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ISSN 0869-5911, Petrology, 2008, Vol. 16, No. 7, pp. 629–651. © Pleiades Publishing, Ltd., 2008.

INTRODUCTION

Modern hypotheses dealing with the origin andinner evolution of solid terrestrial planes are underlainmostly by a diversity of physical and geochemical com-putations and simulations on the basis of certainassumptions. A principal disadvantage of currentlyexisting hypotheses is their abstract nature and com-plete ignoring of information on the tectono-magmaticevolution of these bodies. At the same time, such dataprovide important information on the mechanisms ofthe origin and further evolution of planets, with suchinformatin needed to develop a modern theory of thegenesis and evolution of solid planetary bodies.

In this context, the most comprehensive informationis, of course, available on the Earth, and it will be usedbelow as an example (with data on other planets, mostlyon the Moon, employed where needed) in discussingprincipal problems of the evolution of the planetarybodies. Obviously, the most unbiased method appliedto studying the Earth’s evolution (as well as those ofother planets) is studying its magmatic processes thatprovided molted material to the surface and thusensured the continuous “record” of information on thecharacter of the melted sources throughout the wholegeological history of the planet. It is also self-evidentthat magmatic (petrologic) processes cannot be consid-ered separately from tectonic events, because the devel-

Irreversible Evolution of Tectono-Magmatic Processes at the Earth and Moon: Petrological Data

O. A. Bogatikov and E. V. Sharkov

Institute of the Geology of Ore Deposits, Petrography, Mineralogy, and Geochemistry (IGEM), Russian Academy of Sciences, Staromonetnyi per. 35, Moscow, 109017 Russia

e-mail: oleg@igem.ru, sharkov@igem.ru

Received March 3, 2008

Abstract

—The tectono-magmatic evolution of the Earth and Moon started after the solidification of their mag-matic “oceans”, whose in-situ crystallization produced the primary crusts of the planets, with the compositionof these crusts depending on the depths of the “oceans”. A principally important feature of the irreversible evo-lution of the planetary bodies, regardless of their sizes and proportions of their metallic cores and silicate shells,was a fundamental change in the course of their tectono-magmatic processes during intermediate evolutionarystages. Early in the geological evolution of the Earth and Moon, their magmatic melts were highly magnesianand were derived from mantle sources depleted during the solidification of the magmatic “oceans”; this situa-tion can be described in terms of plume tectonics. Later, geochemically enriched basalts with high concentra-tions of Fe, Ti, and incompatible elements became widespread. These rocks were typical of Phanerozoic within-plate magmatism. The style of tectonic activity has also changed: plate tectonics became widespread at theEarth, and large depressions (maria) started to develop at the Moon. The latter were characterized by a signifi-cantly thinned crust and basaltic magmatism. These events are thought to have been related to mantle super-plumes of the second generation (thermochemical), which are produced (Dobretsov et al., 2001) at the bound-ary between the liquid core and silicate mantle owing to the accumulation of fluid at this interface. Because oftheir lower density, these superplumes ascended higher than their precursors did, and the spreading of their headparts resulted in active interaction with the superjacent thinned lithosphere and a change in the tectonic regime,with the replacement of the primary crust by the secondary basaltic one. This change took place at 2.3–2.0 Gaon the Earth and at 4.2–3.9 Ga on the Moon. Analogous scenarios (with small differences) were also likely typ-ical of Mars and Venus, whose vast basaltic plains developed during their second evolutionary stages. Thechange in the style of tectonic-magmatic activity was associated with important environmental changes on thesurfaces of the planets, which gave rise to their secondary atmospheres. The occurrence of a fundamentalchange in the tectono-magmatic evolution of the planetary bodies with the transition from depleted to geochem-ically enriched melts implies that these planets were originally heterogeneous and had metal cores and silicateshells enriched in the material of carbonaceous chondrites. The involvement of principally different material(that had never before participated in these processes) in tectono-magmatic processes was possible only if thesebodies were heated from their outer to inner levels via the passage of a heating wave (zone) with the associatedcooling of the outermost shells. The early evolutionary stages of the planets, when the waves passed throughtheir silicate mantles, were characterized by the of development of superplumes of the first generation. Themetallic cores were the last to melt, and this processes brought about the development of thermochemical super-plumes.

DOI:

10.1134/S0869591108070011

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opment of melting regions in the upper shells of theplanets was related to the ascent of mantle plumes andthe subsequent spread of their head parts. In fact, tec-tonics and magmatism act as two aspects of a singlegeodynamic process, and hence, understanding of fun-damental relations and tendencies in the evolution ofplanets as self-developing systems relies on the simul-taneous analysis of tectonics and magmatism over thewhole geological history.

Our research was focused on the comparative anal-ysis of the tectono-magmatic evolution of the Earth andMoon carried out in order to identify evolutionary fea-tures of the planets that can be also applied to investi-gating other terrestrial planets (Venus, Mars, and Mer-cury). All of these planetary bodies have generally sim-ilar inner structures (they possess metallic cores andsilicate shells) but differ in size, with the Earth beingand largest and the Moon the smallest among the plan-etary bodies. In this publication we will briefly dwell onissues of the origin of the primary compositions of theEarth’s and Moon’s crusts, distinctive features of tec-tonic processes in the planets during their major evolu-tionary stages, and then, proceeding from these consid-erations, we will discuss the leading mechanisms con-trolling their origin and evolution. We have alreadyaddressed these problems (Bogatikov et al., 2000;Sharkov and Bogatikov, 2001), but recently obtaineddata allowed us to further specify certain aspects.

ORIGIN OF THE PRIMARY CRUSTS OF THE PLANETARY BODIES

Over the Earth’s lifetime (4.6–4.5 Ga) is subdividedinto a number of evolutionary stages, which have irre-versibly modified both the character of deep processesin this planet and the manifestations of these processesat the surface. One of the principally important prob-lems is the composition of the primary crust of theEarth, an issue that have long been discussed and whoseessence is addressed from the following two view-points:

According to the traditional viewpoint, which wasshaped more than 150 years ago, when concepts of geo-synclines were predominant, the crust initially had abasite composition, while the sialic crust was producedlater by either the geosyncline process or, in modernterms, at convergent plate boundaries. Proponents ofthis concept maintain that the continental crust contin-uously grew at the expense of oceanic one.

The primary crust was sialic, and plate tectonicsstarted to act no earlier than the Paleoproterozoic, whenthe oceanic crust began to grow, while the ancient crust,including the continental one, was involved in subduc-tion processes and was “stored” in megaliths (“slabcemetery”) within the mantle, i.e., the ancient continen-tal crust was gradually replaced by secondary oceanicone.

From the standpoint of petrology and physicalchemistry, there is no principal difference between thetwo above viewpoints: both models require the globalmelting of the primary material of the Earth’s upper-most shell (carbonaceous chondrite) to form the pri-mary Earth’s crust. In this situation, the result shouldhave depended only on the degree of melt differentia-tion during the crystallization of the global magmatic“ocean”. The development of the latter could be initi-ated by the energy released during the compaction ofthe material, the presence of short-lived isotopes, tidaleffects, massive meteoritic bombardment, and otherprocesses (Ringwood, 1978).

As was demonstrated in the late 1920s by the prom-inent British physicist Jeffries (1929), the solidificationof large melt volumes of significant vertical extent iscontrolled by two independent parameters: the adia-batic gradient and the temperature gradient of the melt-ing point (the slope of the liquidus line in

P–T

space).Because the slope of the liquidus line of silicate meltsis roughly one order of magnitude steeper than the adi-abatic gradient, solidification can proceed only frombelow upward, because the bulk of the melt is over-heated relative to the liquidus (Fig. 1). This model waslater abandoned, in spite of its universal character andapplicability to both the solidification of molten planets(as was pointed out by H. Jeffries himself) and thesolidification of the Earth’s liquid core or even largeintrusions (Sharkov, 2006). The only exception iswater, whose density is at a maximum at +4

°

C and sig-nificantly decreases, up to the appearance of a relativelylow-density solid phase (ice). Because of this, thefreezing of water bodies proceeds from below, bymeans of the growth of younger ice layers atop olderones, which explains why these bodies only rarelyfreeze completely from their top parts to bottoms.

The in-situ (unidirectional) solidification of planetsshould have resulted in their crusts made up of the low-est temperature derivatives, whose compositiondirectly depended on the depths of the magmatic“oceans”: the deeper the “oceans”, the greater amount

1280 1290 1300 13108

6

4

2

0

Melting-point gradient,

Adi

abat

ic g

radi

ent,

Thi

ckne

ss, k

m

Crystallization zone

T

, ºC

3ºC/km

0.3º

C/k

m

Fig. 1.

Relations between the adiabatic and melting-pointgradients that determine the solidification direction of anylarge melt volumes.

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IRREVERSIBLE EVOLUTION OF TECTONO-MAGMATIC PROCESSES 631

of low-temperature components accumulated untilcomplete solidification.

Geological data on the strong predominance ofgranitoids in the Archean crust, with no analogues ofthese rocks occurring in younger epochs (Sharkov,2003 and references therein), and data obtained onancient detrital zircons from Archean metasediments inAustralia (the mineral crystallized from granitic melt;Valley et al., 2002; Harrison, 2007) are more consistentwith a sialic composition of the primary crust. Becausethe K

2

O concentration in chondrites is one order ofmagnitude higher than the Na

2

O concentration in them,it is reasonable to expect that these should be plagiog-ranites. This is fully consistent with data on the Earth’sArchean crust, particularly the early one, which is dom-inated by granitoids of tonalite–trondhjemite–grano-diorite (TTG) group (Fig. 1), which are usually exten-sively migmatized. This primary sialic crust was multi-ply recycled in the course of the later tectono-magmaticevolution of the Earth. This recycling usually involvedthe partial (migmatization) or complete remelting ofthe material, which inevitably resulted in the partial orcomplete transformation of the isotopic systems and,correspondingly, made the isotopic age of the rocks“younger”.

