1, Richard Ernst2, and Andrew Kerr1
1School of Earth and Ocean Sciences, Cardiff University, Main Building, Park Place, Cardiff CF10 3AT, UK; MinifieMJ@cardiff.ac.uk
2Ernst Geosciences, 43 Margrave Ave., Ottawa, Ontario K1T 3Y2, Canada
Magmatic rocks aged ca. 1880 Ma are located in a number of widespread locations around the world. Magmatic rocks of such age are found in the Aldan Shield, Baffin Island, Baltic Shield, Bastar Craton, Dharwar Craton, Greenland, Kaapvaal Craton, North Australian Craton, Outer Hebrides, Slave Craton, Superior Craton, Wyoming Craton and Zimbabwe Craton. The quality and quantity of geochemical data for each area are variable, which impedes attempts to reconstruct the relative positions of the cratonic blocks at ca. 1880 Ma. The available data suggest that Baffin Island and Superior Craton magmatic rocks do not share a common origin as other ca. 1880 Ma magmatic rocks. Rocks from Australia, India and southern Africa share similar geochemical signatures, which possibly indicates that these regions were in close proximity at ca. 1880 Ma. The Baltic Shield and Slave Craton also share similar geochemical signatures. However, much more geochemical analyses are needed on rocks from nearly all of the regions. In particular, isotopic studies could be a useful tool in continental reconstruction.
The advancement of U-Pb zircon/baddeleyite geochronology in the past 25 years has led to the recognition of many new large igneous provinces (LIPs) and LIP fragments, particularly those of Precambrian age. The method of linking coeval LIPs or LIP fragments on different cratonic blocks is currently being used to reconstruct the position of the continents back through Earth’s history (e.g., Bleeker & Ernst, 2006; Ernst & Bleeker, 2010; www.supercontinent.org). It is now known that there are magmatic suites of ca. 1880 Ma age located on a number of widely separated Archaean cratons (Fig. 1). Such magmatism is found in North America (Slave, Superior and Wyoming Cratons and Baffin Island), Europe (Baltic Shield and Outer Hebrides), Greenland, Siberia (Aldan Shield), India (Bastar and Dharwar Cratons), southern Africa (Kaapvaal and Zimbabwe Cratons) and Australia (North Australian Craton).
Figure 1. Current global distribution of ca. 1880 Ma magmatic rocks.
It may be possible that some of the ca. 1880 Ma magmatic provinces are comagmatic which would indicate that the cratonic blocks where they are now located were once juxtaposed or near neighbours. However, geochronology alone cannot confirm this. Additional palaeomagnetic and geochemical data are required to ascertain if widely dispersed magmatic rocks were once part of the same LIP. This report will assess the available geochemical data from ca. 1880 Ma magmatic rocks to see if there are common signatures which could represent a comagmatic origin and suggest which cratonic blocks were in close proximity in the mid-Palaeoproterozoic.
Australian ca. 1880 Ma magmatism
Magmatic rocks of ca. 1880 Ma age have been recorded in the Halls Creek Orogen of northwestern Australia. These rocks are known as the Biscay Formation and consist of massive metabasalts, mafic volcaniclastics and fragmental deposits, with minor amounts of metadolerite sills, sediments and felsic volcanics. Felsic rocks of the Biscay Formation have a U-Pb zircon age of 1880 ± 3 Ma (Blake et al., 1998, cited in Sheppard et al., 1999). Major and trace element data exist for the Biscay Formation (Sheppard et al., 1999). However, the quality of the trace element data may be questionable given the fractionation observed between elements considered geochemical twins (e.g., Nb-Ta, Zr-Hf) (Fig. 3). The Biscay Formation has a restricted range in MgO of 6.60-7.57 wt% (Fig. 2) but a more varied trace element composition. On the basis of trace element patterns, the Biscay Formation can be split into two groups. One group of samples has relatively flat rare earth element (REE) and multi-element profiles whereas the other group has light-REE (LREE) enriched profiles and negative Nb-Ta-Ti anomalies (Figs. 3 and 4).
