Volcanic Rifted Margins

INTRODUCTION

Volcanic rifted margins (Fig. 1) are produced where conti-nental breakup is associated with the eruption of ood volcan-ism during prerift and/or synrift stages of continental separation(Fig. 2) (Mutter et al., 1982; White et al., 1987; Holbrook andKelemen, 1993; Eldholm and Grue, 1994; Courtillot et al.,1999). These margins are easily distinguished from nonvol-canic margins, like the Iberian margin, that do not contain sucha large amount of extrusive and/or intrusive igneous rock andthat may exhibit unusual features, such as unroofed mantle peri-dotites (e.g., Pickup et al., 1996; Louden and Chian, 1999).Mapping of ood basalt provinces and subsurface seismic vol-canic-stratigraphic analyses show that volcanic rifted marginsborder the northern, central, and southern Atlantic Ocean, the

southern Red Sea, the east coast of Africa, circum-Madagascar,the east and west coasts of India, the western and eastern coastsof Australia, and possibly parts of Antarctica (Cofn and Eld-holm, 1992, 1994; Mahoney and Cofn, 1997; Planke et al.,2000) (Fig. 1). The initiation of a ood basalt province (or of alarge igneous province [LIP]) (Fig. 2) is commonly a preriftphenomenon and takes the form of subaerial basaltic and/orsilicic volcanism (e.g., Cox, 1988; Renne et al., 1992; Menzieset al., 1997a; Larsen and Saunders, 1998). The prerift to synrifttransition is marked by a structural change, in some cases a mag-matic hiatus, erosion of newly formed rift mountains, and theformation of high-velocity lower crust (HVLC), and a seaward-dipping reector series (SDRS) (Mutter et al., 1982; White et al.,1987; Eldholm and Grue, 1994; Planke et al., 2000) (Fig. 2).SDRS comprise subaerial and submarine volcanic rocks and

Geological Society of AmericaSpecial Paper 362

2002

Characteristics of volcanic rifted margins

Martin A. Menzies*

Department of Geology, Royal Holloway, University of London, Egham, Surrey TW20 OEX, UK

Simon L. Klemperer

Department of Geophysics, Stanford University, Stanford, California 94305-2215, USA

Cynthia J. Ebinger

Department of Geology, Royal Holloway, University of London, Egham, Surrey TW20 OEX, UK

Joel Baker

Dansk Lithosfrecenter, Oster Voldgade 10, 1350 Kobenhavn K, Denmark

ABSTRACT

Volcanic rifted margins evolve by a combination of extrusive ood volcanism, in-

trusive magmatism, extension, uplift, and erosion. The temporal and spatial relation-

ships between these processes are inuenced by the plate tectonic regime; the preexist-

ing lithosphere (thickness, composition, geothermal gradient); the upper mantle

(temperature and character); the magma production rate; and the prevailing climatic

system. Of the Atlantic rifted margins, 75% are believed to be volcanic, the cumulative

expression of thermotectonic processes over 200 m.y. Volcanic rifted margins also char-

acterize Ethiopia-Yemen, India-Australia, and Africa-Madagascar. The transition from

continental ood volcanism (or formation of a large igneous province) to ocean ridge

processes (mid-ocean ridge basalt) is marked by a prerift to synrift transition with for-

mation of a subaerial and/or submarine seaward-dipping reector series and a

signicant thickness (to 15 km) of juvenile, high-velocity lower crust seaboard of the

continental rifted margin. Herein we outline the similarities and differences between

volcanic rifted margins worldwide and list some of their diagnostic features.

1

*E-mail: [email protected].

Menzies, M.A., Klemperer, S.L., Ebinger, C.J., and Baker, J., 2002, Characteristics of volcanic rifted margins, in Menzies, M.A., Klemperer, S.L., Ebinger, C.J.,and Baker, J., eds., Volcanic Rifted Margins: Boulder, Colorado, Geological Society of America Special Paper 362, p. 114.

probably variable amounts of sedimentary detritus shed from thevolcanic rifted margin during uplift and tectonic denudation ofthe kilometer-scale rift mountains. The formation of SDRS is as-sociated with the establishment of thicker than normal oceaniccrust, seaward of the rifted margin at the continent to ocean tran-sition (Fig. 2). Eventually stretching and heating lead to effec-tive rupture of magmatically modied continental lithosphere,and seaoor spreading commences. This early oceanic crustmay be thicker than normal owing to hotter asthenosphere as-sociated with the plume and/or steep gradients at the litho-sphere-asthenosphere boundary (e.g., Boutilier and Keen, 1999)(Fig. 2). The interval between the rst expression of volcanicrifted margin formation on the prerift continental margin and theformation of true ocean oor can be tens of millions of years(Fig. 3).

In this paper we focus on evidence from a few of the better-known volcanic rifted margins, Ethiopia-Yemen, the Atlanticmargins, and the Australia-India conjugate margins. Miocene torecent volcanic rifted margins (

margins, the Central Atlantic magmatic province, formed ca. 200Ma (Holbrook and Kelemen, 1993; McHone, 1996; Hames et al.,2000) and records the initial breakup of Pangea. This volcanicrifted margin has been reduced to erosional remnants of intrusiveand/or extrusive complexes (McHone, 1996) and offshore SDRS(Benson, 2002) spread over 106 km2 (Fig. 1). Older intraplateood basalt provinces erupted in the Permian-Triassic occur inSiberia, Russia, and Emeishan, China. However, their relation-ship to rifted margins is unknown, and they are not discussedherein. In addition, we do not consider relics of Precambrianood basalt provinces that are apparent as dike swarms and un-roofed plutonic complexes, because links to continental breakupare even more elusive (Mahoney and Cofn, 1997, and refer-ences therein).

We use these better known volcanic rifted margins as wediscuss some of the controversies surrounding the origin of vol-canic rifted margins (whether associated with active or passiverifting) and describe some of their diagnostic features, i.e., com-mon association of silicic volcanism with the dominant oodbasalts; crustal architecture of HVLC and SDRS at the conti-nent-ocean transition; temporal relation between extension andmagmatism; and rift-margin uplift and mountain building.