The growth of the primary crust during the in-situsolidification of the magmatic “ocean” should haveresulted in the depletion of the upper mantle material inlow-melting components, which can be seen, indeed, inthe most ancient volcanic rocks (among those everfound as of now: they were dated at 3.8–3.7 Ga) of theIsua Belt in southwestern Greenland (Frei et al., 2004).The complementary character of the compositions ofthe depleted mantle and continental crust has long beenregarded by some geochemists as evidence of the sepa-ration of the latter from the primitive mantle early in theEarth’s evolution (Galer and Goldstein, 1991).

From this standpoint, the vertical section of theupper primary lithosphere should have been character-ized by the transition from ultrabasites of the diamonddepth facies to ultramafics of the garnet and spinelfacies, mafites of the lower crust and end with plagiog-ranites of moderate and lower pressures.

The whole gas envelope that surrounded the proto-Earth was likely supplemented with volatiles escapingfrom the solidifying magmatic ocean. Upon the coolingof the Earth’s surface to temperatures below 100

°

C,water should have started to condense and form theocean and primary atmosphere, a concept that findssupport in isotopic and geochemical lines of evidence(Veiser, 1976; Taylor and McLennan, 1985). From thisstandpoint, the water of the ocean can be regarded asthe oldest geological “body”.

In contrast to the Earth, the Moon, which has an ageof 4.56

±

0.07 Ga (Alibert et al., 1994), remains practi-cally unchanged and consists of unusual ferrousanorthosites (FAN). Their genesis is usually thought tobe related to plagioclase flotation in the global mag-

matic ocean that was formed shortly after the origin ofthe Moon (Ringwood, 1978; Snyder et al., 2000; andothers). However, the same result could be produced bythe in-situ solidification of the magmatic “ocean”, andhence, plagioclase flotation cannot be viewed as theonly possible mechanism able to form the primarybasite crust at the Moon. Since the Moon is approxi-mately six times smaller than the Earth, the depths ofthe magmatic oceans of these planets should have beendifferent, which likely accounts for the differences inthe composition of their primary cores. The primaryatmosphere of the Moon could not be retained by thisplanet because of its weak gravity.

EVOLUTION OF TECTONO-MAGMATIC PROCESSES ON THE EARTH AND MOON

Tectono-Magmatic Evolution of the Earth

The geological evolution of the Earth is currentlysubdivided into the following four major stages: pre-geologic, nuclearic, cratonic, and continental–oceanic(Bogatikov et al., 2000).

Pregeologic (hidden or Hadean) Stage

Practically no information about this stage is avail-able as of now. The stage started after the origin of theprimary Earth’s crust and terminated approximately4 Ga. The only data on this stage are provided by detri-tal zircons from Australia, which were dated at 4.4–4.2Ga. It is known that these zircons were formed duringthe crystallization of granitic melt, that liquid water wasalready present, and, possibly, that sialic crust alreadyexisted as long ago as 4.5 Ga (Harrison, 2007).

Nuclearic Stage

The first, nuclearic, stage in the geological history ofthe Earth (from ~4.0 to 2.7–2.5 Ga) spans the wholeArchean (Bogatikov et al., 2000). The occurrence ofsedimentary rocks and pillow lavas produced by sub-marine eruptions (Polat et al., 2003) in the aforemen-tioned Isua greenstone belt testifies that large waterbodies existed as early as 3.7–3.8 Ga, although they didnot necessarily have an oceanic crust; they formed theWorld Ocean.

The major tectonic structures in the Archean weresimultaneously developing large granite–greenstone ter-ranes (GGT), which consisted of rocks of the tonalite–trondhjemite–granodiorite composition (TTG) cut by anunsystematic network of greenstone belts and with gran-ulite belts occurring between the terranes. GGT werecharacterized by uplift, extension, and erosion, whereasthe granulite belts were zones of contraction, subsidence,and sedimentation and were the Earth’s most ancientsedimentation basins (Taylor and McLennon, 1985). Thedistinctive structural features of Archean areas are illus-trated below by the example of the Baltic Shield (Fig. 2).Insight into the evolution of greenstone belts in them is

1

1

632

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BOGATIKOV, SHARKOV

24º 28º 32º 36º

60º

Caledo

nides

BMB

L. Ladoga

L. Onega

BMB KNGBM

KGGT

BARENTS SEA

M

WHITE SEAGulf of Bothnia

L

1 2 3 4 5

6 7 8

a b

a b

28º 32º 36º 40º

68º

66º

64º

62º

70º

68º

100 km

a b

Svekofennides

provided by Rybakov’s (1987) reconstructions for cen-tral Karelia.

The origin of Archean TTG rocks is still disputable.Based on geochemical evidence, many researchers con-sider them to be analogues of arc-type adakites (Martinet al., 2005 and references therein) and believe that theywere produced in relation to subduction processes. Oth-ers believe that these rocks are relics of the primary

sialic crust (see above) extensively recycled by laterprocesses and maintain that greenstone belts are sort ofprotorift structures (Bogatikov et al., 2000;

Early Pre-cambrian…

, 2005).Greenstone belts (GB) were main zones of mantle

magmatism and accounted for no more than 10

−

20% ofGGT by area. These are commonly fragments of longgraben-shaped structures that form an unsystematic

Fig. 2.

Eastern part of the Baltic Shield in the Late Archean (3.0–2.7 Ga).(1) Granite–greenstone terranes (KGGT—Karelian granite–greenstone terrane, M—Murmansk block); (2) greenstone belts;(3) reconstructed by Rybakov (1987) in the southeastern part of the Karelian craton; (4) Belomorian mobile belt (BMB):(a) observed, (b) inferred (L—Lottiiskii domain, T—Terskii domain); (5) Kola–Norwegian granulite belt (KNGB) and relatedenderbites: (a) observed, (b) inferred; (6) Keivy structure; (7) Svecofennides; (8) boundaries: (a) observed, (b) inferred. Contouredpolygonal area is S.I. Rybakov’s study area.The inset shows the location of principal structural domains in the eastern part of the Baltic Shield in the Late Proterozoic.

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IRREVERSIBLE EVOLUTION OF TECTONO-MAGMATIC PROCESSES 633

network in a TTG matrix. The belts are dominated bypillow lavas of the komatiite–basalt series and volcanicrocks that are geochemical analogues of the Phanero-zoic boninite series and that include, along with ultra-mafic–mafic lavas, also volcanics of intermediate andacid composition. Sedimentary rocks are contained inthem in subordinate amounts and are volcaniclasticdeposits, graywackes, quartzites, banded iron forma-tions (BIF), siltstones, limestones, and dolomites. Sed-imentation took place mostly at shallow depths, so thatthe rocks occasionally show ripple marks and dryingcracks. Scarce Fe–Ti basalts appeared in theNeoarchean, but these rocks were rare and did notdetermine the character of Archean magmatism.

The komatiite–basalt series consists of low- andmoderate-Si high-Mg (12 < MgO < 32 wt %) volcanicrocks: komatiites (MgO > 18 wt %, SiO

2

< 52 wt %; LeBas, 2000), komatiite basalts and picrites (12, MgO <18 wt %), and tholeiite basalts (MgO = 5–12 wt %)(Table 2). According to experimental data, the depth atwhich komatiite melts should have separated from theirresidues can be evaluated from the depletion of themelts in Al: komatiite melts not depleted in Al sepa-rated from the harzburgite residue at depth of no lessthan 200 km, and Al-depleted varieties segregated fromthe garnet harzburgite residue at depths of about 250–300 km and from majorite-bearing residue at depths ofno less than 400 km (Arndt and Nesbitt, 1982). As fol-lows from data obtained on Archean GB, they containmelts coming from the whole range of these depth levels.

The other, boninite-like series comprises mafic–ultramafic high-Mg rocks of elevated silicity, whichvary in composition from picrite and basalt (“island-arc”) to andesite and dacite. These rocks are chemicallyvery close to the Phanerozoic boninite series accordingto the LeBas classification: MgO > 8 wt %, SiO

2

>52 wt %, and TiO

2

< 0.5 wt % (Table 3). It is worthmentioning that volcanics of exactly this type composethe oldest (3.7–3.8 Ga) Isua belt (Frei et al., 2004), andthe very first komatiites are known to appear only in thePaleoarchean, for example, in the 3.5-Ga Pilbara craton(Green et al., 2000).

Many researchers thought that these boninite-likerocks provide evidence of subduction processes (Polatet al., 2002; Shchipansky et al., 2004; and others).However, no geological lines of evidence, other thangeochemical ones, are available on the operation ofplate tectonics in the Archean. An important role intheir genesis was rather played by the assimilation ofthe depleted lithospheric mantle and crust by komatiitemelts ascending through them, as was determined formelts of silicic high-Mg series of the cratonic stage (seebelow).

In contrast to greenstone belts, synkinematic mag-matism in granulite belts, which are dominated bymetasediments, produced enderbites and charnockites.

Table 1.

