Protopangaea: Palaeomagnetic definition of Earth's oldest (mid-Archaean-Palaeoproterozoic) supercontinent

Journal of Geodynamics 50 (2010) 154–165

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Protopangaea: Palaeomagnetic definition of Earth’s oldest (mid-Archaean-Palaeoproterozoic) supercontinent J.D.A. Piper Geomagnetism Laboratory, Department of Earth and Ocean Sciences, University of Liverpool, Liverpool L69 7ZE, UK

a r t i c l e

i n f o

Article history: Received 9 October 2009 Received in revised form 18 December 2009 Accepted 1 January 2010

Keywords: Archaean Proterozoic Supercontinents Palaeomagnetism Pangaea Geoid

a b s t r a c t The geological record of consolidation and uplift of crustal protolith identifies three major groupings of Archaean cratonic nuclei (Rogers, 1996). These comprise ‘Ur’ (∼3.0 Ga cratons of Southern Africa, Western Australia), ‘Arctica’ (∼2.5 Ga cratons of Greenland, Fennoscandia, Laurentia and Siberia) and ‘Atlantica’ (∼2.0 Ga cratons of Western Africa and South America). In this paper the ∼2.9–2.0 Ga record from these cratons is used to develop a palaeogeographic model based on (i) the proximities recognised from geological evidence and (ii) the correlation of ∼2.7–2.2 Ga palaeomagnetic poles between Ur and Arctica (Piper, 1982, 2003). The solution shows that Archaean continental crust had aggregated into a crescent-shaped supercontinent, ‘Protopangaea’, by mid-Archaean times. The outer wings comprised core elements of Ur and Arctica which appear to have retained internal quasi-integrity with respect to each other until the end of the Proterozoic eon. Crust in between was still consolidating in early Palaeoproterozoic times and would ultimately include the Atlantica cratonic grouping. Geological tracers, including straight belts and mineral provinces, form lineaments parallel to the arcuate shape of the protocontinent and, as presently defined, the reconstruction indicates that a minimum of 35–56% of the present continental crust had been extracted from the mantle by these times. Derivation of the reconstruction is facilitated because the bulk of the ∼2.7–2.2 Ga poles imply very low rates of apparent polar wander (APW, <10 mm/year) during this ∼5 Ga interval. These low APW rates correlate with the later granite–greenstone tectonics and are presumed to reflect dominant small scale convection. In contrast rapid APW after ∼2.2 Ga records a transition to typical Proterozoic tectonics with large scale mobility reflected in the accretion of mobile belts to the core Ur-Atlantica-Arctica assemblage. The reason why continental crust accretes into supercontinents of symmetrical crescent-shaped form is also addressed. By analogy with Phanerozoic Pangaea and the present-day geoid it is interpreted to have resulted from large-scale, presumably whole mantle, convection systems driving the continental crust towards regions of minimum gravitational potential. © 2010 Elsevier Ltd. All rights reserved.

1. Introduction The earliest assemblages of continental crust comprise the Archaean and early Proterozoic nuclei that have been involved in the accretion and break-up of all the subsequent supercontinents. An insight into their origins can be recognised from the pattern of cratonic stabilisation ages, and from the timing of subsequent uplift and cover by supracrustal successions. From this evidence Rogers (1996) and Rogers and Santosh (2004) propose that Earth’s oldest surviving crust comprised three major agglomerations of ancient nuclei. The oldest, referred to as ‘Ur’, included the only large areas of crust stabilised by ∼3.0 Ga and incorporated five cratons including the Kaapvaal of southern Africa, the Pilbara of Western Australia

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and the Dharwar, Bhandar and Singhbhum cratons of India. Small exposures of ∼3.0 Ga crust along the periphery of East Antarctica together with the protolith histories of metamorphic nuclei in Zimbabwe and Madagascar suggest the possible wider extent to this grouping of cratons. Secondly, primeval crust of Archaean age underlies much of North America, Greenland, Siberia and Fennoscandia and the first three are grouped by Rogers (1996) into the protocontinent of ‘Arctica’. Although isolated relicts as old as ∼3.85 Ga are present here, major protolith formation was concentrated much later near the Archaean-Proterozoic boundary at 2.6–2.4 Ga. There is substantial geological and palaeomagnetic evidence (e.g. Buchan et al., 2000; Mertanen et al., 1999; Windley, 1993a,b; Piper, 1982, 2003) that the core of the Fennoscandian Shield was part of this division by 2.5 Ga. Subsequently Laurentia and Fennoscandia also share a common geological history including Mesoproterozoic marginal accretion and anorogenic magmatism (Bridgewater and Windley,

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Table 1 Summary of selected 2.9–2.0 Ga Palaeomagnetic poles from the constituent shields of Ur and Arctica. No.

Rock Unit

Pole position ◦

Africa, Central-Southern Cratons 1 Agatha Basalt (K) 2 Nhlebela Basalt (K) 3 Usushwana Complex (K)

N

◦

Age (Ma)

Rotated pole ◦

E

N

−9 26 9

333 358 347

2940 ± 22 2984 ± 3 2875 ± 40

−35 6 −13

9

346

(A95 = 24.7◦ )

Database number/reference ◦

E

197 206 203

AF1

Mean 1–3

−14

203

4 5 6 7 8 9 10

Westonara Basalt (K) Belingwe Komatiites (Z) Modipe Gabbro (K) Derdepoort Basalt (K) Allanridge Basalt (K) Allanridge Formation (K) Ventersdorp Lavas (K)

−17 −45 −33 −40 −75 −68 −55

48 303 31 5 0 356 355

2714 ± 8 2692 ± 9 2670–3000 2782 ± 5 2708 ± 4 2700 2699 ± 16

−5 −77 −27 −46 −62 −64 −60

270 165 266 251 307 291 265

AF2

5–10

−52

8

(A95 = 24.1◦ )

−52

267

11 12 13 14 15

Mbabane Pluton (K) Nyanzian Lavas (T) Great Dyke Satellites (Z) Great Dyke (Z) Umvimeela Dyke (Z)

20 −14 23 21 21

106 330 56 58 62

2687 ± 6 2680 ± 10 2574 ± 2 2574 ± 2 2574 ± 2

51+ −41 33+ 32+ 34+

314 195 254 258 261

AF3

Mean 13–15

22

59

(A95 = 4.2◦ )

33

258

16 17

Ongeluk Lavas (K) Mamatwan-type Ore Complex (K)

−1 −8

101 111

2222 ± 13 ∼2200

29+ 24

313 326

AF4

Mean 16,17

17 18 19 20 21 22 23 24

Transvaal System Lavas (K) Gamgara/Mapedi Formation (K) Cunene Anorthosite (C) Phalaborwa Dykes (K) Lower Swaershoek Formation (K) Vredefort Impact, mean result (K) Bushveld Main and Upper (K) Bushveld Upper Zone (K)

AF5

Mean 18–24

25 26 27 28 29

Gaberones Granite (K) Kisii Group Lavas (T) Garauja Gabbro (T) Kenya Granites (T) Post-Kavirondian Granite (T)

Australia, Pilbara and Yilgarn Cratons 1 Millindina complex 2 Nullagine Supersequence P1 3 Nullagine Supersequence P2 4 Nullagine Supersequence P4–7 5 Nullagine Supersequence P8–10 6 Black Range Dyke 7 Cajaput Dyke 8 Mt.Roe Basalt 9 Mt. Jope Basalt

−5

106

−40 2 3 36 37 23 12 16

14 82 75 45 51 42 27 32

2150–2350 2130 ± 92 1950–2200 2061 ± 0.6 2054 ± 4 2023 ± 11 2061 ± 27 2050 ± 12

20

51

(A95 = 18.5◦ )

35 7 −27 −61 41

104 345 341 30 263

2783 ± 2 2531 ± 2 2500 ± 100 2476 ± 50 2420 ± 60

−12 −41 −47 −50 −61 −32 −46 −52 −41

161 160 153 137 193 154 146 178 129

2850 ± 20 2772 ± 2 2766 ± 2 2739 ± 14 2716 ± 2 2772 ± 2 2772 ± 2 2765 ± 3 2765 ± 3

−51 −75 −74 −74 −75 −66 −69 −86 −56

–

a 9141 3431 8717 a b 3403

6216 7538 2361 2362 7553

8250 c

27

320

−41 27 25 37 41 25 13 14

258 291 284 235 239 242 216 238

27

252

303 −16 −48 −48 308

65 202 216 290 −13

3349 2505 1360 8 8122

311 273 243 228 127 282 242 148 241

308 8175 9176 9177 9178 1954 1955 311 309

AUS1

Mean 2–9

−75

241

(A95 = 11.0◦ )

−75

240

10 11

Ravensthorpe Dykes Widgiemooltha Dykes

−38 −8

136 157

2450 ± 100 2410 ± 2

−59 −46

251 307

AUS2

Mean 10, 11

−23

148

–

−56

284

12 13

Paraburdo BIF Wittenoom BIF

−41 −36

225 219

2100–2300 2100–2300

56 57

253 240

AUS3

Mean 12, 13

−39

222

–

56

246

64 46 −10 −43 −20 −15 1 −28 −18 −18

313 207 309 293 279 301 265 260 265 247

2680 ± 3 2440 ± 34 2493 ± 7 2471 ± 35 2458 ± 18 2450 ± 83 2493 ± 17 2449 ± 1 2440 ± 50 2445 ± 4

43 50 −40 −60 −35 −35 −11 −37 −29 −23

278 193 269 241 234 260 226 211 220 202

Fennoscandian Cratons 1 Varpaisjarvi Quartz-diorite 2 Taivalkaski-Syote Gabbro 3 Monchegorsk Pyroxenite 4 Generalskaya Layered Intrusion 5 Karelian Dykes 6 Voche Lambina Metamorphic rocks 7 Monchegorsk Intrusion 8 Burakok (i) Gabbro-diorite 9 North Finland Intrusions 10 Kivakka Intrusion

a a 377

2429 9019 2749 d e f g 6466

1949 1889

7539 7541

1314 5735 7423 h 8464 7410 9225 8522 6625 9226

156

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Table 1 ( Continued ) No.