Figure 2. SiO2 vs. MgO diagram for ca. 1880 Ma magmatic rocks from Australia, Baffin Island, Baltic Shield, India, Slave Craton, southern Africa and Superior Craton. Some samples from Baffin Island, India and Superior Craton plot at higher MgO concentrations that represented on the x-axis and are identified by coloured arrows annotated with the highest MgO concentration. The position of the arrows also mark the SiO2 concentration at the maximum MgO concentration. Data sources are given in main text.
Figure 3. Primitive-mantle-normalised multi-element diagrams for ca. 1880 Ma magmatic rocks. Data sources are given in main text except for southern Africa where data sources are given in the figure. More data exist for the Superior Craton but only a subset is shown for clarity.
Figure 4. (La/Sm)N-(Gd/Yb)N, Nb/Y-Zr/Y, Zr/Nb-Nb/Th and Zr/Ti-Nb/Y diagrams for ca. 1880 Ma magmatic rocks. See Fig. 2 for legend. Data sources are given in main text.
Baffin Island ca. 1880 Ma magmatism
The Bravo Lake Formation is a sequence of metamorphosed igneous and sedimentary rocks that extends for ~120 km across Baffin Island (St-Onge et al., 2004; Johns et al., 2006). There is only a very imprecise U-Pb age determination of 1916 ±35 Ma for the Bravo Lake Formation (Wodicka, 2004, cited in Johns et al., 2006). Johns et al. (2006) have published major and trace element data for the Bravo Lake Formation. The igneous rocks have a wide ranging MgO content (0.57-19.58 wt%) (Fig. 2) and a fairly uniform trace element composition. The REE and multi-element profiles are LREE-enriched and heavy REE (HREE) depleted and show prominent positive Nb-Ta anomalies (Figs. 3 and 4).
Baltic Shield ca. 1880 Ma magmatism
There is a multitude of literature reports on igneous rocks in the Baltic Shield that are thought to be ~1.9 Ga in age. Therefore it is hard to do justice to the geochemical characteristics of Baltic Shield magmatism in this short report. Magmatism of the appropriate age is mainly found in volcanic arc complexes within the Finnish and Swedish components of the Svecofennian Orogen. There is a large range of igneous rocks in the Svecofennian Orogen including ultramafic to felsic lavas, mafic and felsic tuffs, mafic dykes and granodiorite, diorite and tonalite intrusions (e.g., Kahkonen, 1987; Vaisanen et al., 2002). The age ranges of these arc complexes are largely constrained between 1904 and 1867 Ma (Kahkonen et al., 1989; Suominen, 1991; Nironen, 1999; Nironen et al., 2000; Vaisanen et al., 2002).
Despite the seeming abundance of ca. 1880 Ma magmatic rocks in the Baltic Shield there is a distinct lack of good quality major and trace element data from modern geochemical analytical techniques for these rocks, at least in the accessible academic literature. However, Vaisanen & Westerlund (2007) have published a complete major and trace element dataset for mafic volcanic rocks from the Turku area in southern Finland. These rocks vary in MgO from 2.68 to 13.64 wt% (Fig. 2) and have total alkali contents of up to 5.77 wt%. The majority of these rocks have trace element patterns which are LREE-enriched and show negative Nb-Ta-Ti anomalies (Fig. 3 and 4).