Representative analyses of rocks of Archean tonalite–trondhjemite–granodiorite (TTG) association in the VedlozeroBlock, Karelian GGT (after Types of Magmas…, 2006) in com-parison with Phanerozoic adakites (after Condie, 2005)

Compo-nent

Karelian GGT Phanero-zoic adak-

itesK-12/86 K-6/86 K-13/86

SiO

2

70.02 71.79 67.56 65.9

TiO

2

0.35 0.19 0.38 0.47

A1

2

O

3

15.53 15.91 16.87 16.5

Fe

2

O

3

t 2.80 2.41 3.61 4.11

MnO 0.12 0.03 0.05 0.09

MgO 1.01 0.80 1.51 1.67

CaO 4.08 3.23 4.22 4.36

Na

2

O 5.18 4.77 4.62 4.00

K

2

O 0.79 0.75 1.11 2.14

P

2

O

5

0.12 0.13 0.07 0.12

Mg# 0.42 0.40 0.45 45.4

Na/K 6.56 6.35 4.15 2.04

Na + K 5.97 5.52 5.73 6.14

V 43 36 63 –

Cr 9 6 15 32

Co 6 5 8 –

Ni 10 6 15 12

Rb 20 13 31 63

Sr 398 405 417 493

Y 3 4 9 14.9

Zr 129 155 169 122

Ba 486 310 371 716

Hf 3.19 4.00 4.00 3.4

í‡ 0.11 0.07 0.32 0.75

La 25.1 18.0 31.4 17

Ce 46.4 32.8 54.6 34

Nd 15.5 11.1 19.1 16

Sm 1.97 1.66 2.59 3.1

Eu 0.62 0.67 0.71 0.84

Gd 2.2 1.17 1.86 2.8

Tb 0.13 0.13 0.26 0.40

Yb 0.25 0.31 0.73 1.16

Lu 0.04 0.05 0.11 0.18

W 0.11 0.09 0.11 –

Th 4.48 3.26 4.97 7.6

U 0.21 0.21 0.26 1.9

(La/Yb)

n

68.1 39.8 28.8 11.3

Note: Here and below, major elements (oxides) are given in wt %,trace elements are in ppm, Mg# = MgO/(MgO + 0.79 Fe

2

O

3

t).

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BOGATIKOV, SHARKOV

Table 2.

Representative analyses of Al-depleted and Al-undepleted komatiites and komatiite basalts from the Abitibi greenstone belt(Sproule et al., 2002; Polat et al., 1999)

Component

Al-undepleted komatiite basalts and komatiites Al-depleted komatiite basalts and komatiites

Komatiites Komatiite basalts Komatiites Komatiite basalts DH95-2HEG-1 SC96-6 DH96-9 SC95-53 DH96-10 DH96-11 DH96-12

SiO

2

47.27 46.10 47.07 49.11 51.08 42.10 41.60 49.25

TiO

2

0.35 0.38 0.43 0.60 0.82 0.62 0.58 1.11

Al

2

O

3

7.95 7.43 8.75 10.55 10.73 4.05 3.60 5.99

Fe

2

O

3

11.9 12.1 13.5 13.9 12.8 20.7 19.9 14.9

MnO 0.20 0.18 0.19 0.22 0.20 0.25 0.24 0.23

MgO 24.9 21.9 20.4 16.3 14.3 30.3 28.2 17.5

CaO 7.18 7.26 8.75 8.17 9.01 1.93 5.71 10.28

Na

2

O 0.22 0.09 0.78 1.01 0.87 0.10 0.05 0.66

K

2

O 0.01 0.02 0.09 0.08 0.13 0.01 0.01 0.09

P

2

O

5

0.02 0.03 0.04 0.05 0.08 0.04 0.04 0.07

LOI 5.15 4.80 3.09 3.90 0.65 7.12 7.70 1.63

Mg# 81 78 75 70 69 74 74 70

Cr 3028 1631 1142 1770 775 2250 2271 1081

Co 102 105 97 92 65 158 169 99

Ni 972 804 461 546 209 1544 1822 845

Sc 26 28 38 40 47 15 26 45

V 154 171 167 252 252 132 149 259

Ta 0.04 0.06 0.09 0.14 0.13 0.20 0.15 0.34

Nb 0.58 1.08 1.49 1.81 1.96 3.42 2.56 5.89

Zr 20 38 27 49 57 40 29 59

Hf 0.56 0.81 0.77 1.37 1.63 1.02 0.79 1.80

Tb 0.07 0.19 0.13 0.38 0.32 0.28 0.23 0.44

Y 8.0 8.9 11.0 15.1 20.5 6.0 5.0 13.8

La 0.62 1.59 1.51 2.37 3.26 3.31 3.03 4.17

Ce 2.17 4.21 4.10 6.37 8.93 8.48 7.48 12.86

Pr 0.36 0.61 0.61 0.91 1.38 1.19 1.08 1.97

Nd 1.89 3.12 3.16 4.63 6.95 5.60 5.23 9.28

Sm 0.74 0.99 1.02 1.66 2.24 1.57 1.32 2.79

Eu 0.250 0.326 0.350 0.748 0.660 0.320 0.500 0.976

Gd 1.10 1.31 1.54 2.29 3.11 1.62 1.40 3.14

Tb 0.220 0.242 0.250 0.384 0.530 0.220 0.200 0.447

Dy 1.49 1.60 1.82 2.58 3.71 1.36 1.26 2.69

Ho 0.330 0.357 0.360 0.582 0.800 0.270 0.240 0.508

Er 0.98 1.01 1.18 1.65 2.33 0.77 0.58 1.37

Tm 0.150 0.158 0.170 0.249 0.360 0.100 0.090 0.193

Yb 0.98 0.96 1.16 1.48 2.21 0.61 0.50 1.07

Lu 0.160 0.147 0.150 0.177 0.330 0.100 0.060 0.163

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IRREVERSIBLE EVOLUTION OF TECTONO-MAGMATIC PROCESSES 635

Table 3.

Representative analyses of Archean geochemical analogues of rocks of the Phanerozoic boninite series, Hisovaara structure,Karelian GGT, Baltic Shield (Shchipansky et al., 2004)

Component KH-101 N-302 N-333 N-334 N-326 N-323 KH-102 N-335 N-332

SiO

2

45.75 46.59 47.90 50.79 53.96 55.18 60.76 63.63 66.21

TiO

2

0.28 0.40 0.42 0.44 0.44 0.41 0.35 0.36 0.33

Al

2

O

3

6.85 9.49 10.24 11.41 12.13 12.20 9.59 9.46 10.53

Fe

2

O

3

t 12.80 13.43 13.14 11.83 10.95 10.15 9.43 10.13 10.23

MnO 0.18 0.19 0.18 0.18 0.17 0.19 0.16 0.15 0.16

MgO 29.22 21.27 16.86 13.15 10.31 9.20 8.51 8.07 8.38

CaO 4.79 8.22 10.65 11.59 11.09 12.25 10.28 7.79 3.37

Na

2

O 0.00 0.30 0.50 0.45 0.70 0.30 0.65 0.18 0.61

K

2

O 0.02 0.04 0.06 0.09 0.17 0.06 0.14 0.16 0.13

P

2

O

5

0.11 0.06 0.05 0.07 0.07 0.07 0.12 0.06 0.04

LOI 9.82 4.94 1.62 1.75 1.83 2.76 2.59 1.73 0.79

Sc 21 29 40.8 43.8 47.9 34.2 30.8 45.7 48

V 142 213 167 233 254 126 130 179 228

Cr 2840 2150 1412 1199 912 688 1680 1603 977

Co 89.9 74.5 72.2 64.2 66.8 63.7 72.2 77.9 66.7

Ni 905 597 325 241 219 221 452 342 226

Rb 1.82 0.28 0.01 0.42 2.66 1.26 2.52 6.84 7.38

Sr 76.1 26.4 25.5 86.7 127 54.7 102 55.2 37.2

Y 3.74 8.58 12.8 13 14.8 10.7 9.73 13.6 11

Zr 10.8 16.27 25.26 27.78 35.21 15.87 20.7 21.23 27.78

Nb 0.251 0.639 0.821 0.876 0.971 0.81 0.67 0.818 1.1

Ba 13.1 8.3 8.77 17.2 23.8 36.4 47.4 37.6 45.1

La 0.349 0.948 1.17 1.61 1.4 1.08 1.03 1.44 1.99

Ce 0.882 2.46 3.02 3.62 3.58 2.79 2.64 3.51 4.82

Pr 0.131 0.4 0.496 0.606 0.599 0.448 0.441 0.587 0.712

Nd 0.832 2.24 2.58 2.89 2.96 2.43 2.44 2.4 3.26

Sm 0.341 0.799 0.983 0.959 1.07 0.902 0.935 0.987 1.12

Eu 0.086 0.316 0.373 0.448 0.45 0.339 0.447 0.261 0.182

Gd 0.429 1.15 1.22 1.54 1.51 1.27 1.2 1.42 1.46

Tb 0.109 0.235 0.225 0.277 0.298 0.259 0.248 0.274 0.267

Dy 0.693 1.56 1.87 2.13 2.06 1.86 1.74 1.9 1.8

Ho 0.161 0.362 0.387 0.455 0.473 0.416 0.39 0.469 0.491

Er 0.421 1.04 1.3 1.34 1.56 1.3 1.21 1.39 1.45

Tm 0.089 0.167 0.206 0.213 0.244 0.231 0.184 0.229 0.228

Yb 0.449 1.01 1.02 1.37 1.68 1.3 1.09 1.13 1.34

Lu 0.078 0.144 0.176 0.201 0.206 0.173 0.387 0.177 0.221

Hf 0.25 0.46 0.7 0.76 0.89 0.39 0.71 0.68 0.74

Ta 0.023 0.047 0.047 0.06 0.049 0.052 0.041 0.054 0.057

Th 0.052 0.098 0.121 0.083 0.122 0.099 0.081 0.113 0.137

U 0.009 0.026 0.027 0.029 0.041 0.025 0.032 0.026 0.019

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Caledonidis

Barents Sea

White SeaGulf orfBothnia

Karel ian

B. Levgora

Vetreny Belt

Myandukha

OnegaLadoga

Svecofennides

Finnish Gulf

Sediments o

f Russi

an Platform

80 km

LP

1

2

34 5

6

7

8

9

10

T

I - V

LGB

Svecofennides

BMB

UGB TLB

1 2 3 4

5 6 7 8

Kola KratonKarelian Craton

Kola Kraton

Be l o o r i a n B

e l t

GoletsCraton

Lake

Lake

MLT

LGB

Cratonic Stage in the Earth’s Evolution

By the latest Archean–earliest Proterozoic (2.7–2.5 Ga), the Earth’s crust had been cratonized, as fol-lows from the development of rift structures, vast dikeswarms, and large mafic–ultramafic layered intrusions(Fig. 3). As in the Archean, granulite belts developmentbetween cratons and were surrounded by gently slopingzones of tectonic flow. In the Early Paleoproterozoic

(2.5–2.3 Ga), the predominant magma type comprisedmelts of silicic high-Mg series (SHMS), which vary incomposition from low-Ti picrite to andesite, dacite, andrhyolite at the predominance of basalt. In contrast toboninite-LILE Archean rocks, which often were com-plete geochemical analogues of Phanerozoic boninites,the TiO

2

concentration in SHMS volcanics reached0.7–1.0 wt % (Table 4). Some Early Paleoproterozoicrift structure, such as the Vetrennyi Belt in Karelia, con-

Fig. 3.

