The Geoghemical Evolution of the Continental Crustal

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THE GEOCHEMICAL EVOLUTION OF

THE CONTINENTAL CRUST

Stuart Ross Taylor

Department of Nuclear Physics

Research School of Physical Sciences

Australian National University, Canberra

Abstract. A survey is given of the dimensions and

composition of the present continental crust. The

abundances of immobile elements in sedimentary

rocks are used to establish upper crustal composition.

The present upper crustal composition is attributed

largely to intracrustal differentiation resulting in the

production of granites senso lato. Underplating of the

crust by ponded basaltic magmas is probably a major

source of heat for intracrustal differentiation. The con-

trast between the present upper crustal composition

and that of the Archean upper crust is emphasized.The nature of the lower crust is examined in the light

of evidence from granulites and xenoliths of lower

crustal origin. It appears that the protoliths of most

granulite facies exposures are more representative of

upper or middle crust and that the lower crust has a

much more basic composition than the exposed upper

crust. There is growing consensus that the crust grows

episodically, and it is concluded that at least 60% of

the crust was emplaced by the late Archean (ca. 2.7

eons, or 2.7 Ga). There appears to be a relationship

between episodes of continental growth and differen-

tiation and supercontinental cycles, probably dating

back at least to the late Archean. However, such

cycles do not explain the contrast in crustal composi-

tions between Archean and post-Archean. Mecha-

Scott M. McLennan 1

Department of Earth and Space Sciences

State University of New York, Stony Brook

nisms for deriving the crust from the mantle are

considered, including the role of present-day plate

tectonics and subduction zones. It is concluded that a

somewhat different tectonic regime operated in the

Archean and was responsible for the growth of much

of the continental crust. Archean tonalites and trond-

hjemites may have resulted from slab melting and/or

from melting of the Archean mantle wedge but at low

pressures and high temperatures analogous to modem

boninites. In contrast, most andesites and subduction-

related rocks, now the main contributors to crustalgrowth, are derived ultimately from the mantle wedge

above subduction zones. The cause of the contrast be-

tween the processes responsible for Archean and post-

Archean crustal growth is attributed to faster subduction

of younger, hotter oceanic crust in the Archean (ulti-

mately due to higher heat flow) compared with subduc-

tion of older, cooler oceanic crust in more recent times.

A brief survey of the causes of continental breakup

reveals that neither plume nor lithospheric stretching is a

totally satisfactory explanation. Speculations are pre-

sented about crustal development before 4000m.y. ago.

The terrestrial continental crust appears to be unique

compared with crusts on other planets and satellites in

the solar system, ultimately a consequence of the abun-

dant free water on the Earth.

INTRODUCTION

The familiar crust of the Earth is of unique impor-

tance because it formed the platform above sea level

onwhich the later stages of evolution occurred that led

to the appearance of Homo sapiens. If the continental

crust had not provided the mountains, broad plains,

and forests that encouraged the development of the

diversity of land-based life, nonmarine life, restricted

to small islands, would have taken a different course;

birds might have become the dominant land dwellers,

as in prehuman New Zealand. It is thus of interest to

IOn sabbatical in 1994 at Abteilung Geochemie, Max-

Planck-Institut fur Chemie, Mainz, Germany.

inquire how the continental crust, which differs so

greatly in composition from the oceanic crust and the

mantle, came into existence, and whether similar

crusts exist on the other terrestrial planets.

The continental crust comprises only about 0.35%

of the mass of the Earth. Although it might thus appear

to be a trivial geochemical reservoir, the crust contains

substantial amounts, exceeding 30% of the bulk Earth

budget, of several of the most incompatible elements,

such as Cs, Rb, K, U, Th, and La. Because the crust

is not easily recycled back into the mantle, it is a major

geochemical reservoir, and its composition is an im-

portant constraint on all geochemical models of bulk

Earth composition and evolution. Any significant tem-

poral change in either the crust's composition or its

mass will have important implications for understand-

Copyright 1995 by the American Geophysical Union.

8755-1209/95/95RG-00262$15.00

Reviews of Geophysics, 33, 2 / May 1995

pages 241-265

Paper number 95RG00262

.241 •



242 • Taylor and McLennan: CONTINENTAL CRUST GEOCHEMISTRY 33, 2 I REVIEWS OF GEOPHYSICS

ing the evolution of the Earth. Perhaps equally impor-

tant in this context are the details of continental crustal

composition, especially the variations in composition

between upper and lower crust, which provide insights

into the mechanisms of crust formation and differen-

tiation.

The continental crust is especially diverse and het-

erogeneous. Major compositional changes occur,

sometimes on a scale of meters; everyone is familiar

with the complexity displayed on geological maps of

all scales. Arriving at a reliable average composition

might be thought to be difficult, if not impossible.

Nature, however, through the processes of erosion

and sedimentation, has carried out its own overall

sampling of the exposed crust, and this information is

stored within the sedimentary rock record. Some ele-

ments possess the desirable property of having ex-

tremely low concentrations in natural waters and so

are not readily partitioned into fluids during weather-

ing or diagenesis. Such elements, of which the rareearth elements (REE) are a prime example, are thus

transferred quantitatively during erosion and sedimen-

tation from the parent rocks into clastic sediments.

The average composition of the bulk, rather than

the exposed crust, is a far more complicated issue.

Accordingly, there is considerably less of a consensus

about both bulk crustal composition and mechanisms

of crust formation. Constraints from geochemistry,

though model dependent, provide estimates of the

average K, Th, and U content of the bulk crust. In

turn, these data limit models of bulk crustal composi-

tion. The important constraint on K, Th, and U abun-dances in the continental crust relies heavily on the

interpretation of continental heat flowdata, and as will

be noted below, this is proving to be a difficult issue.

Integration of the sedimentary and geophysical

records has helped to resolve the sampling problem

and provided detailed information about the composi-

tion and evolution of the continental crust.

Geological, petrological, geochemical, and geo-

physical observations each produce important con-

straints on the composition and evolution of the con-

tinental crust and combine to allow some broad

generalizations. Thus the composition of the upper

crust cannot persist below about 10 km, so that the

lower crust is in many ways a distinct geochemical

entity. There is much evidence to suggest a fundamen-

tal difference between the crust generated during the

early and middle Archean with that which developed

during and subsequent to the late Archean. The nature

of the earliest crust, though the subject of intense

interest, remains largely a mystery. There has been a

continuing debate over the growth rate of the crust,

the scale of recycling of the crust into the mantle, and

of course its ultimate origin.

Comprehensive reviews of the composition and

geochemical evolution of the continental crust havebeen provided elsewhere [Taylor and McLennan,

1985; Taylor, 1989; McLennan and Taylor, 1991;

McLennan, 1992]. In this paper our purpose is to

provide a general overview of recent developments in

continental crustal geochemistry and to indicate some

of the more promising areas for continuing research.

GLOSSARY

Archean-Proterozoic transition: The time span of

about 3.2 to 2.5 eons (1 eon = 1Ga) that incorporates

the major tectonic stabilization (cratonization) near or

at the end of the Archean. The transition is generally

short-lived on any given craton but is found at different

times in different places. This term is distinguished

from the Archean-Proterozoic boundary that is com-

monly dated at precisely 2.5 eons.

Ce anomaly: Enrichment (positive anomaly) ordepletion (negative anomaly) of cerium compared with

the expected level predicted from a smooth, chondrite-

normalized REE pattern.

ES: European shale composite (a composite of

European shales of Paleozoic age).

Eu anomaly: Enrichment (positive anomaly) or

depletion (negative anomaly) of europium compared

with the expected level predicted from a smooth,

chondrite-normalized REE pattern.

Eu/Eu*: The ratio of measured Eu divided by the

predicted concentration of Eu (Eu*) predicted from a

smooth, chondrite-normalized REE pattern. Eu/Eu*> 1.0 indicates positive Eu anomalies, and Eu/Eu* <1.0 indicates negative Eu anomalies.

First-cycle sediments: Sediments that have been

derived directly from igneous source rocks without an

intervening history of crustal and/or sedimentary re-

cycling.

GdN/YbN: The ratio of chondrite-normalized val-

ues (denoted by subscript N) of Gd and Yb.

HREE: The heavy (or small) rare earth elements,

Gd through Lu.

LaN/YbN: The ratio of chondrite-normalized val-

ues (denoted by subscript N) of La and Yb.

LlLE: Large-ion lithophile elements (e.g., Cs, Rb,

Th, U, LREE).

LREE: The light (or large) rare earth elements, La

through Sm.

Nd model age: The age calculated from the

neodymium isotopic composition of a rock or mineral

when the initial isotopic composition is assumed.

NASC: North American shale composite (a com-

posite of about 40 post-Archean shales and slates,

mainly from North America).

PAAS: Post-Archean average Australian shale

(the average REE abundances in 26 post-Archean

shales from Australia).REE: Rare earth elements. In most geochemical



33,2/ REVIEWS OF GEOPHYSICS Taylor arid McLennan: CONTINENTAL CRUST GEOCHEMISTRY. 243

TABLE 1. Dimensions of Continental Crust

Mean Mean

Area, Thickness, Volume, Height, Flooded

Continent lrf' km2 km 109 km' m Area, %

Major Continents

Eurasia 62.160 38 2.362 361 25.8

North America 39.793 35 1.393 96 30.6Africa 35.722 24 0.857 244 14.5South America 23~982 42 1.007 149 19.0Antarctica 18.984 33 0.626 -344 78.4Australia 14.163 36 0.510 -244 34.9india 4.648 38 0.177 236 14.9Arabia 4.595 23 0.106 470 10.6New Zealand 3.980 21 0.084 -1568 90.3Central America 1.344 18 0.024 -579 52.8

Submerged Microcontinents

Rockall 0.459 33 0.015 -1702 100Seychelles 0.386 32 0.0.12 -1706 99.9Agulhas 0.i34 12 0.002 -3222 100:0Jan Mayen 0.055 0.002* -1600 99.4

Totals 210.405 7.177t 126 30.6

Adapted from Cogley [1984].

*Value for assllmed crustal thickness between 28 arid 45 km."Sum of individual volumes. Product of total area and average thickness of all continents leads to volume of 7.575 x 109km".

discussions, the term is restricted to the lanthanides

La through Lu, but it also includes Y and Sc.

TDM: Nd model age that assumes an initial Nd

isotopic composition equivalent to a model for the

mantle (various models are used) that has been de-

pleted in LREE for much of its history.

TTG suite: A common suite of Na-rich felsic plu-

tonics (tonalite-trondhjemite-granodiorite) found com-monly in Archean terranes and characterized by,

among other things, depletion in the heavy rare earth

elements. Volcanic equivalents are also found and

generally included in discussions of the TTG suite.

S IZE , STRUCTURE , AND AGE

The continental crust occupies 41.2% of the surface

area of the Earth, or 2.10 x 108 knr'; 71.3% or 1.50 x

108 km2 lies above sea level. There are at least 10

major continental blocks and four submerged micro-continents (Table O . The mean elevation of the conti-

nental crust above sea level is 125m. The elevation of

the area which lies above the 200-m isobath (i.e., the

shelf-slope break) is 690 m. The average elevation of

the continents above the mean seafloor (oceanic crust)

is about 5 km. Average crustal thickness is.36 km but

varies between 10 and SOkm, correlating with the size

of the continental block and the age of the last tectonic

event. The crustal volume is 7.35 x 109 knr', a value

that includes sediment on the ocean floor derived from

the continents and the volumes of the submerged con-

tinental masses. This estimate of volume has an error

of at least ±5% because it depends on the value

adopted for crustal thickness. This is not known with

great accuracy; some estimates of average crustal

thickness are as high as 40 km.

The base of the crust is usually defined as the major

seismic discontinuity, the Mohorovicic discontinuity

or Moho, where compressional wave velocities (Vp)

jump from about 7 to about S km s-1. However, the

Moho is not a sharp boundary everywhere and may be

absent locally. Thus the crust is often defined as ma-

terial with Vp < 7.S km S-1 or seismic shear wave

velocities Vs < 4.3 km s-1. These seismic velocities

are consistent with densities less than 3.1 g em- 3• The

base of the crust is likely to be very complicated in

detail because of the possibility of underplating by

basic or ultramafic material and because of the prob-

able interlayering of high-velocity mantle material with

lower-velocity crustal material.