Rock Unit

Pole position ◦

N

Age (Ma) ◦

Rotated pole ◦

E

N

Database number/reference ◦

E

11 12 13 14 15 16 17

Burakok Intrusion (ii) Tolstik Intrusion Main Range Gabbro Kolvitza Porphyrites North Karelia Intrusions Imanda Layered Intrusion Hautavaara Gabbro

−25 −32 −17 −29 −41 −16 −41

248 305 305 317 245 280 306

2439 ± 22 2437 ± 7 2436 ± 25 2423 ± 3 2415 ± 55 2407 ± 39 2400 ± 46

−30 −52 −37 −50 −43 −31 −61

200 262 265 279 189 236 261

F1

Mean 3–16

−25

282

(A95 = 12.7◦ )

−41

236

18 19

Akhmalahti Formation Burakok Intrusion

−60 −40

284 305

2330 ± 38 2300–2400

−73 −60

206 260

F2

Mean 18, 19

−51

297

–

−68

240

20 21 22

Girvas Dykes Matozero Sill Segozero Sill

−31 −38 −41

281 295 311

2150 ± 60 2250 ± 50 2250 ± 50

−46 −56 −62

232 246 269

F3

Mean 20–22

−37

295

(A95 = 20.0◦ )

−60

252

23 24 25 26

Kuetsyarvi Formation Iisalmi Intrusions Iisalmi Dykes Kuusamo Dykes and Greenstone

25 42 59 47

300 249 241 234

2150 ± 50 2050–2250 2050–2250 2050–2250

5 33 50 42

265 226 231 217

F4

Mean 24–26

48

242

(A95 = 13.1◦ )

40

223

−24 37 6 47 23 33 16

212 194 175 152 116 99 85

2700–3000 2700–3000 2600 ± 100 2535 ± 40 ∼2500–2100 ∼2500–2100 ∼2300–2100

65 2 20 −24 −31 −49 −47

176 167 198 186 228 229 261

24

100

(A95 = 26.0◦ )

−43

239

Indian Cratons 1 Orissa Banded Hematite BHJ1 2 Orissa Banded Hematite BHJ2 3 Khammano Quartz-Magnetite 4 Quartz-magnetite rocks 5 Vandallur Charnockites V1 6 Pallavaram Charnockites P1 7 Charnockite Belt B I1

Mean5–7

8 9 10 11

Charnockite Belt A4 Vandallur Charnockites V2 Charnockite Belt A1 Charnockite Belt A2

−18 −26 −8 −45

133 136 163 147

∼2500–2100 ∼2500–2100 ∼2500–2100 ∼2500–2100

9 17 24 36

246 250 217 261

I2

Mean 8–11

−25

145

(A95 = 19.4◦ )

22

244

12 13

Pallavaram Charnockites P2 Charnockite Belt A3

−79 −73

350 75

∼2500–2100 ∼2500–2300

39 30

332 311

I3

Mean 12,13

−70

43

–

35

321

41 13 −47 21 −68 −19 −30 −15 −47 75 7 −62 −51 −37 7 19

170 201 283 182 260 217 190 240 259 222+ 183 292 240 295 188 208

2888 ± 2 2800 ± 200 2800 ± 200 2800 ± 200 2800 ± 200 2738 ± 5 2738 ± 5 2738 ± 5 2736 ± 33 2715 ± 15 2710 ± 2 2705 ± 4 2700 ± 25 2700 ± 100 2700 ± 100 2700 ± 200

Laurentian Cratons 1 Hawk Lake Granites 2 Dundonald Sill 3 Munro Formation 4 Kamiskotia Complex 5 Archaean Rocks Ontario 6 Griffiths Mine Host A 7 Griffiths Mine Host B 8 Griffiths Mine Ore A 9 Kinojevis Tholeiites 10 Red Lake Greenstone 11 Ghost Range Intrusion 12 Stillwater Complex (W) 13 Poohbah Complex N 14 Sherman Volcanics 15 Sherman Iron Ores 16 Steep Rock Iron Ores L1 L2

Mean 1,2,4,6–8,11,15–16 Mean 3,5,9.12–14

5 −54

198 273

(A95 = 15.4◦ ) (A95 = 17.5◦ )

17 18 19 20 21 22 23 24 25

Dogrib Dykes (S) Otto Stock Moose Mountain A Moose Mountain B Adams Mine Ores External and Baldhead Granites Piwitonei Granite Shawmere Anorthosite Deformation Zone rocks

−35 69 −46 −8 −3 27 21 39 −77

310 227+ 274 194 199 195 226 224 266

2692 ± 80 2680 ± 1 2680 ± 30 2680 ± 30 2675 ± 35 2668 ± 2 2665 ± 25 2650 ± 100 2635 ± 33

7425 8524 7424 7404 7782 h 7427

7647 h

8521 7418 8518

7648 1317 1318 2776

i i J 20 k k l

L k l l

k l

6394 2151 2154 2152 1614 5860 5861 5863 5850 7145 1676 3046 1722 5851 5853 5866

2506 2900 5854 5865 5857 6393 8392 393 7146

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Table 1 ( Continued ) No.

Rock Unit

Pole position ◦

N

Age (Ma) ◦

Rotated pole ◦

E

26

Chibougamau Greenstone

−61

273

2600 ± 100

L3 L4

Mean 20–23 Mean 17,19, 24–26

10 −61

203 286

(A95 = 20.3◦ ) (A95 = 27.5◦ )

27 28 29 30 31 32 33 34

Racine Lake Tonalite Kapuskasing Structural Zone Shelley Lake Granite Burchell Lake Granite Red Lake RLG Magnetisation Archaean Gneiss Poohbah Complex R Ptarmigan Dykes

−70 −9 −78 −71 −77 −66 −78 −42

261 239 246 263 266 223 270 220

2500–2700 2500–2600 2580 ± 20 2580 ± 30 2550–2610 2550 ± 100 2550 ± 50 2505 ± 2

L5

Mean 27–33

−70

243

(A95 = 15.4◦ )

35 36 37 38

Matachewan Dykes Matachewan Dykes Matachewan Dykes Chipman Lake Diorite

−45 −53 −40 −51

241 261 216 229

2453 ± 2 2453 ± 2 2453 ± 2 2400–2500

L6

Mean 35–38

−48

236

(A95 = 17.6◦ )

39 40 41 42 43 44 45 46 47 48

Lorrain Formation Coleman Member Lorrain Formation Kaminak Dykes Tulemalu Dykes (C) Maguire Dykes Nipissing Diabase N1 Semeterre Dykes Caribou Gabbro Indin Dykes

−46 −31 −21 −24 1 −9 −15 −15 14 19

268 234 213 238 302 267 263 284 296 284

2350 ± 50 2350 ± 50 2350 ± 50 2170–2570 2255 ± 15 2229 ± 35 2219 ± 4 2216 ± 8 2176–2196 2007–2179

L7

Mean 43–48

−5

277

(A95 = 15.6◦ )

49 50 51 52 53 54

Otish Gabbro Sokoman Iron Formation Wind river Diorite (W) Big Spruce Complex Kapuskasing Dykes Dollyberry Lake Basalts

35 33 84 67 61 59

253 255 215 247 253 200

2000–2200 2100–2200 2170 ± 8 2066 ± 40 2043 ± 14 2040–2400

L8

Mean 49–54

58

244

(A95 = 19.4◦ )

55 56 57 58 59 60 61

Biscotasing Dykes Biscotasing West Dykes Superior Dykes Marathon Normal Dykes Molson Dykes C1 Fort Frances Dykes Fort Frances Dykes