Passive continental rifted margins: Plate-driven

and plume-driven processes

Traditionally, passive (plate driven) and active (plumedriven) rifting models were invoked as an explanation for theformation of nonvolcanic and volcanic passive margins, respec-

tively. Passive or plate-driven rifting models required that con-tinental breakup was initiated by extensional forces, followed bysurface uplift and magmatism related to the passive upwellingof normal asthenospheric material. In this case melts would begenerated by shallow decompression melting processes. In con-trast, active or plume-driven rifting models required deeper meltgeneration and subsequent interaction with the continental litho-sphere. In these models the expectation is that volcanic riftedmargins formed by kilometer-scale surface uplift prior to LIPformation and extension. However, recent research on volcanicrifted margins indicates that such simple chronologies do notapply to many rifted margins, suggesting that their formation isalso not simple. The timing of uplift and extension relative toLIP formation is complex, requiring more detailed observationsin individual provinces (Table 1).

Lithospheric thinning is an uncontested requirement forvolcanic rifted margin formation. Extensional forces largeenough to initiate rifting are generated by the presence of hot,low-density asthenosphere and subsequent heating of mantlelithosphere (Crough, 1978). More controversial is the mecha-nism responsible for the production of large volumes of basalticvolcanism at the Earths surface, which in the majority of casesis spatially and temporally related to continental breakup(Watkeys, this volume). Adiabatic decompression melting dueto active upwelling of normal asthenosphere is triggered bylithospheric thinning, as observed at ocean ridges. This occurswithout a thermal anomaly and/or plume and may have led toconsiderable melt production on rifted margins (e.g., Holbrookand Kelemen, 1993; Boutilier and Keen, 1999; Korenaga et al.,

Characteristics of volcanic rifted margins 3

Figure 2. Schematic volcanic rifted margin based on data from Ethiopia-Yemen and the Atlantic margins. Volcanic rifted margin is characterizedby a subaerial ood basalt province (i.e., large igneous province [LIP]) with upper and lower crustal magmatic systems; prerift to synrift transi-tion with extended continental crust, and formation of a subaerial inner seaward-dipping reector series (SDRS); development of high-velocitylower crust (HVLC) in the transition from continental to oceanic domains; formation of submarine outer seaward-dipping reectors (SDRs), andan ocean basin (i.e., mid-ocean ridge basalt). In this example volcanism has extrusive episodes that are prerift (ood basalts), synrift (inner sub-aerial SDRS) and synrift and/or postrift (outer submarine SDRS).

2000). Alternatively, thermal anomalies or plume processes arebelieved to be a vital prerequisite for the generation of large vol-umes of melt (e.g., White and McKenzie, 1989; Sleep, 1996;Ernst and Buchan, 1997). In such cases the mantle potential tem-perature is elevated above that of the normal asthenosphere(~1300C). However, there is considerable controversy overwhether plumes initiate rifting, or rifting focuses plume activity(e.g., King and Anderson, 1998; White and McKenzie, 1989;Ebinger and Sleep, 1998; Nyblade, this volume). Controversyalso continues with regard to the geometry of plumes; their tem-perature; their depth of origin; and their chemical identity (e.g.,Turcotte and Emerman, 1983; Richards et al., 1989).

Geophysical and geochemical evidence, claimed as proof forthe origin of plumes in the deep or shallow mantle, is equivocal.Mantle tomography models show distinct low-velocity zones atthe core-mantle boundary, but their continuity with upper mantlelow-velocity zones may be ambiguous. Ray dispersion compli-cates the simultaneous resolution of the width and depth of par-ticular features in the upper mantle (Shen et al., 1998). However,as more and more receiver-function studies are undertaken inplume provinces, an exciting new plume detection method hasevolved that allows direct measurement of the 410 km and 670km discontinuities. As a result, the locations of plume stems canbe mapped (Wolfe et al., 1997; Shen et al., 1998). For petrologists

and geochemists the controversy surrounds the identication ofvolcanic rocks with unequivocal primary mantle signatures. Lowmagnesian volcanic rocks that have undergone low-pressure frac-tionation are obviously inappropriate probes of high-pressuremantle processes. Even highly magnesian unfractionated vol-canic rocks can inherit the chemical signature of the lithosphere(crust and mantle) because of their higher temperature. However,advances have been made toward the identication of geo-chemical criteria that may help resolve mantle and/or crustalcharacteristics (e.g., Baker et al., 2000; Breddam et al., 2000;Melluso et al., this volume; Baker et al., this volume).

What may be stated, with some certainty, is that large-scalethermal anomalies, plumes, and/or hotspots exist in the Earth,and their distribution, temporal evolution, and spatial extent arehighly variable. Variations may arise because of preplumelithospheric structure and differences in the velocity of platesover plumes. Slow moving plates may show larger volumes ofmelting than fast moving plates. In addition, their longevity andgenesis may manifest as uplift, subsidence, and extension of thelithosphere and associated magmatism, occurring to differentdegrees and in various sequences. In volcanic rifted marginsmelts are produced by variations in pressure and temperature,and these can be achieved, respectively, by lithospheric thinningand thermal anomalies.

4 M.A. Menzies et al.

Figure 3. Formation of ood basalt provinces (large igneous provinces [LIPs]) on volcanic rifted margins and timing of formation of oceaniccrust. Subaerial or submarine seaward-dipping reector series (SDRS) characterize prerift to synrift (continent to ocean) transition and predateoldest oceanic crust on that volcanic rifted margin. Note that in most instances ood volcanism or LIP formation precedes breakup and forma-tion of ocean crust. High-velocity lower crust and SDRS tend to form after the main ood basalt episode and before the youngest ocean crust.See text and table 1 for references.

Etendeka - Parana Flood Volcanism

Australia - India

Greenland - U K Flood Volcanism

Ethiopia - Yemen Flood Volcanism

CAMP

?

?

?

Oceanic Crust Flood VolcanismKEY

Southern Red Sea Oceanic Crust

0 50 100 150 200 250

0 50 100 150 200 250

Ethiopia -

Yemen

Greenland

- UK

Australia -

India

Africa -

South America

Central Atlantic

Northeast Atlantic Oceanic Crust

Oceanic Crust

Oceanic Crust

Flood volcanism, extension and seafloor spreading

TA

BL

E 1

.C

HA

RA

CT

ER

IST

ICS

OF

VO

LC

AN

IC R

IFT

ED

MA

RG

INS

LIP

:LIP

:LIP

:A

ge o

f sili

cic

LIP

and

LIP

:P

resence

Pre

sence

HV

LC

:

Pre

sent-

day

Peri

od o

f vo

lcanic

rocks:

tecto

nic

s:

Estim

ate

d

and/o

r absence

or

absence

Pre

sence o

f

thic

kness o

f eru

ption o

f pre

-basaltic

, syn-

pre

-rift,

scale

of

of a s

eaw

ard

of hig

h v

elo

city

>10 k

m o

f new

sub-a

eri

al

70%

80%

of th

e

basaltic

or

post-

syn-r

ift

pre

-magm

atic

dip

pin

g r

eflecto

r(~

7.4

km

/s)

mafic igneous

volc

anic

rocks

basaltic

rocks

basaltic

eru

ptions

or

post ri

ft?

uplif

t seri

es

low

er

cru

st

cru

st

1a

>2 k

mB

asalt-r

hyolit

eS

ynchro

nous w

ith

Pre

-rift

Not know

n

S

ub-a

eri

al

Yes

Not know

n

Eth

iopa

(2931 M

a)

basaltic

eru

ptions

magm

atism

buri

ed b

y

inner

SD

RS

(base n

ot date

d)

2631 M

ari

ft a

ctivity

1b

>2 k

m:ori

gin

ally

Basaltic

Post-

basaltic

Pre

-rift

10100 m

(m

ari

ne

denuded r

em

nant

Yes

Not know

n

Yem

en

ca.4 k

m w

ith

eru

ptions

eru

ptions

magm

atism

to c

ontinenta

lof in

ner

SD

RS

ca.2 k

m lost

2931 M

a2629 M

atr

ansitio

n in

and b

uri

ed

to e

rosio

nsedim

ents

)o

ute

rS

DR

S

2a

57 k

m5356 M

aIn

trusio

ns,

Pre

-rift, s

yn-r

ift

100s

SD

RS

n

oYes

Yes

Gre

enla

nd

no v

olc

anic

sand p

ost ri

ftm

ete

rssedim

ents

report

ed

magm

atism

report

ed

2b

ca.1 k

m:heavily

5861 M

a a

nd

Syn-b

asaltic

Pre

-rift

Unknow

n b

ut

SD

RS

m

ostly

Yes

Rockall

Yes

under

United

denuded m

arg

in5356 M

aeru

ptions fro

mm

agm

atism

then

volc

anic

s e

rupte

dvo

lcanic

rocks

(5 k

m thic

k)

continent-

Kin

gdom

5861 M

a a

nd

syn-r

ift and p

ost-

onto

sub-a

eri

ally

ocean

Tert

iary

absent 5356 M

ari

ft (

i.e., S

DR

S)

weath

ere

dtr

ansitio

n

Volc

anic

mari

ne s

edim

ents

Pro

vin

ce

3a

1 k

m:Lim

ited d

ata

;B

asalt/r

hyolit

eS

yn-b

asalts

Not know

nN

ot know

nS

DR

S

Sylh

et

Not know

nN

ot know

n

India

ca.1 k

m (

SD

RS

)ca.95118 M

apro

vin

ce?

3b

12 k

m:>

1 k

mB

unbury

Syn-b

asalts

Syn-r

ift and

Not know

nS

DR

S

Walla

by

7.2

7.3

km

/sN

ot know

n

Austr

alia

(SD

RS

) and 1

km

123

132 M

apost-

rift

Pla

teau

Exm

outh

(Walla

by P

late

au)

magm

atism

Pla

teau

4a

1.8

km

:ori

gin

ally

Basalt e

ruptions

Syn-b

asalts (

and

? P

re-r

ift and

? P

ossib

le p

re-r

ift

Yes

Not know

nN

ot know

n

Bra

zil

ca.4.8

km

with

129133 M

apost-

) basalts

syn-r

ift

ele

vation c

a.500 m

.

Para

na

3 k

m lost to

ero

sio

nm

agm

atism

(tim

ing u

ncle

ar)

4b

0.9

km

:ori

gin

ally

Basalt e

ruptions

Syn-b

asalts (

and

? P

re-r

ift and

? P

ossib

le p

re-r

ift

SD

RS

mix

ture

of

Yes

Yes

Nam

ibia

-ca.3.9

km

with

131133 M

apost-

) basalts

syn-r

ift

ele

vation c

a.500 m

.vo

lcanic

s a

nd

Ete

ndeka

3 k

m lost to

ero

sio

nm

agm

atism

(tim

ing u

ncle

ar)

sedim

enta

ry r

ocks

(Corn

er

et al.,

this

volu

me)

51 k

m198201 M

aN

o s

ilicic

volc

anic

Post-

rift m

agm

a-

?? 0

.92.6

km

SD

RS

volc

anic

sYes,

No, norm

al

Centr

al

rocks

tism

(S

.E.U

SA

)syn-r

ifting o

rw

ith s

om

e inte

r-sig

nific

ant

oceanic

cru

st

Atlantic

and s

yn-r

ift

pre

-volc

anic

bedded s

edim

ents

igneous

(78 k

m)

Magm

atic

magm

atism

(N

.E.

intr

usio

ns

Pro

vin

ce

US

A a

nd A

fric

a)

Note

:LIP

, la

rge igneous p

rovin

ce;S

DR

S, seaw

ard

dip

pin

g r

eflecto

r seri

es;H

VLC

, hig

h v

elo

city low

er

cru

st.

Exam

ple

s o

f sourc

e r

efe

rences for

specific

volc

anic

rifte

d m

arg

ins:E

thio

pia

and Y

em

en:B

erc

khem

er

et al.

(1975);

Davis

on e

t al.

(19

94);

Baker

et al.

(1996a,b

);M

enzie

s e

t al.

(1997a,b

);E

glo

ff e

t al.

(1997);

AlS

ubbary

et al.

(1998);

Georg

e e

t al.

(1998);

Hoffm

ann e

t al.

(1997);

Baker

et al.

(2000);

Ebin

ger

and C

asey (

2001);

Ukstins e

t al.

(2002);

Baker

et al.

(this

volu

me).

Gre

enla

nd a

nd U

K:R

obert

s e

t al.

(1979);

Mutter

et al.

(1982);

White e

t al.

(1987);

White a

nd M

acK

enzie

(1989);

Bro

die

and W

hite (

1994);

Saunders

et al.

(1997);

Lars

en a

nd S

aunders

(1998);

Jolle

y (

1997);

Kore

naga e

t al.

(2000);

Pla

nke e

t al.

(2000);

Kla

usen e

t al.

(this

volu

me).

India

and A

ustr

alia

:Von S

tackle

berg

et al.

(1980),

Von R

ad a

nd T

huro

w (

1992);

Sto

rey e

t al.

(1992);

Colw

ell

et al.

(1994);

Exon a

nd C

olw

ell

(1994);

Miln

er

et al.

(1995);

Fre

y e

t al.

(1996);

Kent et al.

(1997).

Para

na a

nd E

tendeka:H

aw

kesw

ort

h e

t al.