A midcrustal boundary, the Conrad discontinuity, is

sometimes observed to separate upper and lower crust at

about 10-to 20-km depth. It is commonly discontinuous

and often absent or poorly constrained by seismic data;

especially in shield areas. Sometimes it is gradational

over several kilometers. The failure of superdeep drill

holes to identify many of the discontinuities that were

based on interpretations of the geophysical evidence

remains an intriguing area of investigation.

Crustal density estimates range from 2.7 to 2.9 g

cm-3 and increase with depth on average. This param-

eter is difficult to define precisely because of the mod-

el-dependent nature of the lower crust and because of

the great geological complexity of the crust generally.

A crustal mass of2.06 x 1025 g (±7%) is obtained if an

average density of2.S g cm? is adopted. On the basis

of this calculation, the continental crust forms 0.54%



244 • Taylor and McLennan: CONTINENTAL CRUST GEOCHEMISTRY

TABLE 2. Crust Formation Age Distributions

33,2/ REVIEWS OF GEOPHYSICS

Nd Model Age Distributions* This Paper

0% Crust by 4.0 Eons 8% Crust by 4.0 Eons

Distribution, Growth Rate, Distribution, Growth Rate,

Age, eons % % eon" % % eon?'

1 . 0 - 0 . 0 1 8 1 8 1 7 1 7

1 . 6 - 1 . 0 1 2 2 0 11 1 8

2 . 6 - 1 . 6 2 2 2 2 2 0 2 0

3 . 4 - 2 . 6 2 6 3 3 2 4 3 0

4 . 0 - 3 . 4 2 2 3 7 2 0 3 3

4 . 5 - 4 . 0 8 1 6

Total 1 0 0 1 0 0

Mean age, eons t 2 . 2 2 . 3

Model

Distribution,

%

Average

Growth rate,% eon"!

108

2 2

5 0

10

1 0

13

2 2

6 2

1 7

100

2 . 4

Data from McCulloch and Bennett [1993]. There is no identified crust greater than 4.0 eons. Two models are presented, one assuming0% crust present at 4.0 eons and a second assuming 8% crust present by 4.0 eons (a value considered by McCulloch and Bennett [1993] tobe consistent with Nd isotopic characteristics of Archean mantle-derived volcanics).

"In each case, crustal growth is assumed to be episodic, consistent with McCulloch and Bennett [1993] and as described in text.

of the mass of the crust-mantle system and only 0.35%

of the mass of the whole Earth.

The continental crust is of great antiquity when

compared with oceanic crust and thus contains most of

the geological record. An understanding of the mean

age of the continents is important for a number of

geochemical and geophysical models but is a difficult

parameter to estimate with confidence. The most robust

approach to date is based on the Sm-Nd isotope system

because in bulk rock samples it is least prone to resetting

during later metamorphism. The most important episode

of fractionation of Sm from Nd takes place at the time ofmantle melting to form continental crust. Hence the

depleted mantle Nd model age (TOM) of crustal igneous

rocks is thought to reflect the age of mantle extraction.

This simple view is complicated by a number of pro-

cesses that may later change SmlNd or underestimate

the importance of older crustal rocks. These processes

include intracrustal melting, metamorphic resetting, and

assimilation of older material. The overall effect is that

estimates of average crustal age based on the Nd isotope

system, while superior to other isotopic approaches, are

still likely to represent a minimum [McLennan, 1988].

On the basis of Nd model age provinces in NorthAmerica, Europe, and Australia (Table 2), an average

age of continental crust is about 2.0-2.3 eons [Nelson

and DePaolo, 1985;Patchett and Arndt, 1986;McCul-

loch, 1987;DePaolo et al., 1991;McCulloch and Ben-

nett, 1993]. These workers conclude that no more than

about 50% of the continental crust was in place by 2.6

eons, a value that we consider too low because it does

not account for the freeboard constraint that crustal

volumes have only increased to a minor degree since

2.5 eons (see below). A mean age of at least 2.4 eons

for the continental crust is obtained if about 60% of the

crust was in place by 2.6 eons (Table 2), a model thatwe favor here.

U P P E R CONTINENTAL CRUST

Sedimentary Rocks as Crustal Samples

Goldschmidt [1938] suggested that the mixing and

homogenizing effects of sedimentary processes should

produce uniform rare earth element patterns in sedi-

mentary rocks and that these should reflect the abun-

dances in the upper continental crust. This prediction

of Goldschmidt has proven remarkably accurate, es-

pecially for fine-grained clastic sedimentary rocks

such as shales [Taylor and McLennan, 1985]. The useof elemental abundances in sedimentary rocks as in-

dexes of upper crustal composition is based on the

observation that many elements have extremely low

concentrations in river water and seawater and so are

transported and deposited without effective loss by

clastic sediments, such as shales and sandstones (Fig-

ure O . The most important among these elements are

the REE, Y, Sc, and Th (other insoluble elements,

such as Zr, Hf, and Sn, have similar properties, but the

chemistry of these elements in clastic sediments is

totally dominated by very resistant heavy minerals,

such as zircon and cassiterite, that are readily segre-

gated during sedimentation). Shales and their relativesare volumetrically the most important sedimentary

rock types (constituting about 70% of the sedimentary

mass), and so the abundances of these insoluble ele-

ments in shales can be closely related to that of the

upper crust.

Although some diversity exists in REE patterns,

especially sediments derived directly from volcani-

cally active continental margins, the average REE

pattern is very similar for most Phanerozoic and mod-

ern shales [McLennan et al., 1990, 1993; McLennan

and Taylor, 1991].These include estimates of averages

from three continents: post-Archean Australian aver-age shale (PAAS), European shale composite (ES),



33, 2/ REVIEWS OF GEOPHYSICS Taylor and McLennan: CONTINENTAL CRUST GEOCHEMISTRY. 245

and North American shale composite (NASC) samples

[see Taylor and McLennan, 1985]. These REE pat-

terns are enriched in the light REE (LREE) relative to

chondritic patterns, which are taken as representative

of the bulk Earth (see Taylor [1992a] for a discussion

of this), and have a depletion in Eu relative to the

neighboring REE, Sm, and Gd. (This relative deple-

tion is usually referred to as a "negative Eu anomaly";

the converse enrichment in Eu is often referred to as a

"positive Eu anomaly. " See glossary for definitions of

LREE, HREE, LaN/YbN, and Eu/Eu*; for REE nor-

malizing factors, see Taylor and McLennan [1985]).

The REE are not readily fractionated in sedimen-

tary processes (except that Ce3+ may oxidize to Ce4+

and separate as the insoluble phosphate in marine

environments, or less frequently as the hydroxide in

terrestrial environments). Thus the REE pattern is

recording a previous signature of igneous events in-

volving partial melting (or crystal fractionation) and

removal of Eu by plagioclase as a residual phase [e.g.,

Taylor and Mcl.ennan, 1988].

Modern sediments mostly display REE patterns

similar to those of the post-Archean shales. The REE

patterns of most suspended particulate matter in some

of the world's major rivers [Goldstein and Jacobsen,

1988] are also very similar to the PAAS, although

slightly enriched in total REE as a result of their finer

grain size (REE are concentrated in the clay fraction

and in heavy minerals).

Location of REEin Sediments

The bulk of the REE elements reside in the fine-grained (clay and silt) fraction (there is no correlation

with specific clay mineralogy) and also in trace miner-

als such as zircon, monazite, and apatite [Cullers et

al., 1979]. There is some indication that significant

fractions of REE in clay minerals may also occur as

microinclusions of resistate minerals such as apatite

[Caggianelli et al., 1992].Sand fractions within a sed-

imentary rock have significantly lower total abun-

dances than finer-grained fractions. Coarse-grained

sedimentary rocks typically have REE patterns essen-

tially parallel to those of shales, but with lower total

abundances than shales, due mainly to the dilutingeffect of quartz.

The overall shape of the patterns (Eu/Eu*, LaNI

YbN)is generally similar for sandstones and shales. Eu

anomalies are similar in all size fractions. The REE

abundances in quartz-rich sedimentary rocks are typ-

ically very low. The potential of heavy minerals to

distort REE patterns in sedimentary rocks is well

recognized. For example, Archean metaquartzites

from the Western Gneiss Terrain, Australia [Taylor et

al., 1986] show enrichments of light REE due to the

presence of monazite. However, these minerals are

only rarely concentrated in amounts sufficient to cause

perceptible effects on the REE patterns. Thus the

resistant mineral zircon, typically enriched in heavy

--------------------8--

(iJ::l

0W0.. 1.000..

: : : >

O J

(ijs:Cf)

0.10

-10 ~ ~ 4 ~ 0

log (Seawater / Upper Crust)

Figure 1. Plot of the concentration ratios of (average shale)/

(upper crust) versus (seawater)/(upper crust) to illustrate the

elements most useful for estimating upper crustal abun-

dances from sedimentary data. Data are taken primarily

from Taylor and McLennan [1985]; however, since some of

their upper crust values are derived directly from shales

(e.g., REE, Th, and Sc), for the purposes of this diagram, the

independently determined values of Shaw et al. [1976] are

used for these elements. With the exception of Co (and

perhaps Ti, for which seawater data are poorly established),

the ratios of (average shale)/(upper crust) vary by a factor of

< 1.5 for elements that do not partition into water during

geological processes (Iseawaterl/tupper crust) <10-5). Com-

plications exist involving redox controls (Fe and Mn) and

heavy mineral concentrations (e.g., Zr), but in general the

abundances of these elements in shales most closely reflect

upper crustal abundances (notably for REE, Th, and Sc). In

contrast, the ratio (average shale)/(upper crust) varies by afactor of >15 for elements that are partitioned into natural

waters (Iseawaterl/tupper crust) >10-3). Such elements are

readily dissolved in natural waters and are concentrated into

chemical sediments (e.g., Ca, Na, and Sr) or are readily

mobilized during weathering, diagenesis, hydrothermal al-

teration, and so forth (e.g., Mg, U, and B).

REE, affects the bulk rock patterns only when zircon

constitutes more than about 0.06% (or Zr abundances

exceed about 300ppm) because every 100ppm ofZr as

zircon adds only about 0.25 times chondri tic levels ofYb (typical bllik rock patterns are about 10-15 times

chondritic levels for that element).

Sedimentary Processes and REEPatterns

Sedimentary processes may affect REE distribu-

tions in sedimentary rocks. Recent studies of granite

weathering indicate that substantial fractionation may

take place within weathering profiles [e.g., Nesbitt,

1979;Banfield and Eggleton, 1989]resulting in, among

other things, significant Ce anomalies associated with

formation of Ce4+ and stable cerium hydroxides. In

most situations, such fractionation does not translate

into unusual sedimentary rock REE patterns because

the weathering profile is well mixed during subsequent




erosion and transport [Nesbitt, 1979].Exceptions may

exist. For example, McDaniel et al. [1994]recognized

such a process as affecting Early Proterozoic first-

cycle turbidites derived from granitic rocks. In this

case, Ce anomalies were present, and Nd isotopes

clearly record partial resetting ofthe Sm-Nd system at

about the time of sedimentation.

Sedimentary transport may also result in changing

sedimentary REE patterns due to heavy mineral (no-

tably zircon and monazite) fractionation. Such effects

can be modeled with some confidence [McLennan,

1989]and are most likely to be a significant influence in

clastic sediments such as quartzites with low REE

abundances. Recent studies have also demonstrated

that REE may be significantly affected during certain

conditions of diagenesis [e.g., Ohr et al., 1991;

Milodowski and Zalasiewicz; 1991;Bock et al., 1994].

Redistribution ofREE among newly formed diagenetic

minerals, without significant transport, appears to oc-

cur in most cases [Ohr et al., 1991]. However, inexamples of diagenetic ally altered Ordovician tur-

bidites, there is evidence oflarger-scale transport (and

partial resetting of the Nd isotope system) associated

with formation of REE-bearing minerals such as mon-

azite [Milodowski and Zalasiewicz; 1991;Bock et al.,

1994]. In addition to these examples, reports of heavy

REE (HREE, i.e., Gd-Lu) depletion in continent-de-

rived marine sediment continue to appear but in most

cases may be attributed to analytical difficulties [e.g.,

Sholkovitz; 1988, 1990].

Although such effects appear to be comparatively

rare, they are of some interest. Unusual sedimentaryprocesses mostly result in the production of anoma-

lous REE patterns, such as Ce anomalies [McDaniel et

al., 1994] or severe LREE depletion [Bock et al.,

1994]. In each of these examples, evaluation of the

REE systematics permits constraints to be placed on

the original REE patterns [McDaniel et al., 1994;Bock

et al., 1994]. These results also engender caution in

interpreting the numerous Nd isotope studies of sedi-

mentary rocks in which the geochemical systematics

beyond Sm and Nd are not evaluated.