Greenland ca. 1880 Ma magmatism
The Nagssugtoqidian Orogen consists largely of Archaean and Palaeoproterozoic orthogneisses intercalated with some Palaeoproterozoic supracrustal and magmatic rocks (Kalsbeek & Nutman, 1996). The central zone of the orogen contains the Arfersiorfik and Sisimiut charnockite igneous suites. The Arfersiorfik igneous suite mainly consists of a large body of quartz diorite exposed over several hundred square kilometres plus outliers of quartz diorite to tonalite gneiss sheets (van Gool et al., 2002). The Sisimiut charnockite suite is a ~30 km wide complex of tonalites, diorites, granites, leuconorites, gabbros and hornblende pyroxenites (van Gool et al., 2002). Connelly et al. (2000) have reported U-Pb ages from the Arfersiorfik and Sisimiut igneous suites of 1885 +6/-3 Ma and 1873 +7/-4 Ma respectively. In eastern Greenland the Ammassalik mobile belt contains the Angmagssalik igneous complex of ultrabasic to acidic rocks. A mafic sample from the Angmagssalik complex yielded a U-Pb age of 1886 ± 2 Ma (Hansen & Kalsbeek, 1989, cited in Bridgwater et al., 1990). Unfortunately a complete and reliable major and trace element dataset for the Greenland magmatism does not exist.
Indian ca. 1880 Ma magmatism
Magmatism of ~1880 Ma age is found in both the Bastar and Dharwar Cratons which form part of the Archaean shield complex of peninsular India. In the south of the Bastar Craton the northwest-trending BD2 dolerite dyke swarm intrudes basement and supracrustal rocks. The swarm outcrops over a minimum area of ~1000 km2 and has yielded two U-Pb ages of 1891.1 ± 0.9 Ma and 1883 ± 1.4 Ma (French et al., 2008). There are many more northwest-trending dolerite dykes with an areal extent of ~17,000 km2 in the southern Bastar Craton (Ramachandra et al., 1995) which may be the same age. Magmatic rocks of similar age to the dykes in the Bastar Craton are also found in association with the Cuddapah basin in the Dharwar Craton. French et al. (2008) obtained a U-Pb baddeleyite age of 1885.4 ± 3.1 Ma from the Pulivendla sill which intrudes the Tadpatri Formation in the western part of the Cuddapah basin. There are numerous mafic dykes around the basin which are undated using reliable techniques. However, one east-west trending dyke to the southwest of the Cuddapah basin has yielded an Ar-Ar whole rock age of 1879 ± 5 Ma (Chatterjee & Bhattacharji, 2001).
The most complete major and trace element dataset of ca. 1880 Ma Indian magmatic rocks has been obtained using ICP-OES and ICP-MS at Cardiff University, UK, on samples collected by Rajesh Srivastava from dykes in the Bastar Craton. The samples come from various sections of the dykes and have a total range in MgO of 1.79-29.20 wt% (Fig. 2). All the samples share a similar trace element pattern which is LREE-enriched with flat HREEs and negative Nb-Ta-Ti anomalies (Figs. 3 and 4).
Outer Hebrides ca. 1880 Ma magmatism
The ca. 1880 Ma magmatism in the Outer Hebrides, UK, is found in the South Harris Igneous Complex which is comprised of three large meta-igneous bodies a metanorite, a metadiorite and a meta-anorthosite) and numerous smaller metamorphosed basic and ultrabasic intrusions (Fettes et al., 1992). Mason et al. (2004) produced U-Pb zircon ages from the three main igneous bodies of the South Harris Igneous Complex. The age of the meta-anorthosite is 2491 +31/-27 Ma but the ages of the metanorite and metadiorite are 1890 +2/-1 Ma and 1888 ± 2 Ma respectively. There is no complete and reliable major and trace element dataset for the South Harris Igneous Complex in the literature. However, a collection of samples from the metanorite and metadiorite bodies exists at Cardiff University, UK.