(1992);

Renne e

t al.

(1992);

Galla

gher

et al.

(1994);

Turn

er

et al.

(1994);

Renne e

t al.

(1996a, 1996b);

Gla

dczenko e

t al.

(1997);

Peate

(1997);

Cle

mson e

t al.

(1999);

Davis

on (

1999);

Jerr

am

et al.

(1999);

Ste

wart

et al.

(1996);

Hin

z e

t al.

(1999)

and r

efs

there

in;B

auer

et al.

(2000);

Corn

er

et al.

(this

volu

me);

Mohri

ak e

t al.

(this

volu

me);

Tru

mbull

et al.

(this

volu

me);

Watk

eys e

t al.

(this

volu

me);

Centr

al A

tlantic M

agm

atic P

rovin

ce:M

cB

ride (

1991);

Holb

rook a

nd K

ele

men (

1993);

McH

one (

1996);

Liz

arr

ald

e a

nd H

olb

rook (

1997);

Withja

ck e

t al.

(1998);

Ham

es e

t

al.

(2000);

Benson (

2001);

McH

one a

nd P

uffer

(2001);

Schlis

che e

t al.

(2001).

LIP continental basaltic and silicic flood volcanism:

Shallow and deep sources

The birth of volcanic rifted margins (Table 1) is associatedwith the subaerial eruption of basaltic rocks and the minor erup-tion of submarine pillow lavas (e.g., Jolley 1997; Planke et al.,2000) (Fig. 2). Whereas basaltic volcanism normally dominatedthe evolution of the LIP, silicic volcanism may have contributedsignicantly to the total volume of the volcanic pile (e.g., Peate,1997; Bryan et al., this volume; Jerram, this volume). LIPs,which characterize all volcanic rifted margins, are rarely thickerthan 2 km (Table 1) because they represent the erosional rem-nants of earlier sequences estimated to have been as much as 23times as thick at the time of eruption (Table 1) (Cox, 1980;Mahoney and Cofn, 1997). These estimates of the originalerupted thickness on the continental margin take into accountthe amount of subaerial volcanism that has been eroded by syn-rift or postrift processes. However, the erupted thickness differsfrom the actual melt thickness produced during volcanic riftedmargin formation, which must include igneous intrusives addedto the continental crust as dike-sill complexes and plutonic cen-ters. Today these may be evident as unroofed magma chambersextending the length of rifted margins (e.g., Namibia, Scotland),exposed dike swarms (e.g., Saudi Arabia), or overthickenedHVLC. HVLC is never exposed at the surface, but is frequentlyreported from seismic data across volcanic rifted margins.

Many geochronological methods have been applied to vol-canic rifted margins (e.g., Rb-Sr, K-Ar, Ar-Ar), but major ad-vances in argon-argon dating using K-rich phenocryst phases(e.g., sanidine, amphiboles) and lasers have led to an improvedunderstanding of the genesis of silicic and basaltic volcanic rocksin volcanic rifted margins (Renne et al., 1992, 1996a, 1996b;Turner et al., 1994; Hames et al., 2000; Ukstins et al., 2002; Mig-gins et al., this volume). There is considerable debate, however,about the age of individual provinces (see Peate, 1997, for re-view). In the majority of volcanic rifted margins, dating indicatesthat the main pulse (i.e., 70%80%) of subaerial continental mar-gin volcanism, both basaltic and silicic, occurred over a rela-tively short period of time ranging from 1 to 4 m.y. (Table 1).

In some volcanic rifted margins, basaltic volcanic rocks aredominant (e.g., Central Atlantic magmatic province, Greenland),while in others silicic volcanic rocks can constitute a signicantpart of the volcanic stratigraphy (e.g., northeastern Africa, SouthAmerica, Africa) (Table 1). Silicic volcanism can occur earlyduring the main basaltic episode or after the main basaltic erup-tions (e.g., Ethiopia and Parana). Extrusive silicic rocks do notexist in all volcanic rifted margins, but may occur as silicic in-trusives (e.g., Greenland; Table 1). The coexistence of basalticand silicic volcanic rocks or the eventual switch from basaltic tosilicic volcanism reveals the complexity of magmatic processeswithin volcanic rifted margins. Overall the complex relationshipsvary from basalt-dominated volcanic rifted margins, bimodalbasalt-rhyolite volcanic rifted margins, to intermixed basalt-

rhyolite volcanic rifted margins (Table 1). Crustal magma cham-bers play a pivotal role in the formation of silicic magmas, asdoes melting of the lower crust, perhaps fueled by basaltic un-derplating (Cox, 1980, 1988). In Yemen, the silicic volcanismthat postdated basaltic volcanism and lasted >3 m.y. (Baker et al.,1996a) is believed to have originated by processes of assimila-tion and fractional crystallization of mantle-derived melts. Indi-vidual silicic volcanic units can be geochemically linked tonearby intrusive centers, often unroofed as granite-syenite-gab-bro complexes (e.g., UK Atlantic margin, Yemen). These are pre-sumed to have acted as source regions for the silicic volcanicrocks. In contrast, the origin of silicic volcanic rocks fromEtendeka-Paran erupted during the lifetime of the ood basaltprovince (Peate, 1997) may relate more to the formation of large-scale crustal melts. While several igneous complexes in Namibiahave been identied as sources for the volcanic rocks on thebasis of similar ages, the extensive synrift lava cover in manyother examples may hide the identity of associated plutonic com-plexes. In other volcanic rifted margins (e.g., Greenland, Paran,Yemen) the presence of a monotonous basalt stratigraphy onthe rifted margin, and a paucity of plutonic rocks, may indicatethat the plutonic rocks are offshore (e.g., Deccan), or are pre-served on the conjugate margin (e.g., Yemen). It is possiblethat the basalt stratigraphy that dominates the volcanic riftedmargins in Brazil, Ethiopia, and Greenland was inextricablylinked to igneous centers now preserved in their conjugate mar-gins, Namibia, Yemen, and Scotland, respectively.

Along the youthful northeastern African margins, silicicvolcanic rocks were explosively erupted, typically venting102103 km3 of magma (Ukstins et al., 2002). In the DeccanTraps and the North Atlantic Tertiary volcanic province, thepresence of ash layers in the volcanic stratigraphy may indicatesilicic volcanism (Deccan) or alkaline volcanism (Greenland)between periods of basaltic volcanism (e.g., Heister et al., 2001).In other volcanic rifted margins (e.g., Etendeka) (Peate, 1997),individual silicic eruptive units have thicknesses of ca. 100 m,aerial extents >8000 km2, and volumes of 3000 km3. These sili-cic units are comparable in volume to individual mac lava unitsfrom LIPs like the Columbia River. Plinian eruption columns as-sociated with the emplacement of voluminous ignimbrites inthese volcanic rifted margins could have injected large amountsof aerosols into the atmosphere, and so affected global climatemore than basaltic eruptions of similar volume.