Loess

About 10% of the Earth's land surface is covered

with Pleistocene loess. The origin of most loess by

aeolian transport from glacial outwash, particularly

during cold dry climatic regimes is well established.

This combination of widespread production of silt-

sized rock flour and its transport by wind over hun-

dreds of kilometers thus provides geochemists with a

natural sampling of comparatively unweathered mate-

rial from the exposed crust. Loesses from North

America, Europe, China, and New Zealand show very

uniform REE abundances (although sometimes diluted

by quartz or carbonate), with Eu depletions equivalent

to those observed in clastic sedimentary rocks, indi-cating that they are providing a reasonable sample of


the upper continental crust [Taylor et al., 1983;Liu et

al., 1993]. The similarity of REE patterns for the

widely scattered loess deposits to shale averages (e.g.,

PAAS, ES, and NASC) indicates that loess also pro-

vides the same information on the composition of the

upper crust as that provided by clastic sediments.

There is little evidence that these sediments have been

affected by significant weathering. This demonstrates

that upper crustal provenance is the prime control on

REE patterns of clastic sedimentary rocks rather than

the effects of sedimentary processes.

Composition of the Post-Archean Upper Crust

The composition of the post-Archean upper crust is

well established, with several estimates converging on

a composition close to that of granodiorite. The meth-

ods employed have ranged from extensive sampling

and massive analytical programs involving estimates

of the amounts of various rock types to the use of

sedimentary rocks and loess as natural sampling pro-cesses. Some recent studies employing such ap-

proaches include those of Taylor and McLennan

[1985], Wedepohl [1991], Ronov et al. [1992], and

Condie [1992].

The most remarkable feature of the REE abundance

patterns in post-Archean sedimentary rocks is their

uniformity on a worldwide scale. REE patterns for the

post-Archean Australian average shale are similar to

those of composite shale samples from Europe and

North America. These patterns are all characterized

by light REE enrichment and relatively flatheavy REE

abundances about 10 times chondritic, with a ratheruniform depletion in Eu (Eu/Eu* = 0.65). This unifor-

mity both within and between continents is accord-

ingly interpreted to represent the REE abundances in

the upper continental crust exposed to weathering

(Figure 2). Mass balance calculations involving aver-

ages of the various sedimentary rock types (shales,

sandstones, carbonates, and evaporites) indicate that

while REE patterns in fine-grained sedimentary rocks

are parallel to upper crustal abundances, they proba-

bly overestimate the absolute abundances by about

20% [Taylor and McLennan, 1981].

The abundances of the insoluble elements Th and

Sc also reflect upper crustal abundances in the same

manner as the REE (Figure O . Thorium is typically

concentrated in granitic rocks. Scandium, although

trivalent and a member the same group (III) of the

periodic table as the REE, is a much smaller ion, and

is concentrated in basic rocks, entering early crystal-

lizing pyroxenes. The ThlSc ratio thus forms an excel-

lent index of the relative proportions of granitic and

basic rocks.

From sedimentary REE, Th, and Sc data, it is

possible to extrapolate to obtain the upper crustal

abundances of a number of other elements by using

elemental ratios that either are constant across a widerange of igneous compositions (e.g., K/U) or vary



33,2/ REVIEWS OF GEOPHYSICS Taylor and McLennan: CONTINENTAL CRUST GEOCHEMISTRY. 247

(/)

Q)

-;::

"0

§100.0

..co

Ea.a.

V24·73

198A·2·2500.0

- -Ea.a.

10.0

Figure 2. Chondrite-normalized REE diagram showing deep-

sea pelagic muds from the Pacific Ocean basin (data from

Ben Othman et al. [1989] and Lin [1992]). Shown for com-

parison are estimates of post-Archean and Archean upper

crust, from Taylor and McLennan [1985]. Note that pelagic

sediments from the Pacific basin, a region surrounded by arc

volcanism, display negative Eu anomalies characteristic of

the upper crust. This signature is largely due to aeolian

sources. The paucity of such patterns in Archean greenstone

belts is considered strong evidence for the limited extent of

stabilized Archean crust that had undergone intracrustal

differentiation resulting in negative Eu anomalies.

systematically with bulk composition (e.g., KlRb).

Thus the abundance of V can be obtained from gen-

erally accepted upper crustal Th/V ratios (3.8) and K

from KIV ratios (10,000). Rb can be obtained from

KlRb (250); Sr from Rb/Sr (0.3) and so on (Table 3).

Table 4 gives the normative mineralogy of the upper

crust, based on its major element composition.

Origin of the Post-Archean Upper Crust

Geochemical balance calculations indicate that theupper crustal composition cannot be representative of

the entire crust. This observation is also consistent

with heat flow and heat production data, although

these do not uniquely separate the relative crustal and

mantle contributions. The dominance in the upper

crust of compositions effectively equivalent to grano-

diorite (including high abundances of heat-producing

elements and other large-ion lithophile elements) and

the ubiquitous occurrence of the PAAS REE patterns

(including negative Eu anomalies) provide major clues

to the derivation of the upper crust.

The Eu depletion in post-Archean sedimentaryrocks provides strong evidence for an intracrustal or-

igin for much of the upper crust. Igneous rocks of

mantle derivation rarely display Eu anomalies, and no

primitive volcanic rock derived directly from the man-

tle exhibits a relative depletion in Eu. The origin of the

marked depletion of Eu typical of chondrite-normal-

ized REE patterns in terrigenous sedimentary rocks is

not due to surficial processes of oxidation or reduc-

tion. Eu is, with rare exceptions, present as thetrivalent ion in sediments [e.g., Sverjensky, 1984], so

that the depletion in Eu is the signature of an earlier

igneous event. Vnder the reducing conditions typical

of magmas, Eu is divalent. The relatively flat and

uniform HREE patterns seen in sedimentary rocks

(GdNNbN = 1.0 to 2.0) suggests that HREE-enriched

minerals, notably garnet and amphibole, are not im-

portant fractionating phases during the formation of

the upper continental crust.

The generally accepted explanation for the origin of

the Eu-depleted, K-rich granites and granodiorites that

now dominate the upper crust is that they form

through intracrustal melting. Thus the most likely ex-

planation for the upper crustal Eu anomaly is that Eu

is left in residual plagioclase in the lower crust. Pla-

gioclase is stable only to a pressure of 10kbar (a depth

of 40 km on the Earth), so that this sink for Eu is

consistent with the scenario of intracrustal melting for

granite origin that is supported by experimental stud-

ies [e.g., Wyllie, 1983]. There are two sources of heat

available to induce intracrustal melting. The first is the

heat generated by radioactive decay of K, V, and Th.

The second source, underplating of the crust by basal-

tic magmas, is less easily evaluated but appears to be

required because the crustal radioactive sources aregenerally considered to be inadequate for the task. As

we discuss in greater detail below, recognition of Eu-

rich residual material in the available lower crustal

samples has proven elusive and represents one of the

major outstanding issues in crustal geochemistry [e.g.,

Rudnick, 1992a].

Sedimentary rocks provide information about how

long this process has been operating. As was discussed

earlier, the REE patterns of loess from widely scat-

tered localities across the globe are uniform, indicating

that the processes producing the upper crust have not

changed over the period represented by the sourceregions of the loesses. However, loess is derived from

source rocks of widely varying age, extending back

into the Proterozoic. The Nd model age of sedimen-

tary rocks is interpreted as the average age from when

the various provenance components were extracted

from the mantle [McCulloch and Wasserburg, 1978].

The samples studied by Taylor et al. [1983] came from

China, Europe, New Zealand, and North America,

with Nd-depleted mantle model ages TDM ranging from

1060 to 1700 Ma. In fact, such sedimentary REE pat-

terns likely exist for sedimentary rocks with Nd model

ages older than about 2 eons [e.g., McLennan andHemming, 1992]. This variation in average prove-

nance age, coupled with their uniform chemistry, in-



TABLE 3. Chemical Composition of the Continental Crust

Upper Bulk Lower Average Archean ArcheanContinental Continental Continental Andesitic Upper Bulk

Element Crust Crust Crust Crust Crust Crust

Li, ppm 2 0 1 3 1 1 1 0

Be, ppm 3 . 0 1 . 5 1 . 0 1 . 5

B, ppm 1 5 1 0 8 . 3

Na, wt. % 2 . 8 9 2 . 3 0 2 . 0 8 2 . 6 0 2 . 4 5 2 . 2 3

Mg, wt. % 1 . 3 3 3 . 2 0 3 . 8 0 2 . 1 1 2 . 8 3 3 . 5 6

AI, wt. % 8 . 0 4 8 . 4 1 8 . 5 2 9 . 5 0 8 . 1 0 8 . 0 4

Si, wt. % 3 0 . 8 2 6 . 8 2 5 . 4 2 7 . 1 2 8 . 1 2 6 . 6

P, ppm 7 0 0

K, wt. % 2 . 8 0 0 . 9 1 0 . 2 8 1 . 2 5 1 . 5 0 0 . 7 5

Ca, wt. % 3 . 0 0 5 . 2 9 6 . 0 7 5 . 3 6 4 . 4 3 5 . 2 2

Sc, ppm 1 1 3 0 3 6 3 0 1 4 3 0

Ti, wt. % 0 . 3 0 0 . 5 4 0 . 6 0 0 . 4 8 0 . 5 0 0 . 6 0

V, ppm 6 0 2 3 0 2 8 5 1 7 5 1 9 5 2 4 5

Cr, ppm 3 5 1 8 5 2 3 5 5 5 1 8 0 2 3 0

Mn, ppm 6 0 0 1 4 0 0 1 7 0 0 1 1 0 0 1 4 0 0 1 5 0 0

Fe, wt. % 3 . 5 0 7 . 0 7 8 . 2 4 5 . 8 3 6 . 2 2 7 . 4 6

Co, ppm 1 0 2 9 3 5 2 5 2 5 3 0

Ni, ppm 2 0 1 0 5 1 3 5 3 0 1 0 5 1 3 0

Cu, ppm 2 5 7 5 9 0 6 0 8 0

Zn, ppm 71 8 0 8 3

Ga, ppm 1 7 1 8 1 8 1 8

Ge, ppm 1 . 6 1 . 6 1 . 6

As, ppm 1 . 5 1 . 0 0 . 8

Se, ppm 5 0 5 0 5 0

Rb, ppm 1 1 2 3 2 5 . 3 4 2 5 0 2 8

Sr, ppm 3 5 0 2 6 0 2 3 0 4 0 0 2 4 0 2 1 5

Y,ppm 2 2 2 0 1 9 22 1 8 1 9

Zr, ppm 1 9 0 1 0 0 7 0 1 0 0 1 2 5 1 0 0

Nb, ppm 2 5 1 1 6 1 1

Mo, ppm 1 . 5 1 . 0 0 . 8

Pd, ppb 0 . 5 1 1

Ag, ppb 5 0 8 0 9 0

Cd, ppb 9 8 9 8 9 8

In, ppb5 0 5 0 5 0

Sn, ppm 5 . 5 2 . 5 1 . 5

Sb, ppm 0 . 2 0 . 2 0 . 2

Cs, ppm 3 . 7 1 . 0 0 . 1 1.7Ba, ppm 5 5 0 2 5 0 1 5 0 3 5 0 2 6 5 2 2 0

La, ppm 3 0 1 6 1 1 1 9 2 0 1 5

Ce, ppm 64 3 3 2 3 3 8 4 2 3 1

Pr, ppm 7 . 1 3 . 9 2 . 8 4 . 3 4 . 9 3 . 7

Nd, ppm 2 6 1 6 1 2 . 7 1 6 2 0 1 6

Sm, ppm 4 . 5 3 . 5 3 . 1 7 3 . 7 4 . 0 3 . 4

Eu, ppm 0 . 8 8 1.1 1 . 1 7 1 . 1 1 . 2 1.1Gd, ppm 3 . 8 3 . 3 3 . 1 3 3 . 6 3 . 4 3 . 2

Tb, ppm 0 . 6 4 0 . 6 0 0 . 5 9 0 . 6 4 0 . 5 7 0 . 5 9

Dy, ppm 3 . 5 3 . 7 3 . 6 3 . 7 3 . 4 3 . 6

Ho, ppm 0 . 8 0 0 . 7 8 0 . 7 7 0 . 8 2 0 . 7 4 0 . 7 7

Er, ppm 2 . 3 2 . 2 2 . 2 2 . 3 2 . 1 2 . 2Tm, ppm 0 . 3 3 0 . 3 2 0 . 3 2 0 . 4 3 2 0 . 3 0 0 . 3 2

Yb, ppm 2 . 2 2 . 2 2 . 2 2 . 2 2 . 0 2 . 2

Lu, ppm 0 . 3 2 0 . 3 0 0 . 2 9 0 . 3 0 0 . 3 1 0 . 3 3

Hf, ppm 5 . 8 3 . 0 2 . 1 3 . 0 3 3

Ta, ppm 2 . 2 1 . 0 0 . 6

W,ppm 2 . 0 1 . 0 0 . 7

Re,ppb 0 . 4 0 . 4 0 . 4

Os, ppb 0 . 0 5 0 . 0 5 0 . 0 5

Ir, ppb 0 . 0 2 0 . 1 0 . 1 3

Au, ppb 1 . 8 3 . 0 3 . 4

TI, ppb 7 5 0 3 6 0 2 3 0

Pb, ppm 2 0 8 . 0 4 . 0 1 0

Bi, ppb 1 2 7 6 0 3 8

Th, ppm 1 0 . 7 3 . 5 1 . 0 6 2 . 5 5 . 7 2 . 9

V,ppm 2 . 8 0 . 9 1 0 . 2 8 1 . 0 1 . 5 0 . 7 5

Adapted from Taylor and McLennan [1985], with minor alterations [see Esser and Turekian, 1993; Jochum et al., 1993].