Large Early Paleoproterozoic Baltic magmatic province of the silicic high-Mg (boninite-like) series (SHMS).(1) Svecofennides; (2) Paleoproterozoic sedimentary–volcanic complexes (P—Pechenga, I–V—Imandra–Varzuga); (3) intermedi-ate mobile belts (BMB—Belomorian mobile bel, TLB—Terskii–Lotiiskii domain: L—Lottiiskii segment, T—Terskii segment),(4) Lapland (LGB)–Umba (UGB) granulite belt; (5) Archean basement; (6) layered intrusions (circled numerals: 1—Koitilainen,2—Tornio, 3—Kemi, 4—Pinakat, 5—Koilismaa, 6—Olangskaya group, 7—Mount General’skaya, 8—Moncha-Tundra, 9—Fedor-ovo–Pana, 10—Borakovka); (7) Main Lapland Fault (MLL); (8) northern boundary of the Baltic Shield.The inset shows the location of principal structural domains in the eastern part of the Baltic Shield in the Early Paleoproterozoic(Sumian–Sariolian); TLB—Terskii–Lottiiskii belt.

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tain preserved and practically unaltered lavas of thisseries, which bear magnesian olivine and pyroxenes,chromite, and volcanic glass of composition varyingfrom basaltic andesite to dacite. The volcanic rocks,dikes, and layered intrusions of SHMS produce largemagmatic provinces resembling those of Phanerozoicflood basalts (Sharkov et al., 2005). SHMS melts werederived at shallower depths than those from whichkomatiites came: 100–150 km (Girnis and Ryabchikov,1988).

The development of large magmatic provinces(whose variety is Archean granite–greenstone terranes;Ernst et al., 2007 and references therein) implies thatmantle superplumes of the first generation existedbeneath them. In contrast to Phanerozoic superplumes,these consisted of depleted mantle material. Thespreading of their head portions was associated with theadiabatic melting of the rocks at depths from 200 to 400km and even more (see above) and did not result in any

significant dynamic interaction of the superplumes withthe thick overlying lithosphere but merely in the devel-opment of uplift and extension regions above them(GGT in the Archean and cratons in the Early Paleopro-terozoic) and subsidence zones (granulite belts) abovedescending fluxes in the mantle between nearby super-plumes, without break-up of the continental crust. Thesituation can be described in terms of plume tectonics.A hypothetical scheme of the tectonic activity duringthe Earth’s first evolutionary stage is shown in Fig. 4.

Continuation of the Cratonic Stage: Transitional Period at 2.3–2.0 Ga

During the second half of the cratonic stage at 2.3–2.0 Ga, the character of magmatic activity globallychanged from the derivation of SHMS to Fe–Ti picritesand basalts typical of within-plate magmatism in thePhanerozoic. Along with high concentrations of Fe, Ti,

Table 4.

Representative analyses of Early Paleoproterozoic volcanic rocks of the silicic high-Mg series (SHMS) in the Vetren-nyi Belt, eastern Karelia, Baltic Shield

Component Lev‚10 Lev‚19 Gl3b M 33 Gl2b M 1 Lev‚5 M 99 Lev‚16 Gl2e

SiO

2

42.87 44.19 45.08 51.27 48.36 51.65 51.21 51.66 52.05 54.68

TiO

2

0.35 0.42 0.49 0.67 0.66 0.68 0.74 0.73 0.73 0.72

Al

2

O

3

4.62 5.97 11.90 9.84 13.04 10.18 10.45 12.99 13.73 13.46

Fe

2

O

3

t 12.69 12.81 12.84 12.32 12.68 12.21 13.09 11.88 12.38 10.76

MnO 0.18 0.19 0.17 0.18 0.19 0.18 0.19 0.18 0.19 0.17

MgO 33.71 29.23 21.25 14.58 13.74 13.72 12.13 9.99 8.10 8.39

CaO 4.91 6.17 6.41 9.17 9.39 9.44 10.42 9.94 10.82 9.06

Na

2

O 0.49 0.76 0.64 1.54 1.55 1.26 1.65 1.87 1.62 2.43

K

2

O 0.16 0.24 1.16 0.37 0.35 0.62 0.06 0.68 0.34 0.29

P

2

O

5

0.03 0.03 0.06 0.06 0.04 0.06 0.05 0.09 0.04 0.05

Ba 140 150 77 77 381 100 110 71 199 112

Rb 5 6 85 8 23 12 0 10 6 6

Sr 100 119 74 192 186 186 192 195 177 208

Y 13 14 11 14 9 15 20 16 19 13

Zr 44 49 76 76 137 71 69 74 69 87

Nb 2 1 2 9 4 4 3 6 4 2

Ga 14 10 17 23 14 17 17

Zn 85 73 72 87 74 105 80 81 76 70

Cu 63 72 97 98 95 93 96 95 97 109

Ni 938 835 773 257 415 244 141 192 57 192

V 122 140 129 220 168 231 200 231 222 187

Cr 2987 2671 1975 – 1115 – 750 – 367 745

Sc 24 25 23 45 35 28 34 34 34 49

Co 113 97 75 85 51 98 68 69 47 43

Note: Analyses were made by XRF at the Central Chemical Laboratory of the Institute of the Geology of Ore Deposits, Petrography, Min-eralogy, and Geochemistry (IGEM), Russian Academy of Sciences, analyst A.I. Yakushev. Samples: Lev10 and Lev19—olivine–clinopyroxene cumulates; Gl2e and Lev 16—fine-grained clinopyroxene–plagioclase dolerites; other samples are olivine–clinopy-roxene and clinopyroxene boninite basalts.

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and P (Table 5, Fig. 5), these rocks are enriched in alka-lis, Ba, Zr, LREE, and other incompatible elements. Atthe same time, the intensity of the magnetic field dras-tically increased and reached its maximum over thewhole Earth’s history and continued decreasing after-ward (Stevenson et al., 1983).

The style of tectonic activity first did not change. Asin the Early Paleoproterozoic, large magmatic prov-inces were formed, such as the Jatulian–Ludicovian inthe Baltic Shield, where newly erupted lava flows builtup the sequences in the same rift structures, and dikeswarms and large layered intrusions were emplaced,which were Ti-bearing (for example, Gremyakha–Vyrmes in the Kola Peninsula and Elet’ozero in Kare-lia). An analogous situation took place in all other Pre-cambrian shields (Sharkov and Bogina, 2006). In otherwords, large magmatic provinces gave way to others (ofmodern type) in the same territories. Obviously, thiscan be explained by a change in the composition of thesource regions of the mantle superplumes (see below).

The transitional period lasted approximately300 Ma and did not affect tectonic processes. Locallylarge SHMS provinces continued to develop, as atBushveld in the South African Shield or Voronezh Mas-sif in the East European craton.

The change in the composition of the magmaticmelts was simultaneous and likely initiated drasticchanges in the surface environment of the Earth: the

appearance of an oxidizing atmosphere, global glacia-tions, a positive shift in the carbon isotopic compositionof sedimentary carbonates, the appearance of phospho-rites and hydrocarbons, and significant changes in thebiosphere, such as the flowering of cyanobacteria(blue–green algae), which ensured an oxidizing com-position of the atmosphere and the appearance of mul-ticellular organisms (Melezhik et al., 2005).

Continental–Oceanic Stage in the Earth’s Evolution

This stage, which still continues, started at ~2.0 Ga,when a change in the style of magmatic activity wasfollowed by changes in the style of tectonism: the firstfoldbelts (orogens) of the Phanerozoic type started thendeveloping: Svecofennian in the Baltic Shield, Trans-Hudson, Wopmay, Penokean, and others in the Cana-dian Shield, and analogous orogens in the Australian,Ukrainian, Aldan, and other shields. (Sharkov and Bog-ina, 2006). These orogens included ophiolite associa-tions [as they were understood at the Penrose confer-ence (Anonymous, 1972), i.e., consisting of four com-plexes: mantle ultrabasites, lower-crustal gabbroids,parallel dikes, and pillow lavas] and testifying that arelatively thin oceanic crusts had already been formed.Numerous lines of evidence of the existence of islandarcs, backarc basins, active continental margins,

Fig. 4. Hypothetical scheme of tectonics in the Early Precambrian.(1) Spreading head parts of mantle superplumes; (2) regions with descending mantle flows between superplumes, where granulitebelts are formed; (3) sedimentary basins; (4) magma generation regions: (a) in the mantle, (b) in granulite belts; (5) underplatingand greenstone belts; (6) newly generated lower crust in granulite belts; (7) garnetized spinel lherzolites beneath granulite belts;(8) ancient lithospheric spinel peridotites; (9) ancient lithospheric spinel peridotites; (10) ancient lithospheric garnet peridotites.