28 17 27 43 53 43 51

223 225 226 196 180 184 175

2166 ± 14 2166 ± 14 2150 ± 25 2121 ± 14 2091 ± 2 2076 ± 5 2076 ± 5

L9

Mean 55–61

39

205

(A95 = 17.1◦ )

N

Database number/reference ◦

E 2415

394 9005 1716 1718 7146 6023 1723 8308

Footnote (a) Footnote (b) Footnote (c) 6669

8244 8248 8480 2168 1714 8309 Footnote (d) 7190 1174 2898

2230 2421 8994 2034 6444 1410

7189 9213 3040 6470 8548 1739 8034

Rotated poles are in a Laurentian reference frame. African poles are distinguished by craton of origin as follows—C: Congo, K: Kaapvaal, T: Tanzania, Z: Zimbabwe. Laurentian poles distinguished as S, C and W come from the Slave, Churchill and Wyoming terranes respectively; the remainder are from the Superior Craton. (a) ‘Majority’ result; mean of results 1273, 1721, 2153, 3489, 3492, 5876, 6353, 6392, 6401, 6441 (intrudes result 18), 6455 and 64563. (b) Mean of Groundhog, Hornepayne and Ogoki results 5147, 6453 and 6452. (c) Mean of results 6442, 6454 and 6672. (d) Mean of results 1599, 1639, 1666, 2620, 2804, 3071, 3417, and 6355. References to sources are GPDB results numbers of lettered: a, Strik et al. (2003); b, de Kock (2007); c, Evans et al. (2001); d, Morgan and Briden (1981); e, de Kock et al. (2006); f, Salminen et al. (2009); g, Evans et al. (2002) and Hattingh (1989); h, Arestova et al. (2002), i, Das et al. (1996), j, Poornachandra Rao et al. (1989); k, Piper et al. (2003); l, Mondal et al. (2009).

1973; Piper, 1980, 1987). Together with Proterozoic additions, this is embraced by the extended protocontinent of ‘Nena’ (Gower et al., 1990) or ‘Kenorland’ (Williams et al., 1991). Thirdly much of the continental protolith of South America and western Africa appears to have stabilised at ∼2.3–2.0 Ga following Transamazonian and Eburnian orogenic episodes. A most important line of evidence linking the constituent cratons of this assemblage is the widespread development of ∼2 Ga fluvio-deltaic sediments on five cratons (Ledru et al., 1994; Rogers, 1996). The presence of this craton-linking episode effectively excludes the possibility of later widespread dispersal and rewelding of these elements. This cratonic assemblage is referred to as ‘Atlantica’ by Rogers (1996). Preservation of these three groups of cratons within the much later Phanerozoic supercontinent of Pangaea suggests that they

have preserved a strong degree of internal quasi-integrity during their lengthy subsequent histories. This implies that younger orogenic belts separating the cratons involved limited relative motions and did not include the development of wide ocean basins. However, whilst the histories of protolith consolidation and subsequent supracrustal cover can provide strong indications that this was the case, palaeomagnetic data are required to constrain the configurations of the cratons and evaluate their wider global relationship. With this objective in mind, the present paper compiles a database of Archaean and Early Proterozoic palaeomagnetic poles and employs it to evaluate cratonic configurations during these times. An earlier interpretation (Piper, 2003) is updated by incorporating new data, and the links to crust forming processes in the Late Archaean-early Proterozoic are evaluated. The derivation of the supercontinent shape is used to speculate on the wider relation-

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Fig. 1. Reconstruction of Protopangaea, the continental crust in mid-Archaean to Early Proterozoic times, as derived from proximities deduced from geological evidence and palaeomagnetic data (see text). The Laurentian Shield of North America with Greenland rotated to North America) is retained in present day coordinates and other shields are rotated according to Eulerian operations (Euler angle positive anticlockwise, Euler pole ◦ E, ◦ N) Africa (−146◦ , 138◦ , 73.0◦ ), Australia (−161◦ , 167◦ , −69◦ ), India (−149.5◦ , 4◦ , −16.5◦ ), Fennoscandia (−42.5◦ , 211◦ , 60◦ ), Siberia (107.3◦ , 271◦ , 77◦ ). Continental crust, which had consolidated by these times and escaped later reworking, is shown in dense colour; inferred crust in process of consolidating and or subsequently reworked is shown in light colour. Continental assemblages identified as ‘Ur’. ‘Arctica’ and ‘Atlantica’ by Rogers (1996) and Rogers and Santosh (2004) are indicated but note that Atlantica had not fully consolidated by the time period considered here. The distribution of straight belts is after Watson (1973) and the mineral distributions are generalised after Petrasheck (1973), Watson (1978), Berry (1980) and de Wit et al. (1999). Note the concentration of gold and economic pegmatites within the oldest protocontinent of Ur with a second belt of gold emplacement running through Atlantica, whilst the broad zones of Tin and Tungsten are exclusive to Atlantica and comprise broad belts conforming to the form of the supercontinent.

ship of the continental crust to whole-mantle processes operating during this early part of Earth history. 2. The palaeomagnetic data Table 1 summarises palaeomagnetic results assigned to the interval ∼2.9–2.0 Ma with data selected to preserve a testable dataset whilst constrained to four criteria: (i) demagnetisation applied, (ii) A95 ≤ 16◦ , (iii) interpreted as primary by the authors, and (iv) ages of the poles estimated to ±200 Ma or better by stratigraphic or isotopic evidence (but see an exception with the Indian Shield data noted below); poles dated by magnetic correlation are excluded. Unfortunately it is not possible to apply the common criterion N (number of samples) ≥ 24 and leave a substantial dataset from the Fennoscandian Shield where a high proportion of results are based on small numbers of samples; nevertheless the poles from this shield are internally consistent and well grouped and the derivative studies have a minimum N of 13. The data of Table 1 are mostly compiled from the Global Palaeomagnetic Database (GPDB) where they are referred to their result numbers; some new data or results not included in the GPDB are incorporated in the compilation and referenced accordingly. In this short review it is not possible to note specific reservations relevant to individual results (mostly relating to the age assignments). Within Laurentia and Fennoscandia a specific challenge is the isolation of results which have not been overprinted by subsequent widespread Palaeo-Mesoproterozoic Hudsonian and Svecofennian orogenic events; important recent reviews including new data from southern Africa are included in Strik et al. (2007) and Salminen et al. (2009). In spite of the great antiquity of the source rocks, the compilation provides a robust dataset for interpretation. Thus 76% of the 143 poles are derived from igneous rocks linked to radiometric age dating on the same or related units. The value

of the remaining fraction derived mostly from metamorphic rocks or ore deposits is of more limited value; this is specifically because the depressed temperature of remanence acquisition during protracted cooling makes age assignments correspondingly tenuous. The bulk of the poles from India (5–13) for example, come from exhumed granulite-charnockite terranes and, although this cooling was clearly Early Proterozoic in age, no specific assignments within the general range ∼2.5–2.0 Ga are possible. Important reservations also apply to the small (9%) fraction of data from sedimentary units including banded ironstone formations. The complexities of the magnetic record in sediments of this great age are illustrated for example, by study of the Huronian Supergroup where Williams and Schmidt (1997) consider that just two of six components identified are interpretable as quasi-primary. Limitations of the sedimentary and metamorphic results do not however, influence the essential conclusions drawn here due to their relatively small contribution to the database and their general conformity with the igneous data; the exception is the case of the Indian cratons where they comprise the total evidence. To simplify inter-cratonic correlations Table 1 is organised into successive pole position–age groupings; mean pole calculations are shown in bold where common age and positional groups are recognisable. 3. The Late Archaean-Palaeoproterozoic protocontinent The continental reconstruction developed here is based on two criteria. The first is the evident ancient proximity of cratons on the basis of similar geological histories (Rogers, 1996; Rogers and Santosh, 2004) and the second is the constraint provided by the palaeomagnetic poles, and specifically by long-term similarities in the palaeomagnetic records of the five major shields for which a test is possible during the Late Archaean-Palaeoproterozoic interval under consideration. A surprising finding of Precambrian palaeo-

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Fig. 2. Contoured distribution of all palaeomagnetic poles assigned to the interval 2.9-2.2 Ga from the shields of Africa, Australia, Fennoscandia, India and Laurentia (Table 1) plotted following rotation into the reconstruction of Protopangaea; rotation operations are given in the caption to Fig. 1.