Siberian ca. 1880 Ma magmatism
The Aldan Shield is one of the accreted provinces which form the Siberian Platform. Mafic dykes are known to intrude the granite-greenstone and high-grade gneiss terranes of the Aldan Shield (Ernst et al., 1996; Jahn et al., 1998). A northeast-trending mafic dyke swarm known as the Kalaro-Nimnyrsky swarm is intrusive into the Aldan Shield, extending for >300 km and covering an area of ~215,000 km2 (Ernst & Buchan, 2001; Gladkochub et al., 2010). An Ar-Ar hornblende age of 1866 ± 9 Ma has been reported from these northeast-trending dykes (unit 3b in Gladkochub et al., 2010). These dykes may be related to the magmatic rocks of the Akitkan Group and Selenga-Stanovoi superterrane further to the south in the Aldan Shield and which have U-Pb zircon ages ranging from 1878 ± 4 Ma to 1863 ± 9 Ma (Buchko et al., 2006; Didenko et al., 2009). These rocks have not been analysed for their major and trace element chemistry.
Slave Craton ca. 1880 Ma magmatism
The Ghost dolerite-gabbro dyke swarm forms part of the ~1880 Ma magmatic activity in the Slave Craton (Buchan et al., 2010). The Ghost dykes are found in the southwestern portion of the craton to the northeast of Great Slave Lake between Yellowknife and Ghost Lake and trend in a northeasterly direction (Henderson, 1998; Pehrsson, 2002). Davis & Bleeker (2007) obtained a U-Pb age of 1884.4 ± 3.4 Ma for the Ghost dykes. Further north in the Slave Craton than the Ghost dykes is the Wopmay Orogen which is a ~1.9 Ga north-trending orogenic belt that assembled along the western margin of the craton. The Wopmay Orogen contains a variety of igneous rocks ranging from mafic intrusive rocks to felsic plutonic rocks, most of which are either imprecisely dated at ~1880 Ma or are slightly younger than ~1880 Ma. In the Great Bear Magmatic Zone of the Wopmay Orogen the age range of magmatic rocks is from 1876 ± 10 Ma to 1843 ± 5 Ma (Bowring et al., 1984; Ghandi et al., 2001). The Hepburn Intrusive Suite occurs further east in the Wopmay Orogen than the magmatic rocks of the Great Bear Magmatic Zone. The Hepburn Intrusive Suite consists of approximately one hundred plutons of gabbro to granite composition. The best age estimate for this intrusive suite is constrained by U-Pb zircon dating to be 1885 Ma (Hoffman & Bowring, 1984). A volcanic ash bed related to the intrusions has also yielded a U-Pb zircon age of 1882 ± 4 Ma (Bowring & Grotzinger, 1992).
A north-south zone of gabbroic sills ~200 km long and ~10 km wide intrude the Wopmay Orogen and are known as the Morel sills. Davis & Bleeker (2007) regarded the Morel sills to be contemporaneous with the Ghost dyke swarm. This age estimate is largely based on the syncollisional nature of the sills; the Morel sills intrude the Wopmay Orogen deposits but are not deformed. The Mara River sheets of the Kilohigok basin (Fahrig, 1987) occur eastwards from the Wopmay Orogen in the Slave Craton. Davis & Bleeker (2007) obtained a U-Pb age of 1870 Ma for these intrusive sheets.
A complete and accurate major and trace element dataset exists for the Ghost dykes, Morel sills and Mara River sheets. These data were produced by ICP-OES and ICP-MS at Cardiff University, UK, from samples in the archives of the Geological Survey of Canada and the University of Ottawa. The MgO content of the Ghost dykes varies between 5.46 and 8.44 wt% while the respective range in the Morel sills and Mara River sheets is 5.82-8.90 wt%. Negative Nb-Ta anomalies are prevalent in the trace element patterns of all three suites (Fig. 4). The main difference in the trace element patterns is that the Morel sills and Mara River sheets are much more enriched in the LREEs than the Ghost dykes.