Eruption rates in volcanic rifted margins have not been ad-equately dened by volume-time studies of individual eruptiveunits, but as a rst approximation, thickness-time relationshipsreveal a marked decline in eruption rate from the mac to the fel-sic eruptive stages of volcanic rifted margins (e.g., Hawkes-worth et al., 1992; Baker et al., 1996a). This is consistent withthe requirement for longer time periods to allow basaltic mag-mas to pond in shallow magma chambers and to evolve towardsilicic derivatives by a combination of fractionation processesand assimilation of surrounding basement and/or roof rocks.


Continent-ocean transition: HVLC and SDRS

Voluminous subaerial ood volcanism on a continentalmargin lasting for millions of years requires a well-establishedmagma transfer system within the crust and shallow mantle (Fig.2). Cox (1980) rst alluded to the potentially important contri-bution of sill-dike complexes to crustal growth during ood vol-canism. Shallow (i.e., caldera structures) and deeper crustalmagma chambers are a requirement of many models where themineralogy and chemistry of mac magmas indicate fraction-ation at lower crustal pressures and temperatures. The presenceof plagioclase, clinopyroxene, and olivine phenocrysts inbasaltic rocks alludes to fractional crystallization processes inlower crustal magma chambers, and, in many instances the geo-chemistry of these rocks reveals crustal contamination probablyoccurring concomitantly with evolution of the magmas in shal-low or deep crustal chambers (e.g., Cox, 1980; Hooper, 1988).Even more extreme fractionation processes are apparent in therhyolites found within volcanic rifted margins. Such rocks con-tain quartz, mica, and amphibole phenocrysts indicative of high-level processes. While some authors argue for an inextricablelink between underplating and basin inversion (Brodie andWhite, 1994), there are few reports of kilometer-scale, under-plated, high-velocity layers spatially limited below many basinsthat could be analogues of the well-documented HVLC at vol-canic rifted margins (Lizarralde and Holbrook, 1997; Korenagaet al., 2000).

Characteristic features of volcanic rifted margins are zonesof HVLC (Fig. 2) between stretched continental crust and normalthickness oceanic crust (e.g., Kelemen and Holbrook, 1995;Boutilier and Keen, 1999; Korenaga et al., 2000; Benson, 2001;Trumbull et al., this volume). Most likely the HVLC was em-placed during the breakup stage or, if it was a synrift feature, wasassociated with mantle upwelling (e.g., Kelemen and Holbrook,1995; Boutilier and Keen, 1999). In southeast Greenland crustalthicknesses, at equivalent positions on the continental margin,vary from 3040 km thick close to the thermal anomaly (i.e., trackof Iceland hotspot) to 18 km 5001000 km from the anomaly(Korenaga et al., 2000) (Fig. 2). In some volcanic rifted margins,the continent-ocean transition can be abrupt (e.g., Namibia) withentirely new HVLC formed seaward of almost unchanged, per-haps slightly thinned, continental crust. In this case the genera-tion of additional igneous material may have more to do with ex-tension and decompression melting than plumes and/or hotspots.

Current models for volcanic rifted margins are largelybased on the results of geophysical surveying and scienticdrilling in the northeastern Atlantic although few deep wells areavailable to calibrate interpretations (e.g., Korenaga et al.,2000). Scientic drilling in the northeastern Atlantic and indus-try drilling off Namibia (Kudu Field) show that lavas wereerupted subaerially (e.g., Mutter et al., 1982; Clemson et al.,1999). SDRS, rst recognized along the North Atlantic margin,mark the synrift stage in continental breakup and as such are

characteristic of volcanic rifted margins (Roberts et al., 1979;Mutter et al., 1982; White et al., 1987; Larsen and Jakobsdottir,1988; Korenaga et al., 2000; Benson, 2001). Volcanic rifted mar-gins have thick sequences of seaward-dipping volcanic-sedi-mentary strata above, or seaward of, the region of HVLC, andextending landward to the ocean-continent transition zone (e.g.,Mutter et al., 1982; Clemson et al., 1999). Reector packageswithin these SDRS diverge downward and dip oceanward 20 ormore (Fig. 2). Planke et al. (2000) divided the SDRS into in-ner and outer packages (Fig. 2) on the basis of studies of theNorth Atlantic margins (Fig. 1). The inner SDRS were subaeri-ally emplaced ows, the geometry of which was affected bybasin architecture. They proposed that this phase of volcanismoccurred during subaerial seaoor spreading or syntectonicinlling of rift basins. The outer SDRS are believed to representsheet ows in marine basins, and have similarities to subaerialows. Submarine eruptions (i.e., pillowed ows and hyalo-clastites) characterize this developmental stage. SDRS are syn-rift phenomena and are distinct from ood basalts; they straddlethe continent-ocean boundary and can include subaerial andsubmarine volcanic and sedimentary rock types.

On the Namibian margin, modeling of magnetic data fromseismic proles suggests that the SDRS is a mixture of volcanicand sedimentary rocks. Presumably some portion of the SDRSmust comprise sedimentary rocks, given that the volcanicstratigraphy on the uplifted margin can be reduced in thicknessduring synrift erosional processes (Gallagher et al., 1994). How-ever, whether these sediments are argillaceous or arenaceous de-pends on the nature of the material removed from the margin(e.g., metamorphic, sedimentary, or igneous rocks). On the Nor-wegian volcanic rifted margin, seismic sections have been in-terpreted as representing a transition from subaerial to subma-rine volcanic deposits that comprise lavas and volcaniclasticsedimentary rocks (Planke et al., 2000). If we take the Yemenmargin as an indication of what might constitute seaward-dip-ping reector series, it is clear that a signicant proportion (atleast 50%) of these features must be sedimentary in nature. Asediment-budget analysis of the Red Sea margin (Davison et al.,1994) in Yemen indicated that several kilometers of basalticand/or silicic volcanic rock were removed from the volcanicrifted margin during classic synrift extension. This erosionalperiod would have contributed to the SDRS constructed on thestretched continental crust and embryonic oceanic crust.