magnetism has been that a single rotational operation, namely clockwise rotation of 146◦ about a Eulerian pole at 73◦ N, 138◦ E repeatedly brings poles positions from Laurentia and centralsouthern Africa into agreement over an interval of more than 2 Ga; this accordance over such a long interval of time without the need to adjust the primary rotational operation even marginally is a powerful indication that the heart of the continental crust was coherent on a palaeomagnetic scale (Piper, 1982, 2003, 2007). This point is reinforced by the correlation of palaeomagnetic poles here (Figs. 2 and 3): not only are the bulk of the poles from the two regions brought together by this single rotational operation but, as noted below, two key polar trends towards either end of the investigated time interval are recognised within the two datasets. The Kaapvaal Craton of South Africa and the Pilbara Craton of Western Australia are elements of the earliest recognisable protocontinent ‘Vaalbara’ which is identified by impressive lithostratigraphic and chronostratigraphic correlations of 3.6–2.1 Ga granite–greenstone rocks (Cheney, 1996; Sinomson et al., 2009). Palaeomagnetic data supports close proximity of these nuclei by 2.87 Ga (Zegers et al., 1998), with newer evidence indicating that this correlation can be extended back to at least ∼3.5 Ga (Usui et al., 2009). Proximity of these nuclei by Late Archaean times is accepted here using a position with Pilbara sited south rather than north of the Kaapvaal (Fig. 1); this position is permissible within the palaeomagnetic constraints of the above authors and brings the NNE–NE lineaments in the Kaapvaal into broad continuity with N–S to ENE–WSW trending lineaments in the Kaapvaal. It is favoured by three further observations namely (i) terrane accretion at 2.65–2.0 Ga excludes a position along the western margin of the Kaapvaal Craton (de Wit et al., 1992), (ii) consolidation of crust to the north of the Kaapvaal into the Limpopo Belt (cf. Zegers et al., 1998) had welded this craton with the Zimbabwe Craton by ∼2.6 Ga (Windley, 1993a,b), and (iii) the APW link between ∼2.45 and ∼2.2 Ga poles from Australia with those from other cratons is implied by this configuration (Fig. 2 and see poles AUS1-3 in Table 1). The position of India within the Ur assemblage is constrained by an APW swathe defined by the range of Early Proterozoic charnockite poles (Mondal et al., 2009) and by their relationship to two ∼2.6 Ga poles from the Singhbhum Craton in the NE sector of the shield (Figs. 1 and 2); these data locate this shield immediate below Western Australia and Southern Africa. Although now separated

by a rift zone active during Phanerozoic times, the Dharwar and Singhbhum cratons have comparable Archaean histories; similarities between the Bangemall Basin of Western Australia and the Cuddapah Basin of eastern India also provide a case for postulating proximity of these two blocks since Palaeoproterozoic times (Prasad et al., 1987). Hence collectively the Ur nucleus seems to have comprised an alignment of cratons (Fig. 1) embracing at least the Archaean nuclei of India, Australia and southern Africa. A few south-central Africa poles (6, 19, and 26–28) from outside of the Kaapvall-Zimbabwe core suggest a wider inclusion of cratons to the west and north within this group. Whilst the significance of this palaeomagnetic correlation is marginal, the extraordinarily long duration of the palaeomagnetic correlation between Ur and Arctica (Piper, 1982, 2007) must have implications to the integrity of the continental crust in between. It indicates that mobile belts in between such as the Ubendian, Irumide and Lufilian-Zambezi of Mesoproterozoic and Pan-African ages are essentially intracratonic. Broad structural correlations of Mesoproterozoic age cross the younger orogenic belts and also seem to mitigate against large later dislocations (Shackleton, 1973) although a minimum Pan African motion comprised sinistral strike slip (Hanson et al., 1993) on a scale which is unlikely to be detectable by palaeomagnetism; possible oceanic sutures occur to the east and south east of the collective cratonic assemblage (Daly, 1986). The exceptions to the correlations noted above are one pole (25) which does not conform to poles from elsewhere and poles 13–15 from the Great Dyke and satellites of the Zimbabwe Craton (mean AF3) which plot with poles 400 Myr younger (Fig. 3). Poles in the ∼2.9–2.7 and 2.7–2.6 Ga age divisions from Laurentia fall into two groups (summarised by L1–L3 and L2–L4 in Table 1). These are internally consistent but approximately 90◦ apart. Since they are derived from the same shield and all but two results come from the Superior Craton, this difference cannot have a tectonic explanation; also, there is no clear age distinction between the two groups so APW motion does not obviously apply here. The same difference is present between African poles AF1 and AF2 and is also suggested by Indian poles 2–4 and mean pole I1 all with assigned ages between ∼2.9 and 2.6 Ga (Fig. 3). Since these poles from three different nuclei are separated by ∼90◦ an alternative explanation is that they record inertial interchange true polar wander (IITPW). True polar wander (TPW) is caused by migration of the geographic reference frame relative to the dynamically conserved

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Fig. 3. Mean palaeomagnetic poles from Africa, Australia, Fennoscandia, India and Laurentia assigned to the interval 2.9–2.2 Ga (Table 1) plotted following rotation operations given in the caption of Fig. 1 into the reconstruction of Protopangaea with assigned mean ages. A generalised interpretation of apparent polar wander (APW) during mid-Archaean to Early Proterozoic times is also indicated (see text).

angular momentum vector (the spin axis). It can be a response to the redistribution of mass within the Earth caused by mantle perturbations (see Phillips et al., 2009). The rising and sinking of mantle density variations can cause deflection in discontinuities within the planetary interior and perturb the Earth’s inertia tensor. A potential consequence is the wholesale rotation of the solid Earth with respect to the spin axis until the maximum principal inertia axis (Imax ) is aligned with the spin axis (Evans, 1998). Positive mass anomalies represented by the principal axes Iint and Imin are driven towards the equator (Goldreich and Toomre, 1969). Thus, if mantle density heterogeneities change the relative magnitudes rather than the orientations of the principal inertial axes, then Imax can potentially fall below Iint so that the entire tectosphere rotates around Imin to align the new Imax with the spin axis. The signature of such a motion is a rotation of ∼90◦ confined to a circle at right angles to the Imin axis (although differing amounts of rotation could result depending on the character of mantle convection and the internal density structure). Several instances of rapid ∼90◦ APW motions have been attributed to IITPW in Precambrian (Evans, 1998) and Palaeozoic (Piper, 2006) times and might be expected to have occurred more readily in a thermally active Archaean-early Proterozoic Earth. Tsai and Stevenson (2007) estimate that the present day mantle viscosity structure limits the speed of IITPW to no more than ∼8◦ over 10 Myr although rates are likely to have been faster when lower viscosities pertained in Archaean times. The core link between the Laurentia and Fennoscandia shields and the Ur cratonic group shown in Fig. 1 and used to reconstruct the poles in Figs. 2 and 3 is also embraced by Neoproterozoic data (Piper, 2003, 2007). A corollary of this observation is that the combined shield area in between, although apparently not fully cratonised during the time period reviewed here, has remained coherent on a scale resolvable by palaeomagnetism. In general terms the wide spatial separation of the Ur and Arctica cratonic groupings is suggested by totally distinct geological

histories until the widespread deposition of glaciogenic rocks at 2.4–2.2 Ga (Aspler and Chiarenzelli, 1998). The Fennoscandian data of Table 1 are derived mainly from study of igneous episodes linked to superplume events at ∼2.45–2.25 Ga (Abbott and Isley, 2002; Buchan et al., 2000) and approximately equal numbers of poles come from Kola and Karelia cratonic nuclei separated by the Lapland-Belomorian Belt. Although this latter orogen includes a 1.97–1.87 Ga volcanic arc suite interpreted as the site of a collisional suture by Berthelsen and Marker (1986), there is no clear distinction between them on a palaeomagnetic scale and the short swathe of ∼2.45–2.25 Ga Fennoscandian poles (F1–F3) correlates with Laurentian poles L2, L4–L6 (Fig. 3) with mean ages of 2.56–2.36 Ma. A similar palaeomagnetic correlation has emerged from several independent analyses (Mertanen et al., 1999; Piper, 1982, 2003; Buchan et al., 2000; Pesonen et al., 2003) and implies that the KolaKarelia nucleus was in broad continuity with the Archaean terranes of Labrador and Greenland. It further supports the view that a single large igneous province embraced the Superior, Kola and Karelian cratons early in Palaeoproterozoic times (∼2.45 Ga, Heaman, 1997). Contrasting interpretations have sited Siberia either to the north (Condie and Rosen, 1994) or to the west (Sears and Price, 1978,2000) of Laurentia during the Proterozoic but cannot be addressed here due to absence of data from Siberia; the Sears and Price orientation is however, endorsed by Neoproterozoic data (Piper, 2007) and 2.9–2.2 Ga poles from Australia come from steep inclination magnetisations and plot close to this nucleus so that they cannot be matched with the bulk of the remaining data in Fig. 2 by locating Australia in the conventional Rodinian position west of North America. The Sears and Price (2000) configuration is therefore adopted in Fig. 1 and achieves broad continuity of continental crust between Laurentia-Fennoscandia and the area of Atlantica protocontinent (still consolidating during these times). An implication of this assumption is that Arctica comprised a second linear grouping of cratons embracing Fennoscandia, Laurentia and Siberia. The adopted palaeomagnetic reconstruction is strongly endorsed by the collective grouping of ∼2.7–2.2 Ga poles from 5 shields into a quasi-static grouping close to the periphery of Ur (Figs. 2 and 3). In this configuration the subsidiary grouping of ∼2.8–2.6 Ga poles from Laurentia (L1, L3) is also supported by poles from India and the Kaapvaal. In contrast, poles from all five shields falling in the younger age range of the dataset (∼2.2–2.0 Ga) plot far away from this ∼2.7–2.2 Ga quasi-static group and in the vicinity of Laurentia-Siberia (Fig. 2). The south to north APW movement to these latter younger poles is recognised in all five cratonic groups (Fig. 3, AF2 → AF5 in Africa, AUS2 → AUS3 in Australia, I1 → I2 in India, F3 → F4 in Fennoscandia and L6 → L7 → L8 in Laurentia where it moves specifically to the key ∼2.22 Ga Nipissing pole). Hence the analysis of Figs. 2 and 3 identifies an interval comparable in length to the whole of the Phanerozoic eon, when the continental crust exhibited little of no movement with respect to the poles, which was succeeded by a rapid APW movement at ∼2.2 Ga carrying the path across the supercontinent. It seems unlikely that the continental crust ever subsequently became quasi-static with respect to the poles: there is no comparable grouping of ∼2.2–2.0 Ga poles and there is some indication that APW motion crossed Laurentia and moved towards the east (AF4, I3, Figs. 2 and 3) late in the time interval under consideration here.