Granite-greenstone terranegreenstone

belts

spinel peridotites

garnet peridotites

superplume

Granulite-gneissbelt

superplume

0

100

200

300

400

km

1 2 3 4 5

6 7 8 9 10

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IRREVERSIBLE EVOLUTION OF TECTONO-MAGMATIC PROCESSES 639 Table 5. Typical compositions of Middle Paleoproterozoic picrites and basalts in the Baltic Shield: Jatulian (2.3–2.0 Ga) andSuisarian (2.0–1.9 Ga)

ComponentJatulian Suisarian

G12/04 G13/04 G11/04 G17/04 G8/04 Yal-7 Yal-10 Chs-6

SiO2 46.08 50.03 50.34 52.38 53.71 49.46 49.01 50.36

TiO2 0.96 2.11 1.95 2.31 1.65 1.84 2.82 1.67

Al2O3 13.10 14.35 14.80 13.01 12.84 9.97 12.15 13.47

Fe2O3 11.29 13.73 15.10 14.85 14.63 12.64 13.93 12.46

MnO 0.27 0.11 0.17 0.23 0.26 0.16 0.18 0.17

MgO 19.66 7.09 5.49 4.76 6.13 14.73 6.07 7.76

CaO 4.60 6.70 6.85 6.42 5.69 10.17 12.45 9.96

Na2O 2.56 4.60 3.92 5.53 3.08 0.73 3.00 2.94

K2O 1.41 1.02 1.08 0.18 1.85 0.12 0.14 0.99

P2O5 0.07 0.26 0.30 0.33 0.16 0.18 0.25 0.22

Cr 265 63 84 40 38 735 65 136

Sc 36 34 382 37 287 – – –

V 255 384 34 370 67 149 228 216

Co 89 46 92 29 47 70 54 52

Ni 151 54 217 53 953 527 80 125

Cu 86 68 97 51 252 154 149 257

Zn 275 74 27 189 44 108 102 93

Rb 39 27 116 6 290 4 <5 19

Sr 233 113 192 81 157 83 94 345

Y 33 34 414 50 – 20 21 25

Zr 98 159 42 234 519 110 119 187

Nb 5 15 12 15 26 14 16 37

Ba 248 386 – 78 11 67 514 52

Note: Analyses were made by XRF at the Central Chemical Laboratory of the Institute of the Geology of Ore Deposits, Petrography, Min-eralogy, and Geochemistry (IGEM), Russian Academy of Sciences, analyst A.I. Yakushev.

batholiths, etc. indicate that these orogens did not differfrom Phanerozoic foldbelts produced at convergentboundaries between lithospheric plates, i.e., a new typeof tectono-magmatic activity appeared then and contin-ues until nowadays.

Thus, in the mid-Paleoproterozoic, at 2.3–2.0 Ga,the composition of the magmatic melts and tectonicprocesses changed practically simultaneously world-wide, and this was associated with important changes inthe Earth’s atmosphere, hydrosphere, and biosphere.

We believe that this change in the activity wasrelated to the ascent of mantle superplumes of the sec-ond generation (thermochemical). These plumes camefrom the boundary of the liquid core and silicate man-tle, from layer D'', similarly as they do nowadays. Thematerial of these superplumes was enriched in Fe, Ti,alkalis, P, Ba, Zr, REE, and other incompatible ele-ments, and the genesis of such superplumes is thoughtto be related to the periodic accumulation of certainfluid components that come from the liquid core to the

core–mantle interface (Dobretsov et al., 2001; Lobk-ovskii et al., 2004). The age boundary of 2.3–2.2 Gawas likely characterized by the melting of the whole (orthe bulk of) the metallic core, as follows from the max-imum intensity of the magnetic field at that time, whicheventually resulted in thermochemical superplumes.

These superplumes seem to have initially used thepathways of superplumes of the first generation, as fol-lows from changes in the composition of large mag-matic provinces of the transitional period with time (seeabove). However, as the processes intensified withtime, these superplumes started to develop indepen-dently and resulted in oceanic spreading. Because thematerial of these superplumes had a lower density andwas able to ascend to shallower depth levels, thespreading of the head portions of these superplumesresulted in active interaction with the upper parts of theancient lithosphere, including the crust: the breaking ofthe latter, origin of the oceanic crust, origin and

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2000 2200 2400 2600 2800 3000 3200Age, Ma

18001

10

100

1000

10000

Cr

0

0.1

0.4

0.5

0.6P2O5

0

2

4TiO2

7

9

17

19

21Fe2O3tot

0.1

1

10

100(La/Yb)N

0.5

1.5

3.5

4.5

5.5 (La/Sm)N

0

0.5

2.5

3.0(Nb/La)N

2000 2200 2400 2600 2800 3000 320018000

0.5

2.0

2.5

3.0

(Nb/Th)N

15

13

11

0.3

0.2

2.5

2.0

1.5

1.0

1.5

1.0

Age, Ma

Fig. 5. Changes in the concentrations (wt %) of some major components at 2.3–2.2 Ga in Archean and Paleoproterozoic mafic–ultramafic rocks of the Baltic Shield.

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IRREVERSIBLE EVOLUTION OF TECTONO-MAGMATIC PROCESSES 641

motions of plates, etc., i.e., brought about plate tecton-ics.

Inasmuch as superplumes serve as a continuous heatsink from the liquid core, the latter gradually solidifies,which should proceed (according to Jeffries’ model,see above) from below upward and lead to the appear-ance and growth of an inner (solid) core, whose compo-sition and density were different from those of the liq-uid core (Ringwood, 1978). Obviously, this isexplained by the fact that more refractory phases(including stishovite and diamond, according toS.M. Stishov, personal communication in 2007) crys-tallize at the solidification front of the core. The solidi-fication of the liquid core is associated with the releaseof significant amounts of fluids dissolved in it, and thismaintains the ascent of thermochemical superplumes, aprocess that continues for more than 2 Ga.

Earth’s Crust “Continentalization” vs. “Oceanization”

As is known, the principal morphological structuresof the modern Earth are large topographic highs withancient sialic crust and vast lows with young oceaniccrust. This topographic dichotomy is also typical of theMoon (see below) and other terrestrial planets, whichputs forth the question as to what is the general evolu-tionary trend of the Earth’s crust and the crusts of otherplanets.

In other words, what happens now with the crust:does it undergo global “oceanization” or “continental-ization”. It is known that two processes concurrentlyproceed on the Earth: (1) the development of new oce-anic crust in mid-oceanic ridges and backarc spreadingregions due to the derivation of ultramafic materialfrom the mantle and (2) the growth of newly formedoceanic crust in island arcs via the recycling of the oce-anic crust. It is, however, uncertain as to which of theprocesses is currently predominant. Most tectonistsbelieve that the continental crust is continuously pro-duced at the sacrifice of the oceanic crust in island arcs,which evolve from their young to developed and maturestates. The final stage of such evolution is the dockingof a mature island arc (and, accordingly, the newlyformed continental crust) to a continent with a corre-sponding increase in the mass of the latter.

According to another viewpoint, which was formu-lated by Belousov (1989), the aforementioned processis currently coupled with the “basification” of theEarth’s crust (i.e., the development of the oceanic crustat the expense of the continental crust). The proponentsof this viewpoint attract attention to the fact that thechemical composition of foldbelts (according to Ronovet al., 1990) is much more mafic (basalts, andesites,dacites, and rhyolites are contained in them in the pro-portion 6 : 3 : 0.5 : 0.5) than the average composition ofPrecambrian shields (their granite metamorphic shell),i.e., the sialic crust is not restored completely.

The fact that the fraction of the continental crustproduced in the Phanerozoic is only one-third of the allPrecambrian crust obviously implies that the rate ofcrust production is much lower than that of crustdestruction. From this standpoint, the Phanerozoic evo-lution is characterized by the predominance of crust“oceanization”: the ancient continental lithosphere isobliterated and is now a relict.

This is also corroborated by material balance calcu-lations. As was demonstrated above, the fraction of theEarly Precambrian basement of ancient platform isclose to 3% of the solid surface of the Earth. The bulkof its basement (including shields) is metamorphosedto the amphibolite and granulite facies, i.e., sufferedmetamorphism under pressures of 5 to 7–10 kbar (Tay-lor and McLennan, 1985; Early Precambrian Geologi-cal…, 1988). These parameters indicate that the rockswere formed at depths of 15–30 km (~20 km on aver-age), because high-pressure granulite mineral assem-blages are relatively rare. The volume of the granite–metamorphic shell of the modern continental block is(according to Ronov et al., 1990) approximately2995 km3 at an average thickness of 14 km. Thisimplies that erosion in the Late Precambrian throughPhanerozoic has obliterated approximately 3280 km3 ofthe continental crust, whereas the total volume of theEarth’s sedimentary shell (with regard for sedimentaryrocks in foldbelts) is only 1130 km3, i.e., roughly threetimes smaller.

This poses the question as to where is this material.Evidently, some of this material could be returned tothe surface in the form of mantle–crust volcanic rocks,but their fraction in the overall balance of folded andorogenic regions is insignificant and much lower thanthat of basalts, and these regions themselves play a sub-ordinate role in the continental lithosphere. Of course,more accurate material-balance calculations should becarried out, but it is hard to expect they will modify thissituation, which requires that much continental crustalmaterial is continuously removed from the tectono-sphere into the lower mantle. This is consistent with thecalculations in (McDonough, 1992), according towhich crustal material (both oceanic and continental)buried in the lower mantle accounts for 3% of theEarth’s silicate constituent, which is at least twicegreater than the mass of the modern Earth’s crust,which accounts for 1% of the mantle mass. This alsofollows from data in (Artyushkov, 1993), according towhich the volume of volcanic material accumulated inisland arcs is one order of magnitude smaller than thevolume of subducted material, i.e., the bulk of the mate-rial was irreversibly consumed in the mantle.