Several volcanic rifted margins show an abrupt terminationof the SDRS against a high-velocity structural high, which maybe a late synrift intrusion (e.g., Planke et al., 2000), a fault, or anabandoned spreading ridge marking the ocean-continent bound-ary (e.g., Korenaga et al., 2000). Ebinger and Casey (2001) pro-vided a mechanism for synrift emplacement of some SDRS viathe development of high-strain neovolcanic zones and the aban-donment of crustal detachments. Formation of SDRS on vol-canic rifted margins is synchronous with the prerift to synrifttransition on the continental margin. SDRS typically postdate


ood volcanism on the rifted margin, and their formation maybe synchronous with a hiatus in magmatism, a change in mag-matic source area, and a peak in denudation. Because this is asituation that would not be associated with the generation ofmelt, it is likely that strain localization and focused extensionaccelerated melt generation. Although SDRS may predateocean crust formation at a mid-ocean ridge, they are transi-tional between rifted continental margin processes and oceanridge processes (Fig. 2). The continent-ocean transition isdifcult to determine, and therefore considerable controversysurrounds the nature of the crust beneath many SDRS. Thepetrology and geochemistry of both SDRS and the HVLC holda vital clue to a major change in the source of magmas, fromone that fed a LIP to one that produced oceanic crust. SDRSand HVLC are two principal diagnostics of volcanic riftedmargins (Fig. 2).

Breakup extension: Pre-LIP, syn-LIP, or post-LIP?

The relationships between the timing of LIP formation andrifting leading to ocean-oor formation are complex. This mayin part be explained by the fact that some volcanic rifted mar-gins are proximal, others distal, to plume heads and/or stems, soit is unlikely that volcanic rifted margins will show the same re-lationships. It is also complicated by the possibility that magmasources for volcanic rifted margins may reside either in the deepmantle (i.e., plumes) or the shallow mantle (i.e., asthenosphericsmall-scale convection). The temporal relationship betweenmagmatism and extension may differ greatly if, as we believe,in deep-sourced plumes enhanced temperatures triggered meltproduction, whereas asthenospheric melts are decompressionmelts triggered by lithospheric thinning. We envisage plume-de-rived magmatism occurring at any stage in the development ofa rifted continent (prerift, synrift, or postrift), whereas magma-tism derived from the shallow mantle would largely be synriftor postrift.

Another problematic aspect of understanding the relation-ship between extension and magmatism is dening the timing ofrifting and/or extension. Extension may be fault controlled orvia dike injection (Klausen and Larsen, this volume), and maybe identied as the appearance of the rst fault, the rst volcanicrock, or the rst depocenter. Is the onset of extension the timingof the initiation of continental extension, or is breakup markedby the formation of seaoor sensu stricto? Tens of millions ofyears can pass between the initiation of LIP formation (prerift)and the generation of seaoor, so it is important to understandabsolute and relative timing of the geological processes leadingto the formation of new seaoor. Any generalization about theapparent synchroneity of magmatism, extension, and uplift ig-nores the reality that, with the technology available, we can re-solve the relative and absolute timing of these processes and sobetter understand rift processes.

In Figure 3 the relationship between ood volcanism (i.e.,LIP formation) and the formation of oceanic crust is summa-

rized for many of the volcanic rifted margins that formed in thepast 200 m.y. (see also Courtillot et al., 1999). The age of theoldest oceanic crust adjacent to the volcanic rifted margin inquestion can be used as a minimum age of seaoor spreadingbecause it is conceivable that this is not the oldest ocean oor,but merely the oldest seaoor for which samples exist (Fig. 3).The age of oceanic crust can be compared with the age of oodvolcanism on the volcanic rifted margin to better understand therelationship between extension and magmatism.

In Ethiopia-Yemen, magmatism is dated by Ar-Ar methodsas 3126 Ma (Baker et al., 1996a; Hoffman et al., 1997; Ukstinset al., 2002). Extension (leading to the formation of dominofault-block terranes) is dened by Ar-Ar and ssion-track datingof hanging-wall and footwall lithologies (Menzies et al., 2001).Extension in Yemen (i.e., southern Red Sea margin) began in thelate Oligocene (ca. 26 Ma), coincident with a marked hiatus inextrusive activity and signicant tectonic erosion and/or crustalcooling dated by ssion-track methods and validated by Ar-Ardating of unconformities as 1925 Ma (Baker et al., 1996a;Menzies et al., 1997a). On the conjugate margin in Ethiopia,extension occurred along the length of the western escarpment25 Ma, indicating that rifting occurred after the onset of oodbasaltic volcanism ca. 31 Ma (Ukstins et al., 2002). Volcanicrocks were erupted from isolated centers located along the west-ern escarpment in Ethiopia (Kenea et al., 2001; Ukstins et al.,2002). We conclude that much of the Ethiopian-Yemeni oodvolcanism was prerift in character. While the timing will not bethe same for all volcanic rifted margins, the southern Red Sea isan illustration of how breakup and the continent-ocean transi-tion can be protracted.

In the case of the North Atlantic (Greenland-UK) (Fig. 1),LIP formation lasted from 61 to 53 Ma (e.g., Eldholm and Grue,1994; Saunders et al., 1997) and the oldest oceanic crust indi-cates that extension must have taken place before 52 Ma (Fig.3). From this it appears that volcanism in the North Atlanticstraddled breakup with a prerift (LIP) and a synrift stage (SDRS)(Larsen and Saunders, 1998). Such a protracted period of vol-canism may explain the attenuated, heavily intruded nature ofthe broad continent-ocean transition.

The details of the timing are less well known for Australia-India (Fig. 3). Volcanism on the Indian and Australian marginsoccurred between 100 and 130 Ma (e.g., Kent et al., 1997), andbreakup between Australia, India, and Antarctica was 125133Ma (Fig. 3). It appears that volcanism on the rifted margin wassynchronous with continental breakup, but that volcanism con-tinued (sporadically?) during formation of oceanic crust.

In the Paran-Etendeka volcanic rifted margins (Fig. 1),oceanic crust located off Africa is slightly older than that knownoff South America. The age of the oceanic crust indicates thatextension occurred ca. 135 Ma, overlapping with the Paran-Etendeka LIP (Peate, 1997). Because the main pulse of basalticmagmatism occurred ca. 130133 Ma, it can be inferred that theLIP was largely prerift to synrift. A prerift stage is supported bythe fact that the main volcanic units can be traced, and the vol-


canic stratigraphies matched, from the Etendeka across the At-lantic Ocean to the Paran of Brazil (e.g., Milner et al., 1995;Mohriak et al., this volume). Synrift magmatism is supportedby offshore valley systems that appear to be lled with extru-sive lavas with later deformation and faulting-controlled em-placement of the volcanic units (cf. Clemson et al., 1999).Alternatively, both these observations could be explained by asynrift model for the magmatic activity. Initial pulses of mag-matism would ll topographic lows, as described by Clemsonet al. (1999), and further synrift activity would mantle the lledtopography such that units were traceable from South Americato Africa, as reported by Milner et al. (1995).