4. Implications to crustal extraction, rate of continental movement and tectonic style The observation that the shields moved only slightly with respect to the geomagnetic poles over a time interval of some 500 Myr (Figs. 2 and 3) has clearly facilitated the derivation of the primary reconstruction and the contrast with the continuous

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and sometimes rapid APW during subsequent Proterozoic times (e.g. Pesonen et al., 2003; Piper, 2007) is a significant signature of contrasting tectonic regimes. It correlates temporally with the transition from the granite–greenstone belt tectonics characterising the Archaean and Early Proterozoic to the formation of long linear tectono-thermal mobile belts subsequently during the Proterozoic eon (Piper, 1987; Windley, 1993a,b; Condie, 1997). Only relatively small areas of the supercontinent as defined here have escaped subsequent tectonic-thermal reworking and it remains unclear how much of the continental area suggested in Fig. 1, notably within Atlantica, was either still consolidating or not yet extracted from the mantle. The two envelopes surrounding the stabilised cratons in this reconstruction suggest a minimum area of the continental crust formed by Late Archaean and Early Palaeoproterozoic times amounting to approximately 35% of the present crustal area. A calculation of the minimum crustal volume extracted by ∼2.2 Ga must incorporate changes in crustal erosion and crustal recycling since the Archaean, plus estimates of the amount of trapped and congealed melt (Abbott et al., 2000). There are two approaches: the first sums the volume of seismic crust, eroded crust and trapped melt in the lithospheric mantle. The second assumes that the lithosphere root is the residue of a melt column and yields an estimate some 20% higher. Employing the recent estimates of crustal thickness prior to 2.4 Ga (Abbott et al., 2000), the reconstruction of Fig. 1 implies that a minimum of between 35 ± 6% and 56 ± 8% of the continental crust had been extracted from the mantle by the end of the time interval considered here; these estimates compare with figures of between 48% and 72% deduced by Abbott et al. (2000) at ∼2.2 Ga and are broadly consistent with geochemical models for continuing continental growth but with some recycling of early extracted crust (McClennan and Taylor, 1982). Important uncertainties in these calculations relate to the Ukrainian Shield, which includes Archaean crust but was probably not in its present configuration in continuity with Fennoscandia before 1.8 Ga (Elming et al., 2001) and the constituent cratons of South American and West African Shields. The latter are identified as mostly Archaean in origin by Abbott et al. (2000) but considered by Rogers (1996) to be still consolidating until shortly before uplift and cover by shallow fluvio-deltaic sedimentation at ∼2.1 Ga.

5. Why do supercontinents tend to form elongate and crescent-shaped bodies? The subject of pre-Phanerozoic supercontinent reconstructions has been a fertile field of discussion in recent years (e.g. Rino et al., 2008; Cordani et al., 2009; Eriksson et al., 2009) with particular focus on the role of mantle dynamics in the formation and disruption of supercontinents (e.g. Santosh et al., 2009; Senshu et al., 2009). However, a key question that is rarely addressed (although implicit in all of them) is: ‘how did the mantle produce the synergy of conditions responsible for the supercontinent shape?’ If the key premises used here to reconstruct Protopangaea in mid-Archaean to Early Proterozoic times are broadly correct (Fig. 1) a solution is derived which is remarkably similar in form to the much younger supercontinent comprising Wegener’s Pangaea. This observation enables some evaluation of this question. The high degrees of symmetry illustrated by Phanerozoic Pangaea were highlighted by Le Pichon and Huchon (1984) who also noted that it was symmetrically disposed about the poles of rotation as defined by ∼160–300 Ma palaeomagnetic poles (i.e. embracing most of the lifetime of the supercontinent). They linked this configuration to the longest wavelength component of the present day geoid. The geoid is the equipotential surface of the Earth’s gravity field and represents the best fit to global mean sea level. It is dominated by

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Fig. 4. The continental reconstruction of Protopangaea rotated to fit the present day hydrostatic geoid (version EGM96) centred at longitude 0◦ E and illustrated on a Lambert Equal Area projection showing the whole globe. The signature of the present day subduction zones on this version of the geoid is shown in lighter shading to highlight the dominant component of the “tennis ball” form of order 2. The mean 2.7–2.2 Ga palaeomagnetic pole defined by the contoured grouping in Fig. 2 and the mean poles in Fig. 3 is moved to the present south geographic pole of rotation. Within this constraint it is possible to move the continental form of Protopangaea into a geoid negative zone. This suggests that the continental crust formed during the Archaean was concentrated into Earth’s first supercontinent of Protopangaea as large scale, probably whole-mantle, convection moved the crust into regions of minimum gravitational potential.

low-order symmetry of “tennis-ball” shape linked to density variations within the deep mantle (Gough, 1977) and evidently caused by whole mantle convection (Crough and Jurdy, 1980). The present geoid has hemispheric symmetry with an axis located in the equatorial plane and a positive belt lying preferentially along the equator to minimise kinetic energy in the Earth’s rotation (Goldreich and Toomre, 1969). Whilst details are contentious, the modern acceptance that lithosphere slabs can penetrate mantle layers to descend all the way to the core-mantle boundary (e.g. Maruyama et al., 2007; Zhao, 2009), and the parallels between the geoid form and Pangaean shape, support the view that whole mantle convection operates, at least episodically, to aggregate the continental crust into bodies with specific geometrical properties. According to Busse (1983) the present equatorial positive belt is the ascending cell of a mantle wide-convection cell and the polar negative belt is the descending limb. A consequence is that continental crust, probably due to thick lithosphere roots, is governed by deep geoid-related convection whereas plates covered mostly, or entirely, by oceanic crust show no such relationship because they are driven by shallow convection powered by ridge-push and slab-pull forces. Hence most present day continental blocks are moving towards the geoid negative zone. This association appears to have prevailed during the history of Phanerozoic Pangaea (Le Pichon and Huchon, 1984). To test whether Protopangaea might have grown by the movement of continental crust towards regions of minimum gravitational potential controlled by the onset of whole mantle convection, the ∼2.7–2.2 Ga quasi-static pole position (Fig. 3) is moved to the present day rotation axis and superimposed onto geoid EGM96 in Fig. 4. Whilst the positive signature of circumPacific deep subduction remains in this version of the hydrostatic geoid, the dominant tennis ball spherical harmonic of order 2 is

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Fig. 5. (a) Distribution of magnetic inclination |I| in the 2.9–2.0 Ga database compared with the distributions predicted from random sampling of a geocentric axial dipole and derived from the total Precambrian database after Kent and Smethurst (1998). (b) The distribution of magnetic inclination |I| in the 2900–2200 Ma database of Table 1 compared with the distribution predicted from the geoid constrained continental model of Fig. 4.