Southern African ca. 1880 Ma magmatism
Ca. 1880 Ma magmatism is preserved in both the Kaapvaal and Zimbabwe Cratons in southern Africa. Hanson et al. (2004) obtained U-Pb baddeleyite ages of 1878.8 ± 0.5 Ma, 1873.7 ± 0.8 Ma and 1871.9 ± 1.2 Ma from doleritic sills intruding the Waterberg Group sediments in the Kaapvaal Craton. Palaeomagnetic studies conducted by Hanson et al. (2004) on the ~1878-1872 Ma samples found that the poles of the dolerite sills are antipodal to the poles of basaltic lavas and sills belonging to the Soutpansberg Group (Bumby et al., 2001) to the north of the Waterberg Group. This suggests that the Soutpansberg magmatism is coeval with and was emplaced as part of the same event as the ~1878-1872 Ma Waterberg Group magmatism but during a reversed magnetic polarity chron. The Mashonaland sills cover an area of ~160,000 km2 in the northeastern Zimbabwe Craton. These sills have yielded U-Pb baddeleyite ages of 1877 ± 2.2 Ma, 1885.9 ± 1.8 Ma and 1875.6 ± 1.6 Ma (Soderlund et al., 2010). The east-west trending Mazowe River dykes may belong to the same igneous episode as the Mashonaland sills given the similar palaeomagnetic directions and a Rb-Sr whole rock age for the Mazowe River dyke of 1870 ± 600 Ma (Wilson et al., 1987).
The most complete major and trace element dataset exists for feeder dykes to the Soutspansberg Group lavas (Klausen et al., 2010). Good quality data also exist for four samples of post-Waterberg Group dolerites (Hanson et al., 2004). There are other data that exist for other southern African units but these data are incomplete and missing certain crucial trace elements. The overall range in MgO of the southern African magmatism is 2.63-11.09 wt% (Fig. 2). The trace element patterns of all the southern African units are fairly similar and share negative Nb-Ta anomalies (Figs. 3 and 4).
Superior Craton ca. 1880 Ma magmatism
A summary of the ca. 1880 Ma magmatism in the Superior Craton is given in Heaman et al. (2009) and Minifie (2010). The majority of the magmatic rocks circumscribe the Superior Craton margins for >3400 km, although some intrusive rocks are also present within the Craton interior. A range of lavas, dykes, sills and layered intrusions and a range of lithologies and compositions are present in the Superior Craton. Minifie (2010) has produced a large major and trace element dataset for the Superior Craton magmatism. The overall range in MgO content of the lavas is 0.50-24.83 wt% but some cumulative sections of intrusions have MgO concentrations of up to 42 wt% (Fig. 2). The trace element signatures of the Superior Craton magmatism are variable although the dominant signature present throughout much of the craton is that of an oceanic plateau signature (i.e., flat REE and multi-element profiles) (Figs. 3 and 4). However, rocks with LREE-enrichment and negative Nb-Ta anomalies are not uncommon and rocks with positive Nb-Ta anomalies also exist (Fig. 3).
Wyoming Craton ca. 1880 Ma magmatism
The extent of ca. 1880 Ma magmatism is not well constrained in the Wyoming Craton. The Prairie Creek gabbroic sill intruding through muds and carbonates in the Black Hills of South Dakota has a U-Pb age of 1883 ± 5 Ma (Redden et al., 1990). This sill is part of a north-trending, vertically-dipping belt of metagabbro intrusions. However, this belt of intrusions contains at least two episodes of magmatism as Redden et al. (1990) obtained a U-Pb age of 1964 ± 15 Ma from another sill within the belt. A U-Pb age of 1884 ± 29 Ma has also been reported from an alkalic metatuff layer ~20 km west of the Prairie Creek gabbroic sill (Redden et al., 1990). Van Boening & Nabelek (2008) have analysed samples of the ca. 1880 Ma magmatic rocks. However, their data are missing several REEs and show a large fractionation between Nb and Ta and Zr and Hf which calls into question the quality of the data.
Does the geochemistry argue for any of the cratons being nearest neighbours at ca. 1880 Ma?