The systematic destruction of the ancient Earth’ssialic crust likely started at approximately 2 Ga innewly formed systems of volcanic arcs and backarcseas, in which this crust was “drawn” into subductionzones from the backarc regions, and the oceanic crustwas simultaneously taken from the forearc regions

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(Fig. 6) and was then “stored” in megaliths (“slab cem-eteries”), which were identified in the mantle by seis-mic tomography (see, for example, Karason and vander Hilst, 2000, and references therein). Obviously, thistime was marked by the generation of the secondarybasaltic crust (of the oceanic type), which currentlyaccounts for no less than two-thirds of the Earth’s solidsurface. The involvement of continental material insubduction zones also follows from data obtained bystudying ultrahigh-pressure complexes at Kokchetav inKazakhstan, Dabie Shan in China, Dora Mayra in thewestern Alps, and other complexes Norway and else-where, which are regarded as exhumed fragments ofsubducted slabs (Dobretsov et al., 2001; Ernst, 2001,and references therein). Continental rocks (metamor-phosed sediments, volcanic rocks, various gneisses andgranite-gneisses) are often preserved in a metastablestable and retain their texture, structures, and mineral-ogy in spite of metamorphism under ultrahigh pres-sures and moderate temperatures (P > 2.8–4 GPa, pos-sibly, as high as 8.5 GPa, T = 600–900°C).

If this scenario of the Earth’s evolution is realistic,the logical final of this process should be the completerestyling of the planet’s inner structure and surface,which will be dominated by the secondary crust of theoceanic type.

An overall scheme of the tectono-magmatic evolu-tion of the Earth is presented in Fig. 7.

TECTONO-MAGMATIC EVOLUTIONOF THE MOON

General information

The Moon (Fig. 8) is a planetary body approxi-mately four times smaller than the Earth (its diameter is3476 km, i.e., 0.27 of the Earth’s diameter), and itsmass is 8 × 1019 t (0.0123 of the Earth’s mass). Judgingfrom seismic evidence, the Moon has a metal core 340–

700

500

400

300

200

100

600

1 2 3 4 5

6 7 8 9

a ba b

km

I II

III

Fig. 6. System volcanic arc–backarc sea.(1) Material of young low-density (“asthenospheric”) man-tle beneath oceans; (2) material of lithospheric mantle:(a) beneath continents, (b) beneath oceans; (3) material ofthe upper mantle below the discontinuity at 450 km;(5) material of the lower crust: (a) beneath continents,(b) beneath oceans; (6) sialic continental crust; (7) mixtureof sialic and basic crustal material in a subduction zone;(8) magma generation regions; (9) flows of low-densityoceanic upper mantle.Numerals: (I) rocks of the tholeiite series; (II) rocks of theboninite series; (III) rocks of the calc–alkaline series.

Nuclear stage Cratonicstage

Continental-oceanicstage

Ga

Komatiite-basalt series

Geochemical analogs

Mainly depleted mantleMain plumetectonics

Siliceous high-Mg series

Phanerozoic type of magmatism

Enriched and depleted mantle

Mainly plate tectonics

Ear

th f

orm

atio

nM

agm

a oc

ean

Appearenceof liquid

core

Completelyliquidcore

4.6–6.5(?) 4.0 3.0 2.7 2.5 2.8 2.0 1.8 1.0 0

of boninites

Fig. 7. Scheme of the tectono-magmatic evolution of the Earth.

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IRREVERSIBLE EVOLUTION OF TECTONO-MAGMATIC PROCESSES 643

450 km in radius (Konopliv et al., 1998). The Moon’scenter of gravity is shifted from geometric centertoward the Earth, and because of this the Moon alwaysfaces the Earth with the same side (Moon’s nearside).

Similarly to the Earth, which has continents andoceans, the Moon has two major types of its morpho-structures: highlands (“continents”) and maria(“oceans”). The Moon’s surfaces is quantitatively dom-inated by highlands, which occupy practically whole ofits farside. Highlands are heterogeneous: their portionsdistant from maria are characterized by flattenedtopography with large ring structures and topographiclows in between. The largest of the latter are referred toas talassoids, which are large depressions in highlands,whose sizes approach those of maria but which consistof the highland material. The largest tallasoid is theSouth Pole-Aitken Basin, having a depth from 10 to4 km and devoid of mare material.

Rounded maria are younger and occur mostly at thenearside. The total thickness of their lava sequences isusually a few kilometers. Maria are large topographiclows, whose surface lies a few kilometers lower thanthat of the surrounding highlands (Spudis, 1996). Theseterritories usually have a plain topography of muchsimpler structure than at neighboring highlands. Thesemaria sometimes have a concentrically zonal structure:they include two to four ring zones separated by walls

(Valis Schrödinger and Mare Orientale, respectively),which suggest multiple pulses of magmatic activity atthese magmatic centers.

Boundaries with highlands are often accentuated bymarginal highs (lunar mountains), which were discov-ered by Galileo and form arcs around maria. Thesemarginal highs include exposed blocks of deep rocksfrom the upper and lower crust. The vertical span oflunar topographic features reaches 16 km, with thehighest topographic elevations restricted to the farsideof the planet.

The thickness of the lunar crust ranges from 25–35 km in maria to 90–110 km in highlands on the far-side, at an average thickness of 60–70 km (Zharkov,1978; Spudis, 1996). Mass concentrations (so-calledmascons) were detected beneath the thinned crust inmaria. This resembles the distribution of crustal thick-ness beneath oceanic and continental segments on theEarth (at an increase by a factor of two to three).

Characteristics of the Moon’s Tectono-Magmatic Evolution

Data obtained on samples of lunar soil recovered byAmerican and Soviet missions to the Moon (approxi-mately 385 kg of regolith) indicate that the oldest mag-matic activity took place on the Moon at 4.0–4.0 Ga,

A-16

L-16

L-16

A-14A-12

A-11

A-17

A-15

Fig. 8. Moon’s nearside.Dark areas are maria. Rectangles with characters show the landing sites of Luna (Soviet Union) and Apollo (United States) spacemissions.

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and its products are preserved at highlands. These arevolcanic rocks of a low-Ti magnesian series and theirplutonic analogues: layered mafic–ultramafic intru-sions of the ANT series (ANT series: Anorthosite–Norite–Troctolite) (Snyder et al., 1995a). A fragment ofsuch an intrusion is exposed at the Moon’s surface as atectonic nappe in the marginal wall (Silver Spur Range)between the Mare Imbrium and Mare Serenitatis, nearthe landing site of Apollo 15. The chemistry (Table 6),mineralogy, geochemistry, and isotopic composition ofthe volcanic and cumulate rocks of the layered intru-sions are close to those of rocks of the silicic high-Mgseries produced by Early Paleoproterozoic magmatismon the Earth but differ from them in bearing reduced

mineral phases (Sharkov and Bogatikov, 2001; Mokhovet al., 2007, and references therein).

Another type of products of lunar highland magma-tism is medium-Ti basalts enriched (compared to otherlunar rocks) in K, REE, P, and incompatible elements(Zr, Nb, U, Th, La, Ba, Rb, and others) and their intru-sive analogues. Judging from Lunar Prospector data,rocks of this series may compose significant areas in theOceanus Procellarum. According to isotopic geochro-nologic data, these rocks were continuously producedfor at least 300 Ma: from 4.34 to 4.0 Ga (Snyder et al.,2000). Magmatism during this time span generallyresembles magmatism during the transitional stagebetween the cratonic and continental–oceanic stages on

Table 6. Major and trace elements in rocks from lunar highlands

Component 1 2 3 4 5 6 7 8 9

SiO2 48.18 47.49 43.75 47.98 48.15 48.8 46.6 51.4 50.7

TiO2 0.22 0.61 2.70 1.80 1.77 0.69 0.64 2.2 1.9

Al2O3 7.06 10.05 8.34 9.44 9.44 10.0 19.5 15.7 15.4

Cr2O3 0.60 0.51 0.55 0.48 0.63 0.69 0.20 0.31 0.33

FeO 16.30 17.61 21.81 20.23 19.98 17.9 8.6 10.0 9.8

MnO 0.34 0.10 0.30 0.30 0.30 0.30 0.12 0.15 0.14

MgO 18.66 14.54 12.63 8.74 8.85 11.8 10.9 9.0 9.6

CaO 8.43 9.43 8.62 10.43 10.58 9.4 12.6 9.8 9.4

Na2O 0.17 0.34 0.54 0.32 0.37 0.06 0.42 0.79 0.75

K2O 0.00 0.06 0.00 0.06 0.06 0.02 0.14 0.60 0.55

Total 100.46 100.77 99.24 99.78 100.03 99.7 99.72 100.51 99.13

Sc, ppm 37.2 35.7 36.1 – 47 – 21.2 20.2

V 158 160 137 – 130 – – –

Co 47.3 66.3 49.4 39.6 44.6 – 2100 2200

Ni 90 125 85 30.0 8.9 – 19.0 19.8

Sr 12.7 50.7 139 111 111 – 180 200

Ba 7.9 120.2 145 52 45.2 – 780 720

La 0.76 – 22.1 4.86 4.01 – 74 69

Ce 1.93 29.5 57.8 13 13.1 – 190 180

Nd 1.05 29.5 30.9 9.3 8.87 – 110 105

Sm 0.34 16.8 10.2 3.09 2.93 – 33.0 30.3

Eu 0.06 0.30 1.44 0.84 0.48 – 2.70 2.65

Dy 0.61 5.3 9.95 4.51 4.59 – 42 38

Er 0.53 3.2 5.8 1.9 2.7 – – (25)

Yb 0.53 3.0 5.1 2.13 2.35 – 21.8 20.6

Zr 12.0 201 285 280 104 – 980 900

Hf – – – – 2.14 – 23.9 23.2

Th – – – 0.53 0.4 – 12.8 11.7

U – – – 0.14 0.12 – – –

Note: Analyses: (1, 2) droplets of green picrite glass (Apollo 15 and 14); (3) orange glass (Apollo 14); (4, 5) pigeonite basalt (Apollo 15);(6) very low-Ti (VLT) olivine basalt (Apollo 17); (1–6—Papike et al., 1998); (7) high-Al basalt (Luna 20; Prints et al., 1979);(8, 9) KREEP basalts (Snyder et al., 1995b).