clear. The Protopangaean crust can be sited entirely within the latitudinal embrace of a geoid negative zone. This is a tantalising (albeit tentative) indication that the mantle processes that aggregated continental crust during the early history of the Earth were comparable to those during the Phanerozoic. The increase in mantle viscosity and reduction in Rayleigh Number accompanying a decline in radiogenic heat production during the Archaean would have operated to change the pattern of convection from small scale hexagonal, to bimodal, to longitudinal rolls (Richter and Parsons, 1976). The stresses produced by this convection were up to an order of magnitude larger than stresses produced by present day convection (McKenzie and Weiss, 1975) and evidently capable of disrupting the crust to produce the granite–greenstone belts. Thus the progressive transfer of heat release in the form of small scale convection responsible for granite–greenstone tectonics to large scale convection carrying the crust towards gravity minima with a response seen in the development of the long linear mobile belts of Meso-Neoproterozoic times was evidently a key factor in the formation of the Protopangaea shown in Fig. 1. Two isotopic indicators of corresponding mantle depletion (Nb/Th and 4 He/3 He) both decline rapidly from peaks at 2.5–2.0 Ga until after 1.0 Ga (Silver and Behn, 2008). This collapse correlates with a minimum in passive margin development and major emplacement of the anorogenic granite–anorthosite suite (Bridgewater and Windley, 1973). The spatial (mostly in Palaeoproterozoic crust) and temporal (mostly Mesoproterozoic) concentration of this latter suite is difficult to explain in terms of conventional Plate Tectonics and appears to be the consequence of prolonged thermal blanketing of the mantle beneath a long-lived supercontinent (Anderson and Morrison, 2005). These signatures were evidently a direct consequence of the long duration of the Ur-Atlantica correlation. 6. Geomagnetic implications The palaeomagnetic evidence used to reconstruct the configurations of the cratons is the record of a stable and reversing geomagnetic field already established by ∼3.5 Ga (Biggin et al., 2008; Usui et al., 2009). It is less clear that this field conformed to the properties of geocentric axial dipole (GAD) model which accommodates more than 90% of the present day field. This is specifically because the field source is influenced by the Earth’s inner core (Hollerbach and Jones, 1993) which is unlikely to have been present in more than incipient form during the time period considered here (Buffet et al., 1991). Between 58 and 78% of the magnetisations from Africa, Australia, Fennoscandia and Laurentia yield pole positions plotting in the continental hemisphere of Fig. 1 rather than in the opposite oceanic hemisphere

although this marginal bias in just 143 results (Table 1) spanning a period as long as the Phanerozoic, cannot be more than suggestive of field properties. The frequency of polarity inversion is more informative: just 23% of these results show mixed polarities suggesting that polarity reversal was much less frequent during these times. The same observation emerges from the later Proterozoic record (Piper, 1987; Coe and Glatzmeiser, 2006; Aubert et al., 2009) and is the probable signature of a developing inner core. The frequency distribution of palaeomagnetic inclinations over long intervals of geological time can be used to test the dipolar nature of the geomagnetic field provided that the sampling sites are a random sample of the field (Evans, 1976). The distributions for Palaeozoic and Precambrian times show critical departures from the GAD model (Kent and Smethurst, 1998) that have a variety of explanations including the possible contribution of an axial octupole source, late nucleation of the inner core or biasing of continental crust into low latitudes. In Fig. 5a the magnetic inclinations (|I|) of the results summarised in Table 1 are grouped in 10◦ bins to produce a frequency distribution which shows the same trend as that predicted from the GAD model. This dataset does not possess the anomalous bias towards low palaeolatitudes which characterises the later Precambrian and Palaeozoic distributions (Fig. 5a, Kent and Smethurst, 1998). Hence the high frequency of low inclinations in the later timeframe would not appear to be the result of a long term secular change in the geomagnetic field source. Instead it must be due to a limited temporal complexity in the field source or a preferential biasing of continental crust into low latitudes by geotectonic processes. Although the later Precambrian dataset contains a much higher proportion of results from sedimentary rocks, this is evidently not the cause because the anomaly persists in separated datasets of sedimentary and crystalline origin (Kent and Smethurst, 1998). In practice, the mid-Archaean to Early Palaeoproterozoic dataset considered here is clearly not a satisfactory random sample of the ancient field since the strong grouping of ∼2.7–2.2 Ga poles in Figs. 2 and 3 shows that the Protopangaea was preferentially constrained on the Earth’s surface, and probably into the geoid negative form as noted in Section 5. Much of the crust within Ur and Arctica, which had consolidated by these times, lay for protracted periods in intermediate to high palaeolatitudes to produce the predicted |I| distribution shown in Fig. 5b for a static supercontinent constrained as shown in Fig. 4. The departure of the Archaean-Early Protoerozoic |I| distribution from the predictions of random sampling a GAD (cf. Fig. 5a and b) are therefore due to preferential positioning of continental crust on the globe. Nevertheless the contrast with the opposite bias observed later in Precambrian time remains of particular interest: it suggests that explanations for departures of these

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distributions from the GAD are dominated by effects that tend to cluster continental crust into specific regions of the globe. As such they are a potential signature of geotectonic processes. • 7. Metallogenic provinces and ‘straight belts’ There were spikes in gold deposition at ∼2.7 Ga and in volcanichosted massive sulphide deposition at ∼2.6 Ga and ∼1.7 Ga which correlate temporally with peaks in juvenile continental crust production and mantle depletion (Groves et al., 2005). The former are embraced by the interval considered here. Whilst the heterogeneous distribution of metallogenesis in the crust has usefully been linked to Plate Tectonic processes from several decades of study, it has become apparent that these distributions frequently reflect the products of older, and often Archaean and Palaeoproterozoic, deposition so that they can provide key tracers of continental growth (de Wit et al., 1999). One of the most striking is the concentration of tin and tungsten within long zones running through Atlantica (Fig. 1) but absent elsewhere (de Wit et al., 1999). The eastern belt lies mostly within reworked Pan Africa terranes whilst the western belt is within more ancient shields. Although the nature of reworking of Sn and W ores is not understood, the extent of these belts and their broad conformity with the arcuate form of Protopangaea indicates that they had their origin in Palaeoproterozoic crustal accretion. In contrast gold and economic pegmatites are largely restricted to an arcuate belt of Early Archaean crust running through Ur; a second gold belt parallels the Sn-W belts whilst a zone of lesser importance is hinted at by the alignment of similar deposits in the younger crust of Atlantica (Fig. 1). Although the gold often originated in concordant lodes in volcanic succession, remobilisation has re-deposited it as discordant lodes and ultimately as placers. Again the broad axial trend indicates an origin in mantle extraction early in the history of Protopangaea. Other distinctive tracers already established by mid-Archaean times include the regional distributions of nickel beneath the Laurentian and Fennoscandian shields and chromium beneath southern Africa (Watson, 1978). In each of these examples the mineral distribution is intrinsically related to the reconstruction of Protopangaea, and specifically to the configuration of Ur, Atlantica and Arctica (Fig. 1). Whilst the relict Archaean crust tends to exhibit only weak anisotropy over much of its preserved area, volatile transfer accompanied distributed strike slip motions and produced zones of strong ductile transpression or ‘straight belts’ (Watson, 1973). These comprise the oldest pure lineaments in the continental crust where Archaean structures are deflected into broad zones of intense deformation. They are observed to be sub-parallel to the axial trend of the protocontinent (Fig. 1; Watson, 1973; Windley, 1999). 8. Conclusions • The Late Archaean-Early Proterozoic continental crust is defined by protolith histories which define three distinct groups of cratons (Ur, Arctica and Atlantica). The 2.9–2.0 Ga palaeomagnetic data from Ur and Arctic indicate that they occupied opposite limbs of an arcuate protocontinent (‘Protopangaea’). • The palaeomagnetic correlation between these two groups is reinforced by two further observations: firstly a ∼90◦ shift is present between ∼2.9 and 2.6 Ga poles from Laurentia, Africa and India and poles which might record true polar wander. Secondly a rapid APW shift is observed in African, Australian, Fennoscandian, Indian and Laurentian shields at ∼2.2–2.1 Ga between the ∼2.7–2.2 Ga quasi-static grouping and ∼2.2–2.0 Ga poles. • By analogy with the lowest order component of the present day geoid, the symmetrical crescent shape adopted by superconti-

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nents throughout geological time appears to be due to large scale, probably whole, mantle convection carrying the continental crust towards zones of lowest gravitation potential. The quasi-static grouping of 2.7–2.2 Ga poles show that little or no movement of the continental crust with respect to the poles occurred during this long interval. This correlates with younger granite–greenstone tectonism and is interpreted to be the signature of dominant small scale convection in the upper mantle. Surviving areas of Archaean-Early Proterozoic crust suggest that 35–56% of the present crust had been extracted from the mantle by these times in broad agreement with arguments based on geochemical data. The long-term duration of the palaeomagnetic link between Ur and Atlantica has implications to the integrity of the crust in between and provides a general explanation for strong indications of continentality during the Proterozoic eon. The distribution of magnetic inclination in the 2.9–2.2 Ga palaeomagnetic data is consistent with a dominant geocentric dipole during this interval. The low frequency of reversals within the database may be a signature of little or no inner core during these times. Late Archaean-Early Proterozoic zones of strong ductile strike slip and the metallogenic provinces originating during these times show a prominent axial alignment evidently reflecting the accretionary processes that operated to aggregate the oldest surviving crust into the first supercontinent.