TABLE 4. Normative Mineralogy of Continental Crust

Bulk Lower

Upper Crust, Crust,

Crust, % % %

Quartz(Q) 16.1 6.6 3.7Orthoclase (Or) 20.1 6.5 2.0

Albite (Ab) 33.0 26.2 23.7Anorthite (An) 13.9 26.2 30.4Diopside (Di) 5.8 8.7 9.8Hypersthene (Hy) 10.0 24.1 28.6Ilmenite (II) 1.0 1.7 1.9Total 99.9 100.0 100.1

dicates that the upper continental crust was both uni-

form in composition and produced by similar

processes over periods extending back at least 2 billion

years. Accordingly, the processes producing the upper

crustal composition which is being sampled during the

formation of sedimentary rocks have remained the

same well back into the Proterozoic.

Archean-Proterozoic Transition

Compared with post-Archean PAAS-type REE pat-

terns, abrupt changes are observed in clastic sedimen-

tary rocks over time intervals from 3.2 to 2.5 eons in

South Africa [McLennan et al., 1983] and inWestern

Australia [Taylor and McLennan, 1985]. In the early

Proterozoic (2.5-2.2 eons) Huronian sedimentary suc-

cession of Canada, McLennan et al, [1979] found a

progressive change from REE patterns lacking Eu

depletions at the base, to typical post-Archean pat-terns at the top. Similar changes in sedimentary REE

patterns have been noted in Early Proterozoic succes-

sions of the Pine Creek Geosyncline and Hamersley

Basin, Australia, and the Slave Province of Canada

(see review by Taylor and McLennan [1985]). This

change was interpreted to reflect an episodic change in

upper crustal composition from a more primitive state

and related to a massive emplacement of K-rich gra-

nitic rocks, depleted in Eu, toward the close of the

Archean. This is generally referred to as "cratoniza-

tion," the process that produces massive intracrustal

melting to produce granites, transfers heat-producingelements to the upper crust, and generally "stabilizes"

the crust [e.g., Pollack, 1986]. (The development of a

thick lithospheric keel during crustal growth also is

likely to contribute to the inherent stability of cratons.)

It was nonsynchronous over the globe and extended

over several hundred million years. These events mark

the development of extensive fully stabilized continen-

tal crust.

The Archean-Proterozoic transition thus marks a

major change both in the volume of crust, and in

intracrustal differentiation. The crustal processes re-

sponsible for these changes took place during the late

Archean and are recorded in Early Proterozoic sedi-

mentary rocks. The presence of large masses of un-

subductable continental crust changed the tectonic

regime from the multiplate Archean crust and pro-

duced the present linear (e.g., South America) or arc-

uate (e.g., western Pacific arcs) subduction zones.

Many geological events correlate with a major

change at the Archean-Proterozoic boundary. The

widespread occurrence of uranium deposits in basal

Proterozoic sediments can be attributed to the enrich-

ment of the upper crust in incompatible elements due

to intracrustal melting. The dramatic increase in 87Sr

in marine carbonates from that period [e.g., Veizer,

1983] is similarly due to an upper crustal enrichment in

87Rbin the K-rich granites that came to dominate the

upper crust (and a reduction of mantle flux of Sr

resulting from a cooling Earth). It is also interesting to

note the great proliferation of stromatolites after 2.5

eons, which reflects the widespread occurrence of

stable continental marine shelves that formed as a

consequence of events around the Archean-Protero-

zoic boundary (the stromatolites declined early in the

Phanerozoic owing to the emergence of metazoan

grazing species [e.g., Schidlowski, 1992]). The prolif-

eration of banded iron formations, a chemical sedi-

ment that fills the sedimentary environment niche of

carbonate rocks in later geological periods, also can be

related in part to the development of stable shelves

during the late Archean-Early Proterozoic [e.g.,

McLennan, 1980]. The first supercontinent likely

formed at this time [e.g., Nance et al., 1986; Hoffman,

1992; Unrug, 1992], and major climatic changes also

appear to correlate with the Archean-Proterozoic tran-

sition [e.g., Young, 1991].

POST-ARCHEAN LOWER CRUST

Seismic reflection profiling has revealed the very

complex nature of the lower crust, which appears to be

at least as diverse as the upper crust. It is much more

difficult to arrive at a representative composition for

the lower crust than for the upper crust, owing to its

inaccessibility and to the absence of some sampling

technique, such as is provided for the upper crust by

sediments, that might provide a meaningful average.Two types of samples, xenoliths arid granulite facies

rocks, are available, both providing enigmatic infor-

mation [Rudnick and Presper, 1990].

Granulite facies regions are often interpreted as

samples of the lower crust. Archean granulite terranes

commonly possess positive Eu anomalies; however,

these are typically found in the more evolved compo-

sitions rather than in mafic material that could repre-

sent residues after partial melting. Positive Eu anom-

alies are less frequent in post-Archean granulite

terranes [Rudnick and Presper, 1990] and in any case,

such terranes appear, on compositional grounds, to be

mostly upper crust, depressed in Himalayan-type col-

lisions or formed in association with rifts, hot spots,



250 • Taylor and McLennan: CONTINENTAL CRUST GEOCHEMISTRY 33, 2/ REVIEWS OF GEOPHYSICS

and volcanic arcs, and thus not a good model on which

to base lower crustal compositions [e.g., Bohlen, 1987;

Ellis, 1987;Bohlen and Mezger, 1989;Mezger, 1992].

Indeed, Mezger [1992] suggested that most regional

granulites represent the transitional region between

upper and lower crust.

Xenoliths, recording pressure-temperature (P- T)conditions indicative of derivation from the lower

crust, are frequently found in volcanic pipes. Unlike

the granulite facies regions, they are commonly much

more basic in composition and frequently show a rel-

ative enrichment in Eu. They appear to be related to

episodes of underplating by basaltic magmas, and in

most cases the positive Eu anomaly is best related to

the accumulation of cumulate phases rather than to

simple residues from partial melts [e.g., Rudnick and

Taylor, 1987;Rudnick, 1992b].

Current understanding of the petrogenesis of the

majority of granitic rocks and the ubiquitous presence

of negative Eu anomalies in sedimentary rocks indi-

cate that intracrustal partial melting must be a funda-

mental process governing the composition and chem-

ical structure of the continental crust. Although

complementary positive Eu anomalies are commonly

present in the available lower crustal samples (granu-

lite terranes and xenoliths), they typically cannot be

readily characterized as residual material left behind

after intracrustal partial melting.

Accordingly, a fundamental paradox exists in

crustal geochemistry. Upper crustal rocks (granites,

sediments) appear to bear clear evidence that intrac-

rustal partial melting is the dominant process causingcrustal differentiation, but the predicted residues of

such processes are surprisingly uncommon in lower

crustal samples. Paradoxes, of course, arise not from

the operations of nature, but from our imperfect un-

derstanding of them. Rudnick [1992a]has attempted to

resolve this paradox by suggesting a multistage model

whereby basaltic underplating supplies heat to melt

the lower crust (providing melts with negative Eu

anomalies) and creates an environment for the accu-

mulation of crystal cumulates (with positive Eu anom-

alies). Other possible explanations include the follow-

ing: (1) crustal melting takes place at lowertemperatures than commonly thought (i.e., under wet

conditions), and restites are present in lower-grade

(amphibolite) terranes, or (2) neither granulite terranes

nor xenoliths are entirely representative of lower

crustal material, and areas representing residues after

partial melting are not typically sampled at surface by

natural processes. In any case, the solution of this

paradox is a critical step in furthering our understand-

ing of crustal geochemistry.

In summary, the lower crust appears to be essen-

tially the residue left after extraction of the granodior-

itic upper crust together with additions from under-

plating by basaltic magmas. It is thus likely to be very

complex in detail, an example of which appears to be

the Ivrea Zone in northern Italy [Voshage et al., 1990].

On at least a local scale, delamination and sinking of

heavy residual material may also be expected, adding

to the complexity and to the variable nature of the

Moho [Kempton and Harmon, 1992].Estimates of the

average composition and normative mineralogy are

presented in Tables 3 and 4.

THE ARCHEAN CRUST

Archean Upper Crust

The sedimentary record again provides us with a

sampling of the Archean crust and reveals that the

composition of the Archean upper crust stands in

marked contrast to that of the post-Archean crust

(Table 3). The most notable difference is shown by the

REE patterns in the Archean sedimentary rocks,

which in general show less enrichment in the LREE

and no Eu anomalies (Figure 3), relative to those of the

post-Archean crust. These differences inREE patterns

between Archean and post-Archean clastic sediments

have become a crucial observation for models of the

evolution of the continental crust and have been doc-

umented in many studies (see summary of Taylor and

McLennan [1985]).

Most studies have dealt with the relatively well

preserved low-grade sedimentary rocks of the ubiqui-

tous Archean greenstone belts in which turbidites are

the most abundant type of sedimentary rock. There is

often a general superficial resemblance to the REE

patterns of island arc volcanic rocks such as andesites.However, this similarity is merely the result of deri-

vation largely from a mixture of the ubiquitous bi-

modal suite of basaltic and felsic igneous rocks (to-

nalites, trondhjemites, granodiorites, or the "TTG

suite," and their volcanic equivalents) which charac-

terize many Archean terranes.

There is, in detail, an extreme variability of REE

patterns in Archean sediments in contrast to the gen-

erally uniform post-Archean REE patterns. Although

many Archean sediments have intermediate patterns

reminiscent of andesites, both very steep and flat pat-

terns are locally abundant [McLennan and Taylor,1984]. The steep patterns occur in first-cycle sedi-

ments derived directly from tonalites, trondhjemites,

and granodiorites (TTG igneous suites), whereas the

flat patterns are derived from basaltic precursors to

first-cycle sediments. Both REE patterns come from

the Archean "bimodal igneous suite," which largely

dominates the Archean upper crust.

This distinction between the REE patterns of Ar-

chean and post-Archean sediments has led to some

controversy. Thus some workers [e.g., Gibbs et al.,

1986]have argued that this distinction is without age

significance and is the result of differing tectonic set-

tings, implying that post-Archean greywackes should

be identical to those of Archean age. However, these



33, 2 I REVIEWS OF GEOPHYSICS Taylor and McLennan: CONTINENTAL CRUST GEOCHEMISTRY. 251

1.2

-I e

::l1.0

W

-0.80.6

1.2

Post-Archean:'"

l1i: e

i~

: : - - - - - 8f - - - - - - < w £ v : - ' Q ' , , - - - - - - - -

. , " " ://////"~ I

> ~ / » » ~ / : ; .: /~ :'>7-~;~;~~<'/""'" : Cratonic

! Sediments

0.6

0.4 '-'-' .J...O- ~ 0.4 ..........