The relationships for the Central Atlantic magmatic province(Fig. 3) appear more complex, probably because of the size ofthe province and the extent to which it has been eroded. Conti-nental magmatism has been dated as 198201 Ma (Hames et al.,2000). However, along the eastern margin of North America therelationship between magmatism and tectonics is variable (J.McHone, 2001, personal commun.). In southeastern NorthAmerica, volcanic rocks of the Central Atlantic magmaticprovince appear to postdate both the cessation of rifting by ca.10 Ma, and uplift and/or erosion. This should be contrasted withCentral Atlantic magmatic province magmatism in northeasternNorth America and northwestern Africa, where magmatism issynrift and rifting continued for ~25 m.y. after magmatism fol-lowed by Middle to Late Jurassic uplift (J. McHone 2001, per-sonal commun.). SDRS from offshore northeastern UnitedStates are thought to have been emplaced ca. 175 Ma (Withjacket al., 1998; Benson, 2002; Schlische et al., 2002), and oodvolcanism appears to be synrift or postrift. This contrasts withthe North Atlantic margins (Greenland, UK) where a signicantprerift ood volcanic stage is evident. However, there may be abias in the rock record. In the Central Atlantic magmaticprovince, onshore intrusive rocks are used to dene the timingof magmatism on the rifted margin. However, in deeply erodedvolcanic rifted margins, like the Central Atlantic magmaticprovince, these hypabyssal and/or plutonic rocks may bias thedating toward the synrift stage. We use the Yemen volcanic riftedmargin as an illustration of how hypabyssal and/or plutonicrocks may be largely synrift in age, despite a prerift history of45 m.y. of ood basalt volcanism unrepresented in these ex-posed hypabyssal and/or plutonic rocks. In Yemen the originalsubaerial volcanic stratigraphy has an age of 3126 Ma, and isknown to be prerift (Baker et al., 1996a; Menzies et al., 1997a,1997b). Hypabyssal and plutonic rocks underlying or intrudingthe volcanic rifted margin have ages that are primarily youngerthan 25 Ma (Chazot et al., 1998, and references therein) and sointrusive activity, as exposed, is largely synrift. It appears thatpeak extension (and erosion 1926 Ma) was associated with apossible extrusive hiatus, but with signicant intrusive activityexemplied by the dike swarms and granite-gabbro-syenite lac-coliths. This synrift intrusive stage is conrmed by

lengthy inland course because of the uplifted rift margins. Someof the best examples are the Rio de la PlataParana (Brazil),which has its headwaters 31 Ma) close to thepresent location of a 4-km-high mountain range. Paleocurrentinformation and the maturity of the prevolcanic sediments inYemen require a hinterland on what is now the opposite sideof the rift in the Danakil horst, Eritrea (AlSubbary et al.,1998). The prevolcanic sedimentary rocks imply that the con-tinental masses were close to sea level in the southern RedSea and, by inference, Eritrea (AlSubbary et al., 1998). Thepredominance of subaerial volcanic rock units (rather thansubmarine ows or hyaloclastites) also indicates a subaerialcontinental environment at 31 Ma; some rift-related upliftmust have occurred before that time. Possibly the initiation ofuplift is recorded in changes to the orientation of the pale-oshoreline, along with a shallow marine to continental transi-tion that occurred prior to volcanism. These changes indicatethat uplift of that surface (ca. 31 Ma) was tens to hundreds ofmeters. However, the exact age of these sediments relative tothe period of formation of the volcanic rifted margin is un-known. In the unlikely event that these sedimentary rocks areconsiderably older than the volcanic rifted margin, the pale-oenvironmental changes would not relate to the evolution ofthe volcanic rifted margin.

Fission-track ages date crustal cooling and hence rapid tec-tonic denudation as having occurred between 19 and 26 Ma(Menzies et al., 1997a) on the Red Sea margin. We presume that

Oligocene-Miocene denudation required greater topographythan the paleoshoreline inferred to exist at 31 Ma, hence suchtopography was generated, and the period of uplift and exhu-mation is bracketed, in the late Oligocene (2631 Ma). Inde-pendent verication of this period of denudation is to be foundin unconformities in the volcanic stratigraphy that formed be-tween 19 and 26 Ma (Baker et al., 1996a). In contrast to thelargely synvolcanic uplift in Yemen, uplift on a scale of hun-dreds of meters is believed to have preceded volcanism in theNorth Atlantic province (e.g., Larsen and Saunders, 1998). Inwestern Greenland major unconformities beneath the volcanicrocks are associated with uvial peneplanation and valley inci-sion, indicating a period of prevolcanic uplift and erosion.However, in eastern Greenland during the same time the land-scape was close to sea level and, in northwest Scotland-Faeroes,the prevolcanic landscape was a low-relief, vegetated landsurface (Jolley, 1997). Furthermore, in northwest Scotland sub-aerially weathered marine sediments (chalk) underlie the low-ermost ood volcanic rocks (G. Fitton, 2001, personal com-mun.). The prerift North Atlantic volcanic rifted margin (Brodieand White, 1994) could be classied as a low-relief land sur-face with some incised river systems. The proximity of that landsurface to sea level is revealed by studies in Norway (Planke etal., 2000), where the margin is believed to have developed inthe continental to oceanic transition.

In Namibia and Brazil (Etendeka-Parana) the basal vol-canic rocks overlie and are interbedded with continental eoliansandstones, which occasionally overlie uvial deposits. A largeeolian erg system is reported intercalated with the lowermostood basalts (Jerram et al., 1999), but how far above sea level itwas formed is not known. On this volcanic rifted margin upliftand doming prior to rifting could be argued for due to the lackof Upper Karoo sediments (Clemson et al., 1999). This may beconsistent with conclusions based on ssion-track data that ar-gue for a prerift elevation of ~500 m in southeastern Brazil (Gal-lagher et al., 1994).