Acknowledgements I am grateful to Kay Lancaster for drafting the figures, Professor M. Santosh for careful editorial advice and Andrew Biggin and two anonymous reviewers for critical suggestions that materially improved the manuscript. References Abbott, D.A., Isley, A.E., 2002. The intensity, occurrence and duration of superplume events and eras over geological time. Journal of Geodynamics 34, 265–307. Abbott, D.A., Sparks, D., Herzberg, C., Mooney, W., Nikishin, A., Zhang, Y.S., 2000. Quantifying Precambrian crustal extraction: the root is the answer. Tectonophysics 322, 163–190. Anderson, J.L, Morrison, J., 2005. Ilmenite, magnetite, and peraluminous Mesoproterozoic anorogenic granites of Laurentia and Baltica. Lithos 80, 45–60. Arestova, N.A., Khramov, A.N., Guskova, E.G., Losifidi, A.G., 2002. New palaeomagnetic evidence from the Early Proterozoic (2.5–2.4 Ga) Mount Generalskaya and Imandra layered intrusions, Kola Peninsula. Fizika Zemli 38, 65–78. Aspler, L.B., Chiarenzelli, J.R., 1998. Two Neoarchaean supercontinents? Evidence from the Palaeoproterozoic. Sedimentary Geology 120, 75–104. Aubert, J., Labrosse, S., Poitou, C., 2009. Modelling the palaeoevolution of the geodynamo. Geophysical Journal International. doi:10.1111/j.1365245X2009.04361.x. Berry, G.R., 1980. World Atlas of Geology and Mineral Deposits. Mining Journal Books, London, 110 pp. Berthelsen, A., Marker, M., 1986. Tectonics of the Kola collision suture and adjacent Archaean and Early Proterozoic terrains in the northeastern region of the Baltic Shield. Tectonophysics 126, 31–55. Biggin, A.J., Strik, G.H.M.A., Langeris, C.G., 2008. Evidence for a very-long-term-trend in geomagnetic secular variation. Nature. doi:10.1038/ngeo181. Bridgewater, D., Windley, B.F., 1973. Anorthosites, post-orogenic granites, acid volcanic rocks and crustal development in the North Atlantic Shield during the Proterozoic. In: Lister, L.A. (Ed.), Symposium on Granites and Related Rocks, vol. 3. Geological Society of South Africa Special Publication, pp. 307–318. Buchan, K.L., Mertanen, S., Park, R.G., Pesonen, L.J., Elming, S.-A., Abrahamsen, N., Bylund, G., 2000. Comparing the drift of Laurentia and Baltica in the Proterozoic: the importance of key palaeomagnetic poles. Tectonophysics 319, 167–198. Buffet, B.A., Huppert, H.E., Lister, J.R., Woods, A.W., 1991. Analytical model for solidification of the Earth’s core. Nature 356, 329. Busse, F.H., 1983. Quadrupole convection in the lower mantle? Geophysical Research Letters 10, 285–288. Cheney, E.S., 1996. Sequence stratigraphy and plate tectonic significance of the Transvaal succession of southern Africa and its equivalent in Western Australia. Precambrian Research 79, 3–24. Coe, R.S., Glatzmeiser, G.A., 2006. Symmetry and stability of the geomagnetic field. Geophysical Research Letters 33, L21311. doi:10.1029/2006GL027903. Condie, K.C., 1997. Plate Tectonics and Crustal Evolution, 4th ed. Heinemann, Butterworth, 282 pp.

164

J.D.A. Piper / Journal of Geodynamics 50 (2010) 154–165

Condie, K.C., Rosen, O.M., 1994. Laurentia-Siberia connection revisited. Geology 22, 168–170. Cordani, U.G., Teixeira, W.T., D’Agrella-Filho, M.S., Trindade, R.I., 2009. The position of the Amazonian Craton in supercontinents. Gondwana Research 15, 396–407. Crough, S.T., Jurdy, D.M., 1980. Subducted lithosphere, hot spots and the geoid. Earth and Planetary Science Letters 48, 15–22. Daly, M.C, 1986. The intracratonic Irumide Belt of Zambia and its bearing on collision orogeny during the Proterozoic of Africa. In: Coward, M.P., Ries, A.I. (Eds.), Collision Tectonics, vol. 19. Geological Society of London Special Publication, pp. 321–328. Das, A.K., Piper, J.D.A., Mallik, S.B., Sherwood, G.J., 1996. Palaeomagnetic study of Archaean Banded hematite jasper rocks from the Singhbhum-Orissa Craton, India. Precambrian Research 80, 193–204. de Kock, M.O., 2007. Paleomagnetism of selected Neoarchean-Paleoproterozoic cover sequences on the Kaapvaal craton and implications for Vaalbara. Unpublished Ph.D. Thesis. University of Johannesburg, Johannesburg, 276 pp. de Kock, M.O., Evans, D.A.D., Dorland, H.C., Beukes, N.J., Gutzmer, J., 2006. Paleomagnetism of the lower two unconformity-bounded sequences of the Waterberg Group, South Africa: towards a better-defined apparent polar wander for the Proterozoic Kaapvaal Craton. South African Journal of Geology 109, 157– 182. de Wit, M.J., Roering, C., Hart, R.J., Armstrong, R.A., de Ronde, C.E.J., Green, R.W.E., Tedoux, M., Perberdy, E., Hart, R.A., 1992. Formation of an Archaean continent. Nature 357, 553–562. de Wit, M.J., Thiart, C., Doucouré, M., Wilsher, W., 1999. Scent of a supercontinent: Gondwana’s ores as chemical tracers—tin, tungsten and the Neoproterozoic Laurentia-Gondwana connection. Journal of African Earth Sciences 28, 35–51. Elming, S.-A, Mikhailova, N.P., Kravchenko, S., 2001. Palaeomagnetism of Proterozoic rocks from the Ukrainian Shield: new tectonic reconstructions of the Ukrainian and Fennoscandian shields. Tectonophysics 339, 19–38. Eriksson, P.G., Banerjee, S., Nelson, D.R., Rigby, M.J., Catuneanu, O., Sarkar, S., Roberts, R.J., Ruban, D., Mtimkulu, M.N., Raju, S., 2009. A Kaapvaal craton debate: nucleus of an early small supercontinent or affected by an enhanced accretion event? Gondwana Research 15, 354–372. Evans, M.E., 1976. Test of the dipolar nature of the geomagnetic field throughout Phanerozoic time. Nature 262, 676–678. Evans, D.A.D., 1998. True polar wander, a supercontinental legacy. Earth Planetary Science Letters 157, 1–8. Evans, D.A.D., Gutzmer, J., Beukes, N.J., Kirschvink, J.L., 2001. Paleomagnetic constraints on ages of mineralization in the Kalahari Manganese Field, South Africa. Economic Geology 96, 621–631. Evans, D.A.D., Beukes, N.J., Kirschvink, J.L., 2002. Paleomagnetism of a lateritic paleoweathering horizon and overlying Paleoproterozoic redbeds from South Africa: implications for the Kaapvaal apparent polar wander path and a confirmation of atmospheric oxygen enrichment. Journal of Geophysical Research 107 2326. doi:10.1029/2001JB000432. Goldreich, P., Toomre, A., 1969. Some remarks on polar wandering. Journal of Geophysical Research 74, 2555–2567. Gough, D.I., 1977. The geoid and single cell mantle convection. Earth Planetary Science Letters 34, 360–364. Gower, C.F., Ryan, A.B., Rivers, T., 1990. Mid-Proterozoic Laurentia-Baltica: an overview of its geological evolution and summary of contributions made by this volume. In: Gower, C.F., Ryan, A.B., Rivers, T. (Eds.). Mid Proterozoic LaurentiaBaltica, GAC Special Paper 38, Geological Association of Canada, 1-22. Groves, D.I., Vielreicher, R.M., Godfarb, R.J., Condie, K.C., 2005. Controls on the heterogeneous distribution of mineral deposits through time. In: McDonald, I., Boyce, A.J., Butler, I.B., Herrington, R.J., Polya, D.A. (Eds.), Mineral Deposits and Earth Evolution, vol. 248. Geological Society of London Special Publication, pp. 71–101. Hanson, R.E., Wardlaw, M.S., Wilson, T.J., Mwale, G., 1993. U–Pb zircon ages from the Hook granite massif and Mwembeshi dislocation: constraints on Pan-African deformation, plutonism and transcurrent shearing in central Zambia. Precambrian Research 63, 189–209. Hattingh, P.J, 1989. The paleomagnetism of the upper zone of Bushveld Complex. Tectonophysics 165, 131–142. Heaman, L.M., 1997. Global mafic magmatism at 2.45 Ga: remnants of an ancient large igneous province? Geology 25, 299–302. Hollerbach, R., Jones, C.A., 1993. Influence of the Earth’s inner core on geomagnetic fluctuations and reversals. Nature 365, 541–543. Kent, D.V., Smethurst, M.A., 1998. Shallow bias of palaeomagnetic inclinations in the Palaeozoic and Precambrian. Earth and Planetary Science Letters 160, 391–402. Le Pichon, X., Huchon, P., 1984. Geoid, Pangaea and convection. Earth and Planetary Science Letters 67, 123–135. Ledru, P., Johan, V., Millesi, J.P., Tegyey, M., 1994. Markers of the last stage of the Palaeoproterozoic collision: evidence for a 2 Ga continent involving cicumAtlantic provinces. Precambrian Research 69, 169–191. Maruyama, S., Santosh, M., Zhao, D., 2007. Superplume, supercontinent, and postperovskite: mantle dynamics and anti-plate tectonics on the core–mantle boundary. Gondwana Research 11, 7–37. McClennan, S.M., Taylor, S.R., 1982. Geochemical constraints on the growth of the continental crust. Journal of Geology 90, 347–361. McKenzie, D.P., Weiss, N., 1975. Speculations on the thermal and tectonic history of the Earth. Geophysical Journal of the Royal Astronomical Society 42, 131–174. Mertanen, S., Halls, H.C., Vuollo, J.I., Pesonen, L.J., Stepanov, V.S., 1999. Palaeomagnetism of 2.44 Ga mafic dykes in Russia Karelia, eastern Fenoscandian Shield—implications for continental reconstructions. Precambrian Research 98, 197–221.