0.5 1.0 1.5 2.0 2.5 3.0 3.5 4.0 0.5

GdN / YbN

1.0 2.0 2.5

GdN

/ YbN

3.0.5 3.5 4.0

Figure 3. Plot of Eu/Eu* versus GdNlYbN for sedimentary rocks. Three fields are shown: Archean and

post-Archean sedimentary rocks deposited in volcanically active tectonic settings and "cratonic" sedi-

mentary rocks of all ages. Adapted fromMcLennan and Taylor [1991].Each field incorporates at least 90%

of the available data. Post-Archean active margin sedimentary rocks almost completely overlap the fieldof cratonic sediments and thus are significantly more depleted in Eu (lower Eu/Eu*) than are Archean

sedimentary rocks of comparable tectonic association. Archean sedimentary rocks commonly possess

heavy REE depletion (high Gd/Yb), whereas post-Archean sediments very rarely display such features.

Cratonic sedimentary rocks of all ages appear generally similar, with significant Eu depletion, but the

extent of such terranes appears to have been much less in the Archean. Overall, these and other

geochemical data indicate that the Archean and post-Archean upper continental crust differed in compo-

sition, that the nature ofmantle sources ofcontinental crust (or the processes of crust formation) may have

differed, and that the mechanisms of crustal differentiation (notably the role of intracrustal melting) also

differed.

younger sediments display the crucial signatures of Eu

depletion in all post-Archean tectonic environments(e.g., Figure 3), except those of forearc basins of

oceanic island arcs. Many Archean greywackes ap-

pear to have formed in tectonic settings such as back

arc, continental arc, trailing edge, and foreland basins

and are clearly petrographically distinct from their

modern counterparts in being plagioclase-rich (to-

nalitic sources) but having relatively few andesitic

rock fragments [McLennan, 1984].Finally, the general

evidence that Archean turbidites are derived from the

"bimodal basaltic-TI'G suite" rather than from "is-

land arc andesites" renders the supposed analogy be-

tween conditions in the Archean and modern arc en-

vironments superficial.

Others [e.g., Condie, 1992] appear to restrict the

usage of "upper crust" to cratonic settings only and

argue that by this definition the Archean and post-

Archean upper crustal compositions are indistinguish-

able. By this interpretation, Archean greenstone belt

sediments are not considered to be part of continental

crust, even though they are mostly developed on gra-

nitic terranes and are not Archean oceanic crust

[Bickle et al., 1994].We find this definition unsatisfac-

tory because for modern environments, it would ex-

clude volcanically active settings, such as Japan, the

southwest Pacific islands, the Andes, Indonesia, andmuch of New Zealand from being considered part of

the continental crust (also note, for example, that

sediments derived from many of these terranes indeedpossess negative Eu anomalies [e.g., McLennan et al.,

1990]). There is no dispute that the upper crust ex-

posed in cratons of any age is likely to be of compa-

rable composition and possess the distinctive negative

Eu anomaly [Taylor et al., 1986] (see below). In any

case, making such arcane distinctions reduces the dis-

cussion to a question of semantics rather than address-

ing the problems of how the Earth actually evolves.

Apart from the greenstone terrane sediments, small

amounts of Archean crust are preserved in high-grade

metamorphic terranes. Studies of such terranes from

Greenland [Boak et al., 1982;McLennan et al., 1984;

Jacobsen and Dymek, 1988], India [Naqvi et al.,

1983],Montana-Wyoming [Mueller et al., 1982;Gibbs

et al., 1986],Canadian Shield [Sawyer, 1986;Taylor et

al., 1986], Western Gneiss Terrain, Australia, and

Limpopo Belt, South Africa [Taylor et al., 1986;Bo-

ryta and Condie, 1990] have indicated that these re-

gions commonly represent a tectonic environment dis-

tinct from that of the greenstone belts. Typical REE

patterns fall into two groups. One group (e.g., Ka-

puskasing [Taylor et al., 1986]and Quetico Belt, Can-

ada [Sawyer, 1986])are highly metamorphosed equiv-

alents ofArchean greenstone belt sediments. A second

group (Wyoming-Montana [Gibbs et al., 1986] andWestern Gneiss Terrain [Taylor et al., 1986])contains




many samples that are indistinguishable from typical

post-Archean patterns and represent deposition on a

stable shelf environment, most probably occurring on

small, stable minicratons.

Both groups of REE patterns are found in close

proximity in Archean high-grade terranes in India,

Greenland, and the Limpopo Belt. This diversity ofREE patterns demonstrates that the sediments were

derived from highly localized sources, as is the case

for many Archean sediments. In such regions, small

areas of stable crust (minicratons) must exist in close

association with greenstone belts, but the cratons are

not contributing significant amounts of erosional de-

bris to the greenstone belt sedimentary basins. Be-

cause of the tendency of the greenstone belt environ-

ments to undergo destruction by erosion and recycling

[Veizer and Jansen, 1985], the high-grade belts are

preferentially preserved.

The limited extent of these Archean terranes with

"post-Archean" REE signatures is indicated by the

absence in the greenstone belt sediments of such sig-

natures, despite the close association of greenstone

belts with "granitic" terranes [Bickle et al., 1994].

This is striking evidence for the limited extent of "cra-

tonic" terranes in the Archean because wind-borne

dust should have been widespread in that epoch owing

to the absence of land-based vegetation. Thus any

granitic terrane with a PAAS REE signature should

have contributed the characteristic signature of Eu

depletion to the greenstone belt sediments. The ab-

sence of a detectable signature of Eu depletion means

that such high-grade terranes represented less than10% of the exposed Archean crust [Taylor et al.,

1986].Most of the crust as sampled by the greenstone

belt terranes was derived from areas where the bi-

modal suite of basalt and TTG dominated the land area

being eroded to supply the sediments.

Another observation of importance when discuss-

ing the cause of the difference between Archean and

post-Archean sedimentary REE patterns is that in

modern environments where volcanic provenances

should dominate, such as arcs or Pacific deep-sea

environments, there is ubiquitous evidence for contri-

butions of upper crustal material with negative Eu

anomalies [e.g., Ben Othmann et al., 1989;McLennan

et al., 1990;Lin, 1992](see Figure 2). The isolation of

small Archean cratons enabled the survival of distinct

suites of sedimentary rocks with much variation in

REE patterns, in contrast to the situation in post-

Archean time.

A probable scenario for the Archean crust com-

prises many small, fast spreading plates with numer-

ous but relatively short volcanic arcs. Volcanism as-

sociated with plumes is expected to be common,

producing numerous Icelandic-type islands. Regions

of the crust that had undergone intracrustal melting

(generating upper crustal negative Eu anomalies) and

stabilization to form cratonic regions were restricted in


number and of limited extent, perhaps only slightly

greater in size than the present extent of early Archean

terranes such as inWest Greenland-Labrador and the

Minnesota River Valley.

Origin of the Archean Crust

There appears to have been only minor intracrustalmelting in the Archean. This was localized as shown

by the limited extent of PAAS-type REE patterns.

Thus there is not that distinction in the Archean crust

that is observed in the post-Archean crust, which is

highly stratified as a result of extensive intracrustal

partial melting. The steep REE patterns of the Ar-

chean TTG suites are indicative of either (1) garnet in

the residue during partial melting [e.g., Martin, 1986,

1993]or (2) partial melting of mantle sources already

possessing the signature of LREE enrichment [e.g.,

Stern et al., 1989;Evans and Hanson, 1995].

In the latter case, the ultimate origin of the HREE

depletion is likely also due to fractionation of a mineral

such as garnet, but in an earlier event. Because resid-

ual garnet is stable in mafic-ultramafic systems only at

depths below about 40 km, an origin by melting at

mantle depths for the TTG suite is indicated. The

Archean crust thus appeared to form as a mixture of

piled-up basalt-komatiite and tonalite-trondhjemite in-

trusions and extrusives. In the upper crust, sedimen-

tary data suggest that the ratio of basalt to TTG was

about equal. A ratio of 2 basalt/1 TTG suite for the

Archean bulk crust is suggested from a variety of

geochemical and geophysical constraints [Taylor and

McLennan, 1985].In our models the Archean crust be-comes more basic with depth, which we attribute to a

higher ratio of basalt to TTG suites with depth. This

would be expected from simple density considerations.

Production of the tonalite-trondhjemite suite may

be explained by rapid subduction of warm basaltic

crust [Martin, 1986; Defant and Drummond, 1990,

1993; Drummond and Defant, 1990; Abbott et al.,

1994]. This process still operates today in regions

(e.g., southern Andes) where rapid subduction of

young hot oceanic crust occurs; the slab reaches melt-

ing temperatures before complete dehydration occurs,

so that hornblende and garnet are present during melt-

ing of the subducting slab. Similarly, the generation of

the high-Mg andesite (boninite) suite is also considered

to be possibly related to subduction of young, high-

temperature lithosphere, resulting in mantle melting at

low pressure [Crawford et al., 1989].

B U L K CRUSTAL COMPOSITION

It is important to realize that the sedimentary rock

data provide information only on that portion of the

crust exposed to weathering and erosion. Simple mass

balance calculations show that a 36-km-thick crust

with K, Th, and U abundances equal to that of the




present upper continental crust would require about

80-90% of the entire Earth's complement of these

elements to be present in the continental crust. Thus

the upper crust is not representative of the entire

36-km thickness of the continental crust. Evidence

from heat flow data shows that the upper crust (about

10km thick) is strongly enriched in the heat-producing

elements (K, U, and Th).

There are two primary sources of heat: radioactive

decay, principally from 4oK, 238U, 235U, and 232Th,

contributing up to 80%of the terrestrial heat flow, and

other sources, such as secular cooling of the Earth and

core crystallization, responsible for at least 20%.

Some estimates place the radiogenic component of

terrestrial heat flow as low as 50%, and the relative

importance of the various sources of the Earth's heat

remains an important outstanding question. The

present average heat flow from continental areas is

about 50 mW m-2 compared with 100mWm-2 for the

oceans and 84mWm-2 for the whole Earth. However,it is difficult to assign precise values to the contribu-

tions from crust and underlying mantle. Temperatures

at the base of the crust are estimated as 300°-400°C in

shield areas, 500°-600°C in young orogenic areas, and

?650o-750°C in rift regions [e.g., McLennan, 1992].

This heat is derived from three main sources. One is

the so-called "reduced" heat flow, which is heat loss

from the deep interior transmitted mainly by conduc-

tion across the crust-mantle boundary. The second is

heat loss associated with tectonic activity within the

lithosphere. This component decays rapidly over pe-

riods of about 400 m.y. and is unimportant except inregions that experienced Phanerozoic tectonic activ-

ity. The third component is heat generated by radio-

active decay within the lithosphere; this is dominated

by upper crustal heat production. The estimation of

this component and the relative amounts of heat pro-

duced in the crust and mantle are currently under

debate [e.g., Nyblade and Pollack, 1993], and the

interpretation of the heat flow data is not without

problems. Until these problems are resolved, the heat

flow data place only loose constraints on bulk crustal

composition. Previous estimates based on a crustal

heat flow component of 23 mW m-2 were used by

Taylor and McLennan [1985] to arrive at the bulk

crustal composition given in Table 3. These are gen-

erally consistent with the geochemical balance con-

straints that limit the amount of incompatible elements

that can be concentrated in the crust. Until some new

constraints appear from geophysical or geochemical

evidence, and until the nature of the subcontinental

lithosphere is resolved, these values appear to us to

provide the best current estimates.

There is a well-established difference between the

heat flow in Archean and later Precambrian terranes

[e.g., Morgan, 1984;Pollack, 1980,1986;Nyblade and

Pollack, 1993],even when the effects of a tectonic heatcontribution are removed for younger crust. There

appears to be a steep offset in the data at the Archean-

Proterozoic boundary. This is not due to deeper ero-

sion of the Archean crust, removing a surficial hot

layer, because erosional levels are not significantly

deeper in Archean terranes [Watson, 1976].This dif-

ference in heat flow is probably not due to a difference

in composition between the Archean and later crusts

but is attributed by Nyblade and Pollack [1993] to

result mostly from an increase in the thickness of the

subcrustal lithosphere under Archean cratons. This

acts to deflect mantle heat flow around the Archean

cratons, resulting in a lower surface heat flow. Origi-

nally, these differences were attributed to differing

crustal abundances between Archean and post-Archean

stabilized crust [Morgan, 1984; Taylor and McLennan,

1985].Itappears that the principal controllingfactor may

be not the crustal abundances of the heat-producing

elements but the thickness of the subcrustal lithosphere

[e.g., Nyblade and Pollack, 1993].