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the Earth. Note that the same time (4.2 Ga) was markedby the appearance of the Moon’s magnetic field (Ran-corn, 1983).

At approximately 3.9–3.8 Ga, this type of magmaticactivity gave way to mare basaltic magmatism, whichoccurred simultaneously with the development of largelunar mare depressions and mountain ranges (montes).Most researchers believe that mares were produced as aconsequence of catastrophic impact events at ~3.9 Ga(so-called lunar cataclysm; Spudis, 1996, and refer-ences therein). However, mares occur only at theMoon’s nearside (facing the Earth), which is character-ized by a thinner crust. This differences between thestructures of the hemispheres and the eccentricity of theMoon’s center of gravity rather suggest that the Earth’sgravitation played an important role when the primarylunar crust and its inner structure were formed and,later, when large maria developed on the planet (it ishard to explain why giant meteorites should have comeexclusively from the Earth-side). In spite of their rela-tively small sizes, the structures of maria themselvesmost closely resemble those of terrestrial oceans orflood basalt provinces (signiifcantly thinned crust andpowerful basaltic volcanism).

This type of magmatism continued until approxi-mately 3 Ga. As on the Earth, the mare basalts are clas-sified into two types: low- and high-Ti. Their geochem-istry (Table 7) is correlated with that of terrestrialMORB (mid-ocean ridge basalts) and OIB (oceanicisland basalts) (Sharkov and Bogatikov, 2001). Theserocks principally differ from their terrestrial analoguesin having low alkalinity, containing no water-bearingminerals, and the presence of native iron, Fe–Ni alloys,and other reduced phases (Papike et al., 1978; Mokhovet al., 2007, and references therein). These data testifythat magmas on the Moon were derived under morereduced conditions than on the Earth.

By analogy with the Earth, we believe that lunarhighland magmatism was initiated in relation with the

ascent of mantle plumes of the first generation, whichconsisted of the material of the depleted mantle. In thiscase, highland uplifts and talassoids (see above) may beanalogues of terrestrial Early Paleoproterozoic cratonsand granulite belts. Lunar mare magmatism was likelyinduced by the ascent of mantle thermochemicalplumes of the second generation, whose developmentwas initiated at the boundary between the mantle andliquid metal core. The magmatic activity started atabout 4.2 Ga, whereas the intensity of the magneticfield reached a maximum only at approximately 3.9 Ga(when it was as high as 1 Gs) and gradually decreaseduntil 3 Ga (Rancorn, 1983).

Starting at that time, maria could, indeed, be ana-logues of terrestrial oceanic and flood basalts but couldnot be produced by impacts of large meteorites. As onthe Earth, the material of these plumes was lighter thanthat of earlier plumes, and hence, the spreading of theirhead portions occurred at the base of the lunar crust andresulted in its significant transformations and the devel-opment of both lunar mare depressions and surround-ing mountain ranges (montes), where excess crustalmaterial was stacked (Fig. 9). Mascons (concentrationsof dense masses) ubiquitously identified beneath mariaare likely the solidified head portions of these plumes.As on the Earth, manifestations of mare magmatismcoincided with the peak intensity of the magnetic field,which took place on the Moon at 3.9–3.0 Ga.

Magmatism on the Earth and Moon is compared inFig. 10.

DISCUSSIONThe materials and data presented above imply that,

in spite of significant differences in their sizes, theEarth and Moon have similar inner structures and com-positions and evolved according to similar scenarios,although the latter was significantly briefer for theMoon. This is consistent with the ideas of someresearchers (Wasserburg et al., 1977; Galimov, 1995)

0

20

40

60

80

100

km

low-Ti basaltshigh-Ti basalts high-Ti basalts

1 2 3 4 5

Fig. 9. Scheme illustrating the genesis of lunar maria.(1) Mare basalts; (2) mantle plume of the second generation; (3) upper crust; (4) lower crust; (5) lithospheric mantle.

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that the Earth–Moon pair is an example of a duplex sys-tem. Isotopic data testify that these bodies were formedfrom the same starting material (Nemchin et al., 2006).It is commonly believed that this material was a pro-toplanetary gas–dust nebula, from which the Earth, as amore massive body, could more efficiently “extract”volatile components, first of all, water. This likelyexplains the depletion of lunar material in these compo-nents and its enrichment in refractory components, as

this follows from geochemical data (Kuskov and Kro-nrod, 1999). Correspondingly, the Moon could hardlybe produced of the material of the terrestrial mantle asa result of the catastrophic collision between the Earthand a body of Martian size, as this was proposed bysome researchers (Spudis, 1998, and referencestherein). Hence, available petrological–chemical datasuggest that both planets were formed as independentbodies of somewhat different composition.

Table 7. Major and trace elements in rocks from lunar maria

Component12006 12017 21b 25c B3 A B2 D

1 2 3 4 5 6 7 8

SiO2 44.23 47.27 38.25 34.54 42.2 40.7 40.0 41.4

TiO2 2.59 3.37 10.17 16.70 10.3 11.0 8.93 8.43

Al2O3 7.67 10.0 4.46 4.63 9.45 8.20 11.5 11.4

FeO 20.94 19.72 24.32 21.86 17.6 20.2 19.6 19.2

MnO 0.29 0.29 0.26 0.30 0.25 0.23 0.26 0.26

MgO 14.67 7.63 14.84 13.12 8.50 8.01 7.57 7.33

CaO 8.13 10.97 7.23 7.59 11.2 10.5 11.2 11.5

Na2O 0.20 0.27 0.29 0.21 0.34 0.52 0.43 0.38

K2O 0.05 0.09 0.00 0.10 0.05 0.30 0.07 0.09

P2O5 – – – – 0.07 0.19 0.11 0.21

Total 98.77 99.61 100.50 99.84 99.96 99.85 99.67 100.20

Sc 40.1 52.8 54 58 76.5 82.2 76.7 78.0

V – – 122 295 106 71.8 70.7 97.0

Cr 6250 3550 4926 5611 2813 2258 1505 2258

Co 60 32 59 38 17.5 27.7 14.3 16.7

Ni 110 – 30 – – – – –

Sr 89 118 147 255 – 164 160 –

Rb – – – – – 5.68 0.63 –

Ba 56 75 48 270 60 297 130 203

La – – 4.14 15.4 5.70 26.5 15.9 33.1

Ce 15.7 – 13.0 44.9 23.0 78.5 49.7 93.0

Nd – – 13.3 38.9 23.0 65.3 43.0 75.3

Sm 3.77 5.1 5.85 2.35 8.10 20.9 14.6 23.3

Eu 0.72 – 1.49 14.3 1.50 2.24 1.89 1.93

Tb 1.02 – – – 1.96 4.61 3.21 4.73

Dy – – 9.9 18.3 13.0 31.2 21.4 31.7

Yb 3.3 4.4 4.48 7.50 7.65 17.3 11.8 16.8

Lu 0.47 0.66 – – 1.14 2.49 1.64 2.39

Zr 97 – 206 399 – – – –

Nb 6.4 – – – – – – –

Hf 3.0 – – – 6.4 16.5 10.1 12.6

Ta – – – – 1.5 2.4 1.9 1.7

Th – – – – 0.3 3.1 1.2 2.5

Note: (1, 2) Apollo 12 mission, low-Ti basalts: (1, 2) olivine basalts, (3, 4) picrite glass, Apollo 11 and 12 missions, respectively;(5−8) Apollo 11 mission, high-Ti basalts: (5) “primitive” basalt B3, (6–8) basalts of groups A, B2, and D.

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An exceptionally important fact is a drastic changein the evolution of tectono-magmatic processes duringthe intermediate evolutionary stages of the Earth andMoon, when high-Mg magmatic melts gave way togeochemically enriched Fe–Ti basalts and picrites withelevated concentrations of incompatible elements, andthen the character of tectonic processes has alsochanged. Data presented above show that the early evo-lutionary stages of the planetary bodies were character-ized by the development of mantle superplumes withinthe depleted mantle, whose depletion in incompatibleelements intensified with the passage of time. The onsetof the next stages was marked by the appearance of liq-uid cores, which maintained the existence of mantlethermochemical superplumes of the secondary genera-tion. These plumes were formed at the boundary of theliquid metallic cores and silicate mantles. Their ascentbegan ~2.5 Ga after the origin of the Earth and ~1.5 Gaafter the origin of the Moon and has fundamentallychanged the style of tectonic processes in the uppershells of these planets.

At principal differences in tectono-magmatic pro-cesses before and after this change, these processes sys-

tematically and irreversibly evolved for a long time dur-ing each of the stages. At the Earth, and likely, also atthe Moon (KREEP basalts), transitional period existedand lasted for ~300 Ma. This likely suggests that thedriving forces maintaining the evolution of both plane-tary bodies stemmed from their interiors and were notexternal.