Mondal, S., Piper, J.D.A., Hunt, L., Bandyopadhyay, G., Basu Mallik, S., 2009. Palaeomagnetic and rock magnetic study of charnockites from Tamil Nadu, India, and the ‘Ur’ protocontinent in Early Palaeoproterozoic times. Journal of Asian Earth Sciences 34, 493–506. Morgan, G.E, Briden, J.C., 1981. Aspects of Precambrian palaeomagnetism, with new data from the Limpopo Mobile Belt and Kaapvaal Craton in southern Africa. Physics of the Earth and Planetary Interiors 24, 142–168. Pesonen, L.J., Elming, S.-S., Mertanen, S., Pisarevsky, S., D’Agrella-Filho, M.S., Meert, J.G., Schmidt, P.W., Abrahamsen, N., Bylund, G., 2003. Palaeomagnetic configuration of continents during the Proterozoic. Tectonophysics 375, 289–324. Petrasheck, W.E., 1973. Some aspects of the relations between continental drift and metallogenic provinces. In: Tarling, D.H., Runcorn, S.K. (Eds.), Implications of Continental Drift to the Earth Sciences, vol. 1. Academic Press, London, pp. 567–572. Phillips, B.R., Bunge, H.-P., Schaber, K., 2009. True polar wander in mantle convection models with multiple, mobile continents. Gondwana Research 15, 288–296. Piper, J.D.A., 1980. Analogous Upper Proterozoic apparent polar wander loops. Nature 283, 845–847. Piper, J.D.A., 1982. The Precambrian palaeomagnetic record: the case for the Proterozoic Supercontinent. Earth Planetary Science Letters 59, 61–89. Piper, J.D.A., 1987. Palaeomagnetism and the Continental Crust. Open University Press, 488 pp. Piper, J.D.A., 2003. Consolidation of Continental Crust in Late Archaean-Early Proterozoic times: a Palaeomagnetic test. Gondwana Research 6, 435–448. Piper, J.D.A., 2006. A ∼90◦ Late Silurian-Early Devonian apparent polar wander loop: the latest inertial interchange of planet Earth? Earth and Planetary Science Letters 250, 345–357. Piper, J.D.A., 2007. The Neoproterozoic Supercontinent Palaeopangaea. Gondwana Research 12, 202–227. Piper, J.D.A., Mallik, S.B., Bandyopadhyay, G., Mondal, S., Das, A.K., 2003. Palaeomagnetic and rock magnetic study of a deeply-exposed continental crustal section: the charnockite belt of Southern India. Precambrian Research 121, 185– 219. Poornachandra Rao, G.V.S., Mallikarjuna Rao, J., Reddi, A.E.B., 1989. Preliminary palaeomagnetic and geochemical studies of granulites from Tamil Nadu and Kerala. In: Proceedings of the IASPEI Meeting on Structure and Dynamics of Indian Lithosphere, Hyderabad, pp. 1–3. Prasad, C.V.R.K., Pulla Reddy, V., Subbarao, K.V., Radhakrishnamurty, C., 1987. Palaeomagnetism and the crescent shape of the Cuddapah Basin, Purana Basins of Peninsula India. Geological Society of India Memoir 6, 331–347. Richter, F.M., Parsons, B., 1976. On the interaction of two scales of convection in the mantle. Journal of Geophysical Research 80, 2529–2541. Rino, S., Kon, Y., Sato, W., Maruyama, S., Santosh, M., Zhao, D., 2008. The Grenvillian and Pan-African orogens: World’s largest orogenies through geologic time, and their implications on the origin of superplume. Gondwana Research 14, 51–72. Rogers, J.J.W., 1996. A history of continents in the past three billion years. Journal of Geology 104, 91–107. Rogers, J.J.W., Santosh, M., 2004. Continents and Supercontinents. Oxford University Press, 289 pp. Salminen, J., Pesonen, L.J., Reimold, W.U., Donaldini, F., Gibson, R.L., 2009. Palaeomagnetic and rock magnetic study of the Vredefort impact structure and the Johannesburg Dome, Kaapvaal Craton during the Mesoproterozoic. Precambrian Research 168, 167–184. Santosh, M., Maruyama, S., Yamamoto, S., 2009. The making and breaking of supercontinents: some speculations based on superplumes, superdownwelling and the role of tectosphere. Gondwana Research 15, 324–341. Sears, J.W., Price, R.A., 1978. The Siberian connection: a case for the Precambrian separation of the North American and Siberian cratons. Geology 6, 267–270. Sears, J.W., Price, R.A., 2000. New look at the Siberia connection: no SWEAT. Geology 28, 423–426. Senshu, H., Maruyama, S., Rino, S., Santosh, M., 2009. Role of tonalite–trondhjemite–granite (TTG) crust subduction on the mechanism of supercontinent breakup. Gondwana Research 15, 433–442. Shackleton, R.M., 1973. Correlations of structures across Precambrian orogenic belts in Africa. In: Tarling, D.H., Runcorn, S.K. (Eds.), Implications of Continental Drift to the Earth Sciences, vol. 2. Academic Press, London, pp. 1091–1095. Silver, P.G., Behn, M.D., 2008. Intermittent Plate Tectonics? Science 319, 85–89. Sinomson, B.M., Sumner, D.Y., Beukes, N.J., Johnson, S., Gutzmer, J., 2009. Correlating multiple Neoarchaean-Palaeoproterozoic impact spherule layers between South Africa and Western Australia. Precambrian Research 169, 100–111. Strik, G., Blake, T.S., Zegers, T.E., White, S.H., Langereis, C.G., 2003. Palaeomagnetism of flood basalts in the Pilbara Craton, Western Australia: Late Archaean continental drift and the oldest known reversal of the geomagnetic field, Journal of Geophysical Research 108. EPM 2, 1–21. Strik, G., de Wit, M.J., Langereis, C.G., 2007. Palaeomagnetism of the Neoarchaean Pongola and Ventersdorp Supergroups and an appraisal of the 3.0–1.9 Ga apparent polarwander path of the Kaapvaal Craton, Southern Africa. Precambrian Research 153, 96–115. Tsai, V.C., Stevenson, D.J., 2007. Theoretical constraints on true polar wander. Journal of Geophysical Research 112, B5, B05415. doi:10.1029/2005JB003923. Usui, Y., Tarduno, J.H., Watkeys, M., Hofmann, A., Cottrell, R., 2009. Evidence for a 3.45-billion–year-old magnetic remanence: hints of an ancient geodynamo from comglomerates of South Africa. Geochemistry, Geophysics, Geosystems 10, Q09207. doi:10.1029/2009GC002496. Watson, J.V., 1973. Effects of reworking on high-grade gneiss complexes. Philosophical Transactions of the Royal Society of London A273, 433–456.

J.D.A. Piper / Journal of Geodynamics 50 (2010) 154–165 Watson, J.V., 1978. Ore deposition through geological time. Proceedings of the Royal Society of London A362, 305–328. Williams, G.E., Schmidt, P.W., 1997. Palaeomagnetism of the Palaeoproterozoic Gowganda and Lorrain Formations, Ontario: low palaeolatitude for Huronian glaciation. Earth and Planetary Science Letters 153, 157– 169. Williams, H., Hoffman, P.F., Lewry, J.F., Monger, J.W.H., Rivers, T., 1991. Anatomy of North America: thematic portrayals of the continent. Tectonophysics 187, 117–134.

165

Windley, B.F., 1993a. The Evolving Continents, 3rd ed. John Wiley, New York, 526 pp. Windley, B.F., 1993b. Uniformitarianism today: plate tectonics is the key to the past. Journal of the Geological Society of London 150, 7–19. Zegers, T.E., de Wit, M.J., Dann, J., White, S.H., 1998. Vaalbara, Earth’s oldest assembled continent? A combined structural, geochronological and palaeomagnetic test. Terra Nova 10, 250–259. Zhao, D, 2009. Multiscale seismic tomography and mantle dynamics. Gondwana Research 15, 297–323.