It is hard to conclusively prove from the geochemistry of the various ca. 1880 Ma magmatic provinces where each of the cratons were positioned relative to each other at ca. 1880 Ma. The Baffin Island magmas have drastically different trace element compositions and ratios to the other provinces. Given the large uncertainty of the age of the Baffin Island magmatism, it is possible that it is completely unrelated to other ca. 1880 Ma magmas. This notion is not contradicted by the trace element geochemistry. It is also feasible that the Superior Craton magmatism is unrelated to other coeval magmatic rocks. Flat, oceanic plateau-like trace element profiles and high MgO lavas (>12 wt% MgO), with the exception of just one 13.64 wt% sample from the Baltic Shield, are only found in the Superior Craton. This could suggest that the Superior Craton was isolated from the other cratons at ca. 1880 Ma.
Previous studies on the petrogenesis of the magmatism in the Slave Craton and Baltic Shield have assigned a subduction-related volcanic arc origin to these magmatic rocks. This is certainly consistent with the lithologies present in these regions and to the prevalent negative Nb-Ta anomalies in the trace element profiles. On the basis of geochemistry it is possible that the Slave Craton and Baltic Shield magmatism were part of the same arc system. However, if the Palaeoproterozoic continental reconstructions of Zhao et al. (2002) and Hou et al. (2008) (Fig. 5) are correct, then it seems unlikely that the Slave Craton and Baltic Shield could have been neighbours given the length the arc system would need to be.
Figure 5. Two possible reconstructions of the supercontinent Columbia at ~1.85 Ga from (a) Hou et al. (2008) and (b) Zhao et al. (2002). The black segments in both figures represent ~2.1-1.8 Ga orogenic belts. Precise estimates of palaeolatitude are not given in the reconstruction of Zhao et al. (2002).
Despite the prevalence of negative Nb-Ta anomalies in the trace element profiles of the magmatic rocks from Australia, India and southern Africa, previous studies of these rocks have not assigned to them an origin related to arc volcanism. Instead these rocks are considered to have originated from enriched subcontinental lithospheric mantle or asthenospheric melts contaminated by enriched subcontinental lithospheric mantle. As the magmatic rocks from these three regions have similar trace element profiles and ratios, it seems possible that they could have originated from the source region and, therefore, that the regions were in close proximity to each other ca. 1880 Ma. If a large, hot mantle plume head was the trigger for melting of the asthenosphere or lithospheric mantle, then the relative positions of Australia, India and southern Africa in the reconstruction of Zhao et al. (2002) (Fig. 5) may have some merit.
A problem with using the geochemistry of igneous rocks in palaeocontinental reconstructions is that magmatic processes such as partial melting and fractional crystallisation can alter geochemical signatures so that magmas originating from the same mantle source region can have vastly different geochemical compositions. Radiogenic isotopes (e.g., 87Sr/86Sr, 143Nd/144Nd, 206Pb/204Pb, 208Pb/204Pb, 176Hf/177Hf, 187Os/188Os) are not altered by differing degrees of partial melting and fractional crystallisation and could be an efficient tool in linking up once contiguous magmatic provinces. Unfortunately, a Sr-Nd-Pb-Hf-Os isotopic dataset only exists for the Superior Craton magmatism (Minifie, 2010).
Recommended future work
Funding from NERC PhD studentship NER/S/A/2006/14009 and BHP Billiton is acknowledged. This work represents part of the Matthew Minifie’s PhD at Cardiff University (Minifie, 2010). We are very grateful to all those who contributed samples towards this study, including: Bob Baragar, Guy Desharnais, Don Francis, Henry Halls, Larry Hulbert, Steve Kissin, Andre Lalonde, Jim Mungall, Steve Sheppard, Rajesh Srivastava and Vince Vertolli. Andrey Bekker, Gerry Benger, Ken Buchan, Bill Cannon, Tim Corkery, Peter Dahl, Norman Halden, Richard Herd, Garth Jackson, Dan Layton-Matthews, Mike Lesher, Richard Ojakangas and Sally Pehrsson were also very helpful in obtaining samples for this study. Iain McDonald and Ley Woolley were very efficient at running sample solutions on the ICP-OES and ICP-MS at Cardiff University.
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