Thus there may be little difference in bulk compo-sition between the Archean and present bulk crust

(that incorporates a 75%Archean component) despite

a significant difference in the upper crustal composi-

tions between Archean and post-Archean time. The

bulk crustal compositions listed by Taylor and McLen-

nan [1985] (see also Table 3) in fact are quite similar

(present values of 0.91% K, 3.5 ppm Th, and 0.91 ppm

U compared with bulk Archean crustal values of

0.75% K, 2.9 ppm Th, and 0.75 ppm U). These differ-

ences are probably within the error limits.

Nevertheless, there are some significant distinc-

tions between Archean and post-Archean crustal com-positions that have important implications for crust-

mantle evolution. For example, HREE depletion

(GdN/YbN > 2.0) is a common feature in Archean

crustal rocks but is virtually absent from rocks formed

during the post-Archean. The implications for crustal

evolution are discussed below. In our models, most of

the crust is generated in the Archean, with lesser

additions from later island arc volcanism, to make up

the present crust.

THE ORIGIN OF THE CRUST

The geological record indicates that there is a basic

distinction between the igneous activity that contrib-

uted to the formation of the continental crust in the

Archean and post-Archean epochs (Figure 4). The

basic cause of this difference is higher heat flowduring

the Archean. This results in the subduction of young

hot oceanic lithosphere during the Archean. There are

a number of possible fates for this subducted material.

In most models, such basaltic crust reaches melting

temperatures before dehydration has occurred. Partial

melting occurs under these conditions, leaving a horn-

blende-garnet residue [e.g., Arth and Hanson, 1975;Martin, 1986]resulting in production of the TTG suite



254. Taylor and McLennan: CONTINENTAL CRUST GEOCHEMISTRY 33, 2/ REVIEWS OF GEOPHYSICS

Post-Archean Growth of New Continental Crust

Andesite volcano Sedimentary

Sea level

Trench

t

FromMid-ocean ridge-

Plume headof basaltunderplatedb e ne a th c ru s t?

Mantle

Growth of New Continental Crust during the Archean

--

Felsic volcano

First cycle sediments

Sea level

25

Oceanic crust

Tonalite (sodium granite)

intrusion

Mantle km

Zone of mantle melting?

50

'---J--'-_- Zone of dehydration and/or partialMelt ing of downgoing slab

Figure 4. Schematic models illustrating the difference between Archean and post-Archean crustal devel-

opment at subduction zones. During the Archean, subducting oceanic crust was younger and hotter on

average, and the pressure-temperature conditions of dehydration/melting of the subducting slab and

overlying mantle differed considerably.

with steep REE patterns and either a zero or an occa-

sional positive Eu anomaly.

Recently, modeling of REE and other elements inundeformed posttectonic TTG suite rocks has led to

suggestions that they may be derived directly as man-

tle melts that have been enriched in a large-ion litho-

phile element (LILE) component shortly before«100-200 m.y.) formation [e.g., Shirey and Hanson,




1984; Stern and Hanson, 1991; Evans and Hanson,

1995]. In this model, two observations are particularly

significant.

First, in cogenetic suites, HREE are commonly

subparallel (i.e., uniform Gd/Yb ratios) over a wide

range of Si02 content. Such an observation is incon-

sistent with garnet's being a significant residual phasebecause even minor differences in the amount of re-

sidual garnet would lead to highly variable levels of

HREE depletion (i.e., variable GdlYb ratios).

Second, the Mg/(Mg + Fe) ratios in some suites of

TTG rocks are uniformly higher than those generated

by melting of basaltic compositions under a wide va-

riety of experimental conditions and compositions

[Evans and Hanson, 1995]. Thus melting of basalt

does not appear to be capable of generating the appro-

priate compositions for at least some of the TTG

suites.

Evans and Hanson [1995] have noted some similar-

ities in the compositions of high-Mg (most primitive)

TTG rocks (part of the so-called Sanukitoid Suite

[Shirey and Hanson, 1984]) in the southwestern Supe-

rior Province and modern boninites (high-Mg andes-

ites) in island arc environments. Generation of such

material appears to require low-pressure and high-

temperature conditions. In the Archean, such condi-

tions could be achieved by mantle melting associated

with dehydration and/or partial melting of subducted

young and hot oceanic lithosphere [Martin, 1986; De-

fant and Drummond, 1990, 1993; Drummond and De-

fant, 1990].

The Archean crust is interpreted as being derivedfrom mixtures of the two dominant ("bimodal") igne-

ous lithologies in the Archean: basalts and Na-rich

igneous rocks such as tonalites, trondhjemites, and

granodiorites, or the TTG suite (and their volcanic

equivalents). Rarely, REE patterns with Eu depletion

similar to PAAS are observed in the sedimentary

record, but they appear to be restricted to cratonic

sediments preserved in some high-grade metamorphic

terranes. These are interpreted as being derived from

minicratons [Taylor et al., 1986], forerunners of the

massive cratonic development in the late Archean.

A massive increase in the growth rate of the conti-

nental crust occurred OVer an extended period be-

tween 3.2 and 2.6 eons (with the exact age differing for

individual cratonic regions). This is well documented

by the Nd isotopic evidence [e.g., McCulloch and

Wasserburg, 1978; Chase and Patchett, 1988; Galer

and Goldstein, 1991; McCulloch and Bennett, 1994].

Toward the end of the late Archean, massive intra-

crustal melting of newly formed late Archean crust

produced an upper crust dominated by K-rich grano-

diorites and granites, typically showing a relative de-

pletion in Eu. This change is reflected in the REE

patterns observed in the clastic sediments. Archean-

type REE patterns are swamped, and the upper crustassumed its present composition.

There is a little-understood relationship between

the growth of continents and the intracrustal melting

that shortly followed. A significant observation is that

such intracrustal melting mostly occurs within 50-100

million years of the derivation of new crust from the

mantle [e.g., Moorbath, 1977]. Perhaps the litho-

spheric keel beneath older cratons deflects mantle heatflow or rising mantle plumes toward such areas. The

development of extensive crustal areas may produce

an insulating trap for mantle-derived magmas to accu-

mulate at their base and so provide the necessary heat

for crustal melting. Another factor that may be impor-

tant in providing additional heat is the thermal pertur-

bation associated with the continental collisions that

ultimately must terminate the subduction of the oce-

anic plate.

As global heat flowdiminished in the late Archean,

the number of plates became fewer, and modern-style

plate tectonics became dominant. The result was that

the oceanic crust was subducted on a much longer

timescale and so on average was both older and colder

by the time it reached the subduction zone [e.g., Da-

vies, 1992]. Under these conditions it was dehydrated

at higher pressures during subduction, becoming re-

fractory before reaching the melting temperatures of

the basaltic crust. This older oceanic crust thus returns

to the mantle to greater depth without being remelted.

The fluids resulting from dehydration ascend into the

overlying mantle wedge, where they induce melting,

and result in the production of the present-day calc-

alkaline suite [Martin, 1986; Defant and Drummond,

1990, 1993; Drummond and Defant, 1990]. These con-ditions, which result in the addition of material of

andesitic composition to the crust, appear to have

persisted since the early Proterozoic.

The Freeboard Constraint

Flat-lying Proterozoic sediments provide evidence

that for the past 2 billion years, sea level has been

within about 1 km of the present level. This is strong

evidence that the volumes of continents relative to

oceans have been similar over this period [e.g., Wise,

1974; Armstrong, 1981; Kasting and Holm, 1992].

However, the sedimentary record in the Archean is

much too fragmentary to make freeboard a restriction

on crustal volume. Although the constraints from the

freeboard record are now considered considerably less

stringent than was previously thought [McLennan and

Taylor, 1982; Schubert and Reymer, 1985; McLennan,

1988;Armstrong, 1991; Galer, 1991], the general argu-

ments appear valid at least back into the Early Pro-

terozoic, or about halfway through the geological

record. Thus although the freeboard observation has

been one of the cornerstones of the no-growth early

crust school, it is valid for only the latter part of the

geological record and provides no constraints of

crustal volume before the Proterozoic.The freeboard constraint is accommodated in most



256. Taylor and McLennan: CONTINENTAL CRUST GEOCHEMISTRY

models by proposing that there have been relatively

minor additions to the continents since the Early Pro-

terozoic and that most crustal growth occurred by the

late Archean. It is noteworthy, however, that a num-

ber of models requiring significant crustal growth at

2.1-1.8 eons, such that the amount of crust present at

2.5 eons was less than 50% (in some models as little as

30-40%), have not come to terms with this constraint.

Andesite Models

In early estimates of crustal composition [e.g., Tay-

lor, 1967, 1977]it was assumed that the bulk crust was

equivalent in composition to that of average island arc

volcanic rocks. This formed the basis for the so-called

"andesite" model. Island arc volcanism is the princi-

pal mechanism at the present day by which material is

derived from the mantle and accreted to form the

continental crust (see McCulloch and Gamble [1991]

andHawkesworth et al. [1991, 1993]for recent reviews

of island arc magmatism). The model has major prob-

lems in that it contributes too much potassium (i.e.,

the composition generates too much heat, although

note the difficulties with interpretation of heat flow

discussed above) and other incompatible elements and

too little nickel and chromium to account for the ob-

served crustal abundances [Taylor and McLennan,

1985].Furthermore, the major growth of the continen-

tal crust appears to have occurred episodically, mostly

in the late Archean where andesites are comparatively

scarce and the TTG suite was the chief nonbasaltic

product of igneous activity. In contrast, the produc-

tion of andesites in island arcs has probably beenoperating in its present mode only for about the past 2

billion years. This process has produced only about

25-40% of the crust.

The overall crustal bulk composition in our models

was calculated from a 75/25 mixture of the Archean

bimodal and the post-Archean andesitic compositions

[Taylor and McLennan, 1985], and as the crustal

growth curve is further refined, these proportions will

also require adjustment. In this model, any addition of

material by basaltic underplating is not taken into

account.

Plumes and Basaltic Underplating

There is growing appreciation that there may be

additional mechanisms for the growth of continents,

mainly by basaltic underplating at convergent margins

[e.g., Kempton and Harmon, 1992]and by magmatism

and underplating largely associated with mantle

plumes [Hill et al., 1992;Campbell and Griffiths, 1992;

Hill, 1993]. The volumetric scale of such processes

and the implications for crustal composition are much

less understood.

Basaltic underplating at convergent margins or as-

sociated with plume activity has been suggested as an

important mechanism to explain the chemistry andmetamorphic history of lower crustal xenoliths [e.g.,


Kempton and Harmon, 1992;Rudnick, 1990, 1992a].

Rudnick [1990] suggested that as much as 20% (or

more) of the crust in the Ivrea Zone [Voshage et al.,

1990]was derived from basaltic underplating, in this

case associated with convergent margin magmatism.

On the other hand, underplating may not be a signifi-

cant process at times substantially after the major

episode of crustal growth [Wendlandt et al., 1993].

Estimates of crustal additions derived directly from

magmatism associated with continental intraplate

mantle plumes or obducted ocean plateaus are re-

stricted to no more than a few percent [Hill et al.,

1992].

Although there are a number of caveats, if basaltic

underplating is an important process for the growth of

the continents, then all models of bulk crustal compo-

sition may have to be reevaluated. It has long been

argued on petrological grounds that the average com-

position of island arcs is more mafic than andesite

[e.g., Arculus, 1981;Ellam and Hawkesworth, 1988;Pearcy et al., 1990]. Recent evidence from the

geochemistry of turbidites derived from the lzu-Bonin

Arc adds some support to this suggestion [Gill et al.,

1995].Incorporation of basaltic material at the base of

the arc crust would enhance such a model. The major

difficulty with a basaltic continental crust is that it is

difficult to differentiate such material in order to gen-

erate sufficient granite to provide for the well-con-

strained upper crustal compositions. At least one quar-

ter of crustal volume, is needed to provide for the

granitic composition observed in the upper crust. If

such basaltic material is calc-alkaline in nature, theremay be barely sufficient concentrations of heat-pro-

ducing elements to be consistent with the observed

heat flow data [Basaltic Volcanism Study Project,

1981].

Various models have been proposed to avoid such

dilemmas. One is to maintain an intermediate overall

composition by delaminating the more ultramafic com-

ponents at the base of the crust and resorbing them

into the mantle [e.g., Turcotte, 1989; Pearcy et al.,

1990; Kay and Kay, 1991]. Such mechanisms, while

worthy of study, are at this stage rather ad hoc. Ifthe

crust does indeed evolve through such convoluted

mechanisms, then estimates of its bulk composition

would be largely unconstrained.