Hence, a principally important feature of the evolu-tion of both the Earth and the Moon was a relativelyquick change in the character of the melted mantlesource material from depleted to geochemicallyenriched, which implies that principally other types ofmaterial were involved in tectono-magmatic processesduring the intermediate evolutionary stages of the plan-ets and geological catastrophes.

This poses the questions as to where this materialwas stored, how was it activated, and why this led tosuch consequences. This scenario of the events couldbe ensured only by a combination of two independentfactors: (1) the planetary bodies were initially heteroge-neous, i.e., were produced by heterogeneous accretion,and (2) their heating proceeded from their upper levelsinward, i.e., from the surface to core, and was associ-

0

1.5

4.5

4.0

3.5

3.0

2.5

2.0

MoonEarthPrimordial crust

Headen

Nuclearicstage

Cratonicstage

Continental-oceanic stage

Continental stage

Maria stageMare magmatism

Magnesian suiteKREEP-basalts

1

2

3

4

5

Ga

Fig. 10. Comparison of the evolution of tectono-magmatic processes on the Earth and Moon.(1) Granite–greenstone terranes at the Earth with mantle komatiite–basalt and boninite-like magmatism; (2) silicic high-Mg seriesof the Earth and magnesian series of the Moon; (3) Fe–Ti picrites and basalts at the Earth and KREEP basalts at the Moon; (4) Phan-erozoic type of tectonic activity at the Earth; (5) mare magmatism at the Moon.

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ated with the cooling of the upper shells. Only thismodel is able to explain why this material has been“conserved” for such a long time.

The global drastic change in the tectono-magmaticevolution of the planetary bodies (this change wasrelated to the substitution of depleted sources forgeochemically enriched ones) should not be possibleduring homogeneous accretion, because this scenarioimplies the opposite succession of the events. Thisplaces serious limitations onto the hypothesis of plane-tesimals, which is so popular among astrophysicists butdoes not take this issue into account.

The systematic directional heating of the planetarybodies is a more complicated process. Although thisproblem is still very poorly explored, it is quite clearthat the evolution of tectono-magmatic processes on theEarth cannot be explained without invoking a “heatwave” passing from the peripheral to inner parts of theplanetary body. This “wave” should have graduallypassed inward the newly formed planet and heated pro-gressively deeper levels in its mantle. It is not necessarythat the “wave” induced melting of material in theEarth’s inner shells. It could merely initiate the ascentof mantle superplumes of the first generation with adi-abatic melting in their head portions. The core was thelast to be affected by the “wave”, and hence, the corematerial, which still maintains the development of ther-mochemical plumes on the Earth, was the last to berecycled. After the core was completely (or mostly)melted (the maximum magnitude of the magnetic fieldwas reached at 2.3–2.2 Ga and gradually decreasedafter this), the situation stabilized. An analogous transi-tion occurred at the Moon at approximately 3.9 Ga.Evidently, the passage of the “energy wave” inward theplanet was associated with the cooling of its outershells, and hence, the existence of an atmosphere at theEarth became possible as long ago as the Eoarchean.

It can be hypothesized that such a scenario of theevents was caused by a centripetal passage of a wave ofdeformations (energy wave), which was experimentallymodeled in rotating bodies (Belostotskii, 2000). Thisenergy transfer is the most active during the accelera-tion of rotation and is practically absent at a stationaryrotation regime. Inasmuch as deformations are alwaysaccompanied by heat release, this “wave” should, infact, have been a heating wave.

This means that the probable reason for the revealedsuccession of events was the passage of a heating wave(zone) from the peripheral to inner parts of the planetsafter their formation was completed. In this situation,the termination of accretion in these bodies during theirfirst evolutionary stage was immediately followed bythe acceleration of their rotation, until the modern rota-tion regime was reached after ~2.5-Ga evolution of theplanets.

Conceivably, this “energy pumping” into the Earth’sinner shells explains the principal difference in the tem-perature regime between the mantle and core: while the

temperature at the lower mantle boundary at depths of~2900 km is 2900°C, the temperature at the surface ofthe solid core is ~5000°C (McDonough, 2003). Thegenesis of the liquid core is usually explained by therelease of much potential gravity energy and its primarydissipation through the whole Earth (Ringwood, 1978).As follows from the data presented above, this canhardly be the case with the Earth. The resolution of thisproblem evidently calls for further research in the fieldof mechanics of celestial bodies and extends outside thescope of our research.

Now superplumes of the second generation deter-mine practically all tectono-magmatic activity at theEarth, maintaining mantle convection and serving asthe main driving force of tectonic processes. Thismeans that liquid metallic cores of planetary bodies aretheir energetic “hearts” during the secondary, finalstages of their evolution. Upon the solidification ofthese cores, tectono-magmatic processes cease to oper-ate, as is the case with the Moon.

The occurrence of two major types of morphostruc-tures on Venus and Mars (extensive areas covered withbasalt flows and old and more elevated territories ofcomplicated topography, so-called tesserae on Venusand terra and planitia on Mars) suggests that theseplanets were also formed in two stage. During the ear-lier one, the primary lithosphere of the planets was pro-duced via the solidification of their global magmatic“oceans” and the activity of mantle plumes of the firstgeneration. The second stage was marked by dynamicprocesses associated with extensive basaltic magma-tism related to the ascent of thermochemical super-plumes from the boundaries of the liquid cores andmantles. Mercury is the least know planet, whoseradius is trice smaller than that of the Earth’s but onwhich two types of morphostructures were distin-guished: those resembling lunar highlands and maria(Ksanfomality, 2008).

Practically no magnetic field generated by a liquidcore is present on Venus and Mars (Marov, 1981),which are also devoid of modern volcanism. This sug-gests that the endogenic evolution of these planets iscompleted, and they are, similarly to the Moon,“extinct” bodies that gradually lose the rest of theirenergy to space. The situation with Mercury is not asclear, because this planet still has a magnetic field(which is possible dipolar) but, likely, no liquid core(Ksanfomality, 2008). Conceivably, this can beexplained by the unusually large size of Mercury’score: the ratio of the core’s radius to the surface is 0.8,the highest among those of all terrestrial planets.

CONCLUSIONS

The crusts of the Earth, Moon, and, likely, also otherterrestrial planets started developing during the solidifi-cation of their global magmatic “oceans”, a process thatproceeded from below upward due to a difference in the

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adiabatic gradient and the gradient of the melting point.As a result of this process, the lowest melting compo-nents accumulated during melt crystallization andenriched the near-surface levels of the planets. This ledto the origin of the primary Earth’s sialic crust and theMoon’s mafic crust because of the different depths ofthe magmatic “oceans”. The same processes wereresponsible for the primary depletion of the mantle ofthese planets in low-melting components. The compo-sition of the primary crusts of other terrestrial planets isstill uncertain.

All of the planetary bodies in question are character-ized by the occurrence of the crust of two types: old“continental” crusts, which composed elevated areas,and younger basite (basaltic) ones, which makes upextensive lowland areas at large planets and equant“seas” at smaller ones (Mercury and the Moon). Judg-ing from data on magmatism on the Earth and Moon,the typical melts of the initial stages of their evolutionwere rich in Mg and were derived from depleted mantlematerial. The secondary evolutionary stages were char-acterized by melts rich in Fe, Ti, and incompatible ele-ments, which were melted out of geochemicallyenriched sources.

The intermediate evolutionary stages of the Earthand Moon were marked by an irreversible drasticchange in the course of their evolution because of theappearance of liquid metallic cores (which is consistentwith the maximum strength of the magnetic fields at2.3–2.2 and 3.9 Ga on the Earth and Moon, respec-tively). This initiated the ascent of mantle superplumesof the second generation: thermochemical super-plumes, whose material was enriched in fluid compo-nents. The head portions of these superplumes spread atrelatively shallow depths and notably transformed theupper shells of the planets, leading to the gradualreplacement of the primary crusts by secondary ones(basaltic).

The change in the character of tectono-magmaticactivity at the Earth in the Middle Paleoproterozoic wasassociated with a change in the ecological environmentat the planet’s surface, including the development of anoxygen-bearing atmosphere, hydrocarbons, phospho-rites, changes in the water chemistry in the ocean andthe isotopic composition of carbonate sediments; theseprocesses were accompanied by the origin of multicel-lular organisms.

The facts and considerations presented above sug-gest that the intermediate evolutionary stages of theplanetary bodies were marked by the involvement of aprincipally new type of material in tectono-magmaticprocesses. This material was previously “conserved” atdeep levels of the planets, which could be possible onlyif the planets initially had a heterogeneous structure(with a metallic core and silicate mantle, which con-sisted of the material of carbonaceous chondrites), andthe heating of the planets proceeded inward and wasassociated with the cooling of the outer shells. As of

now, the Earth’s core is partly solidified, and the coresof the Moon, Mars, and Venus are completely solid, asfollows from the absence of magnetic fields and mod-ern volcanism at these planets.

The centripetal heating of the planetary bodies isthough to have been related to an energy wave, whichwas generated after the planets had been completelyformed during the acceleration of their rotation beforethe onset of stable rotation during intermediate stagesof their evolution.

Available data suggest that terrestrial planets areself-evolving systems, whose evolution was associatedwith irreversible changes in tectono-magmatic pro-cesses and the composition of the atmospheres. Theevolution of all of them except the Earth, the largest ofthe planets, is now completed, and they are “extinct”bodies that gradually lose their energy to space.

ACKNOWLEDGMENTS

This study was financially supported by the RussianFoundation for Basic Research (project no. 07-05-00496), Project ONZ RAN 4, and program for the sup-port of leading research schools.

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