One end-member proposes that because basalts are

the primary silicate melt from the mantle, and because

continents are ultimately derived from the mantle,

their bulk composition must be basaltic. The answer to

this paradox is of course that the siliceous composition

of the continental crust is mostly a consequence of

recycling and remelting of the primary basalts follow-

ing subduction. It is this process, apparently unique to

the Earth, that has resulted in the production of the

fractionated continental crust. The lack of such recy-

cling on other planets and satellites has precluded thedevelopment of continental crust on those bodies.




Subcrustal lithosphere

A crucial problem, closely related to crustal

geochemistry, is the nature of the subcrustal litho-

sphere. Are there deep roots to the continents? If so,

what is their relationship to the overlying crust? There

is good evidence that fast seismic P wave velocities

extend to depths of 220 km. Lateral variations in S

wave velocities extend much deeper, to 400 km. Thus

there appear to be deep keels to the continental re-

gions, which are relatively cold and appear to be re-

fractory in composition [Jordan, 1981, 1988;Boyd and

Gurney, 1986].One interpretation of the heat flowdata

[Nyblade and Pollack, 1993;Bickle, 1986]is consistent

with a much thicker lithosphere beneath Archean cra-

tons than beneath those of younger age. This view has

received considerable support from studies of inclu-

sions in diamonds that indicate that low temperatures

have persisted beneath shields for up to 3000 m.y.

[Richardson et al., 1984;Walker et al., 1989;Pearson

et al., 1995]. The general interpretation from the seis-mic, petrological and isotopic evidence is that these

deep roots represent refractory, lower-density mate-

rial, consistent with their being the residue remaining

from the extraction of a Fe-rich partial melt, leaving

behind an Mg-enriched zone of lower density.

Whether these roots are residual from the extraction of

basaltic oceanic crust or are connected with the ex-

traction of continental crust is an interesting but

largely unresolved question.

T H E N AT U RE O F C R US T AL G R OW TH

Two major hypotheses have been advanced to ex-

plain the evolution of the continental crust. The first

proposes that the present mass of the crust formed

very early in Earth history and has been recycled

through the mantle in steadily decreasing fashion such

that new additions are balanced by losses, resulting in

a steady state system [Armstrong, 1968, 1981, 1991].

The second proposes that the crust grew throughout

geological time, with variations from a steady rate to

growth in major episodic pulses [e.g., Moorbath, 1977;

Taylor and McLennan, 1981]. The REE abundances

shed considerable light on these problems. As was

noted above, there is a major dichotomy in the sedi-

mentary record at about the Archean-Proterozoic tran-

sition (ca. 2.5 eons), and this has been correlated with

an earlier major pulse of crustal growth during the late

Archean.

There is growing evidence (e.g., from lOBeand Pb

isotopes) that sediments have been recycled into the

mantle and are important in controlling the composi-

tion of island arc and perhaps intraplate volcanics (see

reviews by McLennan [1988]and White [1989]). The

evidence from lOBe and various other geological,

geochemical, and isotopic constraints limit the amountof subducted sediments to a few percent. However,

those models that propose massive recycling of the

crust through the mantle, on a scale large enough to

maintain a steady state crustal mass, encounter vari-

ous difficulties. Data from long-lived radiogenic iso-

topes for mantle-derived rocks provide no indepen-

dent constraints on this problem, a point long made by

Armstrong [1968, 1981, 1991]. One constraint, from

the mass-age distribution and Nd isotopic characteris-

tics of sedimentary rocks, is that during various sedi-

mentary recycling processes the sedimentary mass

receives only about 15%new material from the mantle

[Veizer and Jansen, 1985;McLennan, 1988].Thus the

mass of sediment available for subduction is <1.6 x

1015g yr " (about 0.5 krrr' yr-I of crust), Such an

amount likely provides insufficient material to support

a steady state crustal mass [McLennan, 1988].

Two observations inform us of the limited extent of

very early "granitic" crusts. First, the small expo-

sures of early Archean potassic granites do not con-

tribute a negative Eu signature to the greenstone beltsediments. There was no land vegetation in the Ar-

chean [Cloud, 1988]. Hence there must have been

wide exposures of bare rock, with the consequent

formation and wind transport of dust. Present-day

mid-ocean sediments show the telltale signature of Eu

depletion of the upper continental crust derived largely

by wind transport [Ben Othman et al., 1989;Lin, 1992]

(Figure 2). This signature of Eu depletion is missing

from virtually all Archean sediments, including those

from greenstone belts (Figure 3) and from many high-

grade terranes, except for the local cratonic areas

discussed, for example by Taylor et al. [1986].Accord-ingly, most of the exposed crustal rocks from which

the Archean sediments were derived were basalts and

the TTG suite, rather than the now familiar granites

and granodiorites that constitute the present exposed

upper crust of the continents.

The second crucial observation is that of Stephen-

son and Patchett [1990],who failed to find evidence of

extensive pre-3.5-eon recycled zircon populations that

would have been indicative of large areas of old gra-

nitic crust. In this study they analyzed zircons, mainly

from quartzites, from the Canadian Shield and the

Wyoming, North Atlantic (Labrador, Greenland,

Scotland) and Kaapvaal (Southern Africa) cratons.

Zircon, being a highly durable mineral, survives

through many cycles of weathering, erosion, and dep-

osition of sediments. It also appears to be aerodynam-

ically suited for large-scale wind transport, as is wit-

nessed by enrichments seen in loess deposits [Taylor

et al., 1983] and the presence of substantial exotic

zircon in ancient soils [Brimhall et al., 1990].Stephen-

son and Patchett found that the age of the zircon

populations, dated by the 176Lu_176Hfechnique, in

early Archean quartzites, was usually the same age as

the terrane in which they are found and that there is a

scarcity or absence of zircons of significantly olderage. If there was an early sialic crust, then a popula-




tion of ancient zircons derived from it by erosion

should have survived and been recycled into younger

Archean sediments. A. P. Nutman (personal commu-

nication, 1994) finds no older zircon populations in

3.8-eon metasediments from Isua, Greenland, using

U-Pb dating by ion microprobe.

Armstrong[1991]attempted to refute this argument

by noting a similarity between the sedimentary zircon

ages and the average age of crust (at the time of

sedimentation) predicted by Armstrong's [1981]

steady state model. The difficulty with this is that

sedimentary rocks are not derived solely from contem-

poraneous "average" crust but instead are typically

derived substantially from older, cannibalistically re-

cycled sedimentary rocks [e.g., Veizer and Jansen,

1985] and would be expected to have components

older than "average" crust. Thus younger Proterozoic

sediments, in contrast, commonly contain a significant

population of zircons derived from rocks of late Ar-

chean age. The conclusion from this study is that old

Archean sialic continental crust was rare and not sub-

stantially more extensive than the few remnants that

are currently exposed.

The efficient averaging carried out by the processes

of sedimentation and the buffering effects of cannibal-

istic sedimentary recycling probably conceals small-

scale variations in the rate of continental growth and

perhaps even in upper crustal composition. Thus there

is considerable evidence for a period of growth at

about 1800m.y. ago that is not clearly reflected in the

sedimentary geochemical record. The major change in

upper crustal composition in the late Archean and therecycling of sediments effectively swamp out the ef-

fects of later, smaller episodes. Detailed sedimento-

logical, geochemical and isotopic studies should un-

cover these effects [e.g., McLennan et al., 1995].The

situation was different in the early Archean. Small

cratonic areas were isolated, resulting in the preserva-

tion of individual episodes of cratonization, now re-

flected in the local variations in REE patterns.

Nd model ages are now well established as a means

of estimating crustal formation ages [McCulloch and

Wasserburg, 1978;DePaolo, 1981].During extraction

of material from the mantle and its incorporation into

the continental crust, major chemical fractionation of

Sm and Nd occurs. Intracrustal melting events lead to

further fractionation of Sm and Nd, although this gen-

erally follows crust formation within 100-200 m.y. and

is comparatively minor. The Nd model ages of sedi-

mentary rocks in general give the average time of

formation of the various sources of the sediments

rather than the age of deposition. This assumes that

Sm and Nd are not fractionated during sedimentation

or diagenesis (see above).

Evidence regarding possible changes in upper

crustal composition during the post-Archean can be

obtained from the Nd model ages of sediments, asdiscussed earlier (see section on loess). For post-Ar-

chean sediments these range from about 2.5 to 1.0

eons and represent the average age of upper crustal

material which was the source of these sediments.

However, post-Archean sediments have very uniform

REE patterns and other key geochemical features

(e.g., Th/Sc) despite the wide range of provenance

age. Minor compositional changes in the upper crust

may be masked to the sedimentary record by the

buffering effects of cannibalistic sedimentary recycling

[McLennan, 1988;McLennan and Hemming, 1992].

Nevertheless, within these uncertainties, this indicates

that the processes forming the upper crust must have

been uniform over that interval and that any additions

to the upper crust during the post-Archean could not

have differed significantly in composition.

In summary, the crust continues to grow at present

by island arc volcanism and related magmatism, fol-

lowed by episodes of intracrustal melting (e.g., eastern

Australia). There has been no significant change in this

process since the Archean-Proterozoic boundary that

is detectable in the sedimentary REE record. The

growth of the continental crust has proceeded in an

episodic fashion throughout geological time [e.g.,

Moorbath, 1977, 1978; Taylor and McLennan, 1981,

1985], with a major increase in the growth rate in the

late Archean (Figure 5).

Supercontinental Cycles

There is a growing consensus that periodically the

continental masses unite to form one or two supercon-

tinents [e.g., Nance et al., 1986;Hoffman, 1988, 1992;

Unrug, 1992; Anderson, 1994]. These episodes arefollowed by times, such as the present, when the

continents are fragmented and dispersed. The recog-

nized periods of supercontinental accumulations

[Nance et al., 1986; Young, 1991; Unrug, 1992] are

during the Late Paleozoic (Pangea; 0.3 eons), Late

Proterozoic (0.9 eons), possibly at about 1.8 eons (for

North America/Baltica at least) and perhaps during the

late Archean (ca. 2.7 eons). During these episodes, sea

level appears to be generally low (i.e., lowstands of

first-order sea level fluctuations) probably due to the

maturing of ocean basins associated with relatively

minor plate motion. Among other things, this conti-

nental configuration appears to correlate with major

climatic fluctuations [e.g., Young, 1991] and funda-

mental biological changes such as the radiation of

stromatolites and mass extinctions.

Apart from the geographical configuration of conti-

nents, supercontinental cycles also likely have an im-

portant connection with crustal evolution [McLennan

and Taylor, 1991].For example, supercontinental cy-

cles may explain the apparent periodicity of crust-

forming events during the post-Archean. Apart from

the early Archean, these episodes of crustal growth,

generally taken to have occurred at about 3.8-3.5,

2.9-2.7,2.0-1.7, 1.3-1.1, and 0.5-0.3 eons [e.g., Me-Culloch and Bennett, 1993], appear to correlate rea-




IntenseMeteoriteBombardment

High T / Low P Modern Island arc regime< "" ;:>-

Subduction regime10 0

4.0 3.0 2.0

A ge (A e)

Q)

80E~(5

>60

C ii

-J)

~. . .40 o

20

0

0.0

Major episodes of

crustal growth

Major supercontinents In place

1.0

Figure 5. Schematic model for the growth and evolution of the continental crust. The actual values of crust

present at any given time are not well constrained; however, a value of 50% crust by about 2.5 eons is

likely a minimum value to satisfy freeboard constraints. Although major global episodes of crustal growth

and differentiation are well documented during the late Archean and at about 2.1-1.7 eons, it is less clearif crustal growth is episodic on a global scale during younger times. In general, there appears to be a

relationship between crustal growth episodes and assembly phases of supercontinents.

sonably well with supercontinental assembly phases.

During supercontinental assembly, subduction of old

oceanic crust would be at a maximum, leading to an

increase in arc volcanism [Nance et al., 1986; Zhong

and Gurnis, 1993]. In addition, the continental colli-

sions that terminated the assembly phases would

promote crustal thickening, intracrustal melting and

differentiation, granulite facies metamorphism, conti-

nental underplating, and, accordingly, preservation of

such crust in the geological record.

On the other hand, the strong contrast in crustal

geochemistry seen at the Archean-Proterozoic transi-

tion does not appear to be related simply to the accu-

mulation of a late Archean supercontinent [McLennan

and Taylor, 1991]. Comparable changes in composi-

tion are not seen during younger cycles. The massive

increase in crustal growth in the later Archean is most

probably related to higher mantle heat production.

Intracrustal differentiation of this newly formed crust

to form stabilized continents resulted in a widespread

"granitic" upper crust (as opposed to localized earlyArchean differentiated cratons) possessing the charac-

teristic negative Eu anomaly, enrichment in heat-pro-

ducing elements, and other geochemical characteris-

tics. Subsequent episodes of crustal growth appear to

have followed a similar pattern because they have

produced little change in upper crustal compositions

[e.g., McLennan and Hemming, 1992].

Causes of Continental Breakup

Another interesting question, although somewhat

peripheral to this review, is the cause of continental

breakup [e.g., Duncan and Turcotte, 1994]. Because

such rifting is frequently accompanied by vast out-

pourings of basaltic lavas, a major debate has arisen

about whether such basaltic magmatism is the cause of

the breakup or merely a consequence of decompres-

sion melting of the mantle following the thinning of the

overlying crust [e.g., Storey et al., 1992]. Related to

this question are the role that deep mantle plumes play

in this process [Hill, 1993] and the influence of super-

continental cycles in controlling mantle convection

patterns [Gurnis, 1988; Zhong and Gurnis, 1993]. Do

mantle plumes initiate splitting of the crust by doming?

These chicken and egg arguments are extensively dis-

cussed by Storey et al. [1992]. Thus White [1992]

distinguishes various types of rifting that occur during

continental breakup. The extreme cases may be called

cold and hot rifts, respectively. During cold rifting

there is very little volcanism, and the continental rift is

"stretched like gum." When it breaks, this allows the

formation of a new oceanic spreading center. Hot rifts,

in contrast, form as a consequence of massive volca-

nism originating from mantle plumes. Very large vol-

umes (several kilometers thick) of lava are extruded.Thus mantle plumes, although not the cause of exten-

sion, may provide a trigger in the shape of additional

heat that boosts the splitting and breakup of the con-

tinents. Perhaps the most difficult aspect for plume-

related models is the linear nature of most rifting and

of many large igneous provinces, such as the 3000-km-

long Ferrar- Tasman Province.

T H E EARLIEST PRE-3.9 EON CRUSTS

The oldest preserved terrestrial rock is the 3.96-eon

Acasta Gneiss, in the Northwest Territories of Canada




[Bowring et al., 1990]. What happened in the 500-m.y.

gap between the formation of the Earth and that of the

Acasta Gneiss? In order to place this question in an

appropriate perspective, it is necessary to consider theproblems encompassing both the formation of this

planet and some associated events.

The gaseous parts of the solar nebula were dis-persed very quickly on timescales of a few million or

less years following To (4.56 eons) by early violent

solar activity [Strom et al., 1989]. The Earth then

accreted from a hierarchy of rocky planetesimals in a

gas-free environment [Wetherill, 1986, Taylor, 1992a,

1993a] between 10 and 100 m.y. after To. The Moon

was probably formed during the final stages of plane-tary accretion by the collision with the Earth of a body

0.15 Earth masses, somewhat larger than Mars [Benz

et al., 1989],with derivation of the Moon from the rocky

mantle of the impactor [Newsom and Taylor, 1989].

High temperatures are a consequence of the formation of

the Earth from planetesimals, and total melting of theEarth seems unavoidable. Indeed the lunar-forming

event was sufficiently energetic to have accomplished

this on its own [Benz and Cameron, 1990].

In an initially molten Earth, separation of the me-

tallic core is effectively coeval with accretion. The

subsequent crystallization history of the terrestrial

mantle is a matter of intense geochemical debate. A

number of things did not happen. The crystallization of

large bodies of molten rock, such as the Skaergaard or

Stillwater layered intrusions, might lead one to expect

the production of zones of differing mineralogy whose

elemental and isotopic fingerprints would inform us ofthis early event in the mantle. This was the course ofevents on the Moon. However, no terrestrial an-

orthositic crust, analogous to the lunar crust, formed

for various reasons: the Earth is less well endowed

with Ca and Al than the Moon, so that plagioclase willnot be an early crystallizing phase from a terrestrial

magma ocean; the massive Eu anomalies produced, as

on the Moon, by such an event, are missing; plagio-

case floated on the bone-dry lunar magma ocean but

would have sunk in the wet terrestrial magma; plagio-

clase is stable on the Earth only to 40 km, below whichit will react to form garnet. In addition, there is no sign

of a reservoir of primitive 87Sr/86Srvalues from such a

feldspathic crust [Taylor, 1982, 1992b].

The molten mantle was so large that conventional

terrestrial analogies, derived from layered intrusions

or the Moon, which is 2 orders of magnitude smaller,

are inadequate to describe the solidification of a mol-

ten terrestrial mantle. Probably, the mantle crystal-

lized without producing separate mineral zones [Tonks

and Melosh, 1990].

No good evidence exists for that enduring geologi-cal myth of a primordial world-encircling crust of

"sial" or granite. Such models originated through false

analogies between the production of a silicic residuumduring crystallization of basaltic magmas and conditions

in an early molten Earth, and a lack of appreciation of

the difficulties of producing granite. The absence of ev-

idence for such a crust is discussed earlier. A primitive"sialic" crust is refuted, among other evidence, by the

absence of an old zircon population in younger sedi-

ments [Stephenson and Patchett, 1990].

The conditions for the production of massive granitic

crusts are probably unique to the Earth and require three

or more stages of derivation from a primitive mantle

composition. The Earth has transformed less than 0.5%

of its volume to continental crust of intermediate com-

position and less than 0.2% of its volume into granitic

continental crust (i.e., upper continental crust) in over 4

billion years, so that the process is inefficient. The high-

land feldspathic crust of the Moon, about 12% of lunar

volume, formed in contrast during crystallization from a

magma ocean within a few million years.Regardless of how the terrestrial mantle crystal-

lized, there is extensive evidence for some probably

localized early mantle differentiation. Rocks derivedfrom the upper mantle from 3900 m.y. ago onward

come from a source with high Sm/Nd ratios. Although

the evidence refers specifically to Sm and Nd, by

inference the other LREE and other incompatible el-ements (e.g., K, Rb, Cs, Ba, Th, U, Zr, and Hf) were

also depleted in parts of the Archean mantle at a very

early stage [Galer and Goldstein, 1991]. This depletion

must have occurred well before the time of formation

of the oldest surface rocks. Although many Archean

rocks are derived from such a LREE-depleted (Sm/

Nd-enriched) reservoir, this may merely indicate that

during the Archean the mantle depletion in LREE wasrestricted to a small mantle reservoir (that presumablygrew in approximate proportion to crustal growth) and

not a planet-wide affair [McCulloch and Bennett, 1993,

1994]. After all, caution is warranted, since Archean

rocks now cover only a trivial percentage of the sur-face of the Earth and so provide only limited informa-

tion on the state of the bulk mantle.

Apparently, portions of the upper mantle under-

went partial melting, which led to an increase in the

Sm/Nd ratio in the mantle as the LREE were extracted

into the melt. Extraction of an early crust is an obviousway to carry this out, but early granitic or anorthositic

crusts are unlikely candidates as noted earlier. Basalts,

being the primary melt from silicate planetary mantles,

are the most likely candidates for an early crust.

However, the basic problem is that if this primor-

dial early basaltic crust ever existed, it seems to have

vanished into the mantle without providing the finger-

print of its enriched LREE signature to rocks subse-

quently derived from that region. No Archean or later

rocks have been derived from such a primitive crust

(which would resemble an alkali basalt and be en-riched in LREE, with low Sm/Nd ratios [e.g., Galer

and Goldstein, 1991]).

Another possibility is that the earliest crust wascomposed of komatiite [Campbell and Jarvis, 1984;




Chase and Patchett, 1988].Such high-Mg melts might be

produced under the early high-temperature conditions inthe Hadean (pre-4.0 eons [Cloud, 1988])Earth. Such a

crust would be too refractory to melt during subduction,

so it could have been recycled into the deep mantlewithout the telltale signature of LREE and incompatible

element enrichment subsequently reappearing in surfacerocks. However, komatiitic rocks are produced by highdegrees of partial melting and are unlikely to be enriched

in incompatible elements to the degree required.

The probable explanation is that only small regions

of the early Archean mantle underwent partial melting

and that the early LILE-enriched mafic crust was

buried deep in the mantle and is no longer accessible.

Any crust existing before 4.0 eons was probably

highly disrupted and perhaps destroyed by the ongoing

meteoritic bombardment which was endemic through-

out the solar system down to about 3.9 or 3.8 eons.

Major basin-forming (>300 km diameter) events

ceased on the Moon only at that time. In the interval

between about 4.4 and 3.8 eons, at least 200 such

major collisions occurred on the Earth. This evidence

may need to be reevaluated if most of the basin-

forming events on the Moon were due to a late spike or

"cataclysm" in the cratering record around 3.8-4.0

eons [Grieve and Parmentier, 1985; Ryder, 1990].

U NIQ UE NE SS O F T E RR EST R IAL C O NT IN EN TAL

C R U ST IN T H E S O L AR SY ST E M

The formation of crusts distinct from the bulk plan-etary composition is a common planetary phenome-

non. However, no crusts similar to the continental

crust of the Earth appear to have formed on the other

terrestrial planets or on the icy satellites of the giant

planets. On the inner planets these crusts are basaltic,

which are the typical partial melts derived from the

Fe-Mg silicate mantles. On the satellites of the giant

planets, crusts involving water, methane, or ammonia

ices form by melting of ice-rock mixtures. All such

crusts have been termed secondary [Taylor, 1989] be-

cause they form as a result of melting in planetary

interiors subsequent to planetary formation.The Moon presents a special case of a primary crust

[Taylor, 1982] composed essentially of anorthosite,

which floated on an anhydrous magma ocean. This

was a consequence of total melting at the time offormation of the Moon. There is no evidence of a

similar crust forming on the Earth, as was noted ear-

lier, nor any sign of a lunar granitic crust. The largest

"lunar granites" are tiny «2 g) fragments of siliceous

differentiates [Taylor, 1982].

Venus is the closest planet in size and composition

to the Earth, and accordingly might be expected to

have had a rather similar geological history. However,

both planets show major differences [e.g., Kaula,

1990; Taylor, 1993b] and the Venusian surface appears

to be mostly basalt, perhaps with a few scattered

"pancakes" possibly of more siliceous differentiates

[e.g., Basilevsky et al., 1990]. The high-standing re-

gions (e.g., Ishtar Terra, Aphrodite Terra) appear to

consist of crumpled-up basaltic crust. The planet is

apparently unable to subduct its basaltic crust, possi-

bly because of the greater depth of the basalt-eclogitetransition (70-80 km) than on Earth (40 km), a conse-

quence of the slightly smaller size of Venus. The

planet appears to have choked itself on basalt without

being able to produce more differentiated rocks.

The northern crust of Mars appears to be basaltic,

consistent with the geomorphic evidence for basaltic

plains and volcanoes. The two Viking landers, 4000

km apart in the northern hemisphere, gave similar

basaltic compositions for the material analyzed. The

fine material analyzed by the Viking landers will con-

tain a significant component from the ancient cratered

terrane that dominates the southern hemisphere,

where most of the global dust storms originate. How-

ever, no component more siliceous than basalt appears

in the Viking compositions. Accordingly, there is no

evidence for a granitic terrane [Taylor, 1992a].

There is little evidence about the crust of Mercury,

except that it has a reflectance spectrum similar to thatof the Apollo 16 highlands [Vilas et al., 1988, p. 71]

and hence is unlikely to be granitic. Enigmatic plains

units may be basaltic, but differ in albedo from the

lunar maria.

The significant feature about the Earth, in contrastto the other terrestrial planets, appears to be the pres-

ence of liquid water at the surface, coupled with platetectonics and subduction, that enables recycling of

subducted basaltic crust through the mantle. It is this

process that permits the slow production of the conti-

nental crust [e.g., Campbell and Taylor, 1983]. In

other planets the absence of subduction leads to the

persistence of barren basaltic plains such as we ob-

serve on Venus and Mars.

ACKNOWLEDGMENTS. We are grateful to Owen

Evans, Gil Hanson, Klaus Mezger, and Roberta Rudnick for

helpful discussions. Useful reviews were provided by Alan

Chave, Bill Curry, Roberta Rudnick, Steve Shirey, and Tom

Torgersen. S.M.M. was supported by the National Science

Foundation under grant EAR-8957784. He is also grateful for

the support from the Deutscher Akademischer Austausch-

dienst and the Max-Planck-Institut fur Chemie(Abteilung

Geochemie) during the final stages of preparing the manuscript.

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