The degree of preservation of volcanic rifted margins is in-extricably linked to climate, elevation, and the amount and/orrate of erosion. The youngest volcanic rifted margins survive as34-km-high mountain ranges in the desert climate of north-eastern Africa, and Cretaceous-Tertiary volcanic rifted marginsare characterized by ~57-km-thick volcanic sections in themountain ranges of subpolar Greenland, deeply eroded riftmountains in the west maritime climate of the UK, and majorscarp retreat in tropical India and Brazil. The western Ghat es-carpment (Deccan) is believed to have an erosional, rather thana tectonic origin, and scarp retreat is believed to be the major de-terminant of landscape with the original continental margin ~75km west of its present location. Erosion over 200 m.y. has re-duced the subaerial portion of the Central Atlantic magmaticprovince in the eastern United States and western Africa to adike swarm. However, submarine equivalents to these marginshave survived offshore as SDRS. Just as in the relationship be-tween magmatism and rifting, we have a complex history of up-


lift and denudation operating on a variety of prevolcanic land-scapes.

Volcanic rifted margins: Summary

Volcanic rifted margins evolved in response to local ther-motectonic conditions, and consequently marked differencescan be found in the temporal and spatial relationships betweentectonics, magmatism, uplift, and erosion. Of all passive conti-nental margins around the world, 90% are volcanic rifted mar-gins to varying degree, the exceptions being continental marginsin eastern China, Iberia, the northern Red Sea, South Australia,the Newfoundland BasinLabrador Sea, and possibly the Gulfof California. Although the Arctic and Antarctic margins havelargely unknown status, parts of the Antarctic margin are be-lieved to be volcanic (Fig. 1).

A considerable variation exists in volcanic rifted margins.The prevolcanic environment can vary from shallow marine-continental (e.g., North Atlantic), uvial-continental (e.g.,Yemen-Ethiopia) to eolian continental (e.g., Etendeka). Floodvolcanism can be thick (e.g., 7 km, Greenland) or relatively thin(e.g., 1.5 km, Deccan). Volcanism can be represented by pre-dominantly basaltic volcanic rocks at the base and by mainlysilicic volcanic rocks at the top (e.g., Yemen), or silicic volcanicrocks may be found throughout the volcanic stratigraphy (e.g.,Ethiopia) or be essentially absent (e.g., Deccan). Processes re-lated to volcanic rifted margin can vary in time and space. Mag-matism can predate breakup extension by several million years(e.g., Yemen-Ethiopia), magmatism and breakup can be syn-chronous (e.g., GreenlandNorth Atlantic Tertiary volcanicprovince), or magmatism can postdate breakup by several mil-lion years (e.g., Australia-India). Magmatism in volcanic riftedmargins may have originated by decompression melting associ-ated with lithospheric thinning and/or upwelling of thermalanomalies modied by melting of lithospheric rocks. In mostvolcanic rifted margins prevolcanic uplift can vary from tens ofmeters (Yemen) to hundreds of meters (North and South At-lantic), but it appears that kilometer-scale prevolcanic uplift isnot as widespread as some models have predicted.

The magmatic and structural evolution of individual vol-canic rifted margins is complex and may not t simple models.This may be due to the geology, age, and thickness of the pre-rift lithosphere and proximity to plume heads, which are poten-tially variable in temperature, longevity, and dimensions. Thereappears to be a continuous gradation from volcanic rifted mar-gins to nonvolcanic rifted margins; a possible continuum is ev-ident in the southern Red Sea. Much remains to be learned aboutthe extent to which plumes drive, or are focused by, lithosphericextension and the exact geophysical and geochemical nature ofplumes. Perhaps we can better understand the process of for-mation of volcanic rifted margins by comparing and contrastingtheir geological characteristics with nonvolcanic rifted marginswhere continental rifting occurs without thermally enhancedmantle (e.g., Newfoundland, Iberia), or the formation of in-

traplate large igneous provinces in ocean basins (e.g., Ontong-Java oceanic plateau) and continents (e.g., Siberian oodbasalts) where widespread rifting is absent. Clearly a single ac-tive rifting model cannot explain the formation of all volcanicrifted margins around the world.

Volcanic rifted margins: Characteristics

The following characteristics are sufciently common involcanic rifted margins to be diagnostic.

1. Flood volcanism may have reached 47 km thicknessprior to erosion, which has reduced several margins to thick-nesses of 12 km.

2. Basaltic and silicic volcanic rocks are erupted subaeri-ally. Of the exposed subaerial basaltic rocks, 70%80% oc-curred in

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MANUSCRIPT ACCEPTED BY THE SOCIETY OCTOBER 15, 2001


Printed in the U.S.A.

Geological Society of AmericaSpecial Paper 362

2002

Crust and upper mantle structure in East Africa:

Implications for the origin of Cenozoic rifting and volcanism

and the formation of magmatic rifted margins

Andrew A. Nyblade

Department of Geosciences, Penn State University, University Park,

Pennsylvania 16802, USA

ABSTRACT

Crust and upper mantle structure in East Africa, together with the tectonic his-

tory of the region, is used to evaluate geodynamic processes commonly associated with

the formation of magmatic rifted margins. Cenozoic rifting and volcanism in East

Africa represent the earliest stage in the development of a rifted continental margin,

and East Africa is one of the few places where geodynamic processes may be active

that could lead to the development of a magmatic rifted margin. The Precambrian tec-

tonic framework of East Africa is characterized by an Archean craton (Tanzania Cra-

ton) surrounded by Proterozoic mobile belts. An extensive nonmagmatic rift system

associated with the separation of Madagascar from Africa developed in the mobile

belts during the Permian-Cretaceous. In the Cenozoic, two rift branches (Western Rift

and Eastern Rift) formed in the mobile belts. Volcanism is present in both branches,

but most of it is concentrated in the Eastern Rift. Crustal structure away from the

Cenozoic rifts is typical for Precambrian crust. Similarly, uppermost mantle structure

across East Africa does not appear to have been altered, except beneath the rift val-

leys. Deeper in the upper mantle a thick (~200 km) lithospheric keel is found under

the Tanzania Craton, and a broad (200400 km wide) thermal anomaly extending to

a depth of at least 400 km is beneath the Eastern Rift. Several models cannot fully ac-

count for the tectonic history of East Africa and/or the structure of the crust and up-

per mantle, including edge ow in the convecting mantle around the keels of Archean

cratons, broad (>500 km wide) thermal upwellings originating in the lower mantle, arelatively stationary plume head, and a plume head that ows outward along topog-

raphy on the underside of the lithosphere. Consequently, these models are not strong

candidates for the origin of volcanism and rifting in East Africa. Models with two or

more plume heads, however, can explain the relevant observations, and therefore a

multiple plume head explanation for the rifting and volcanism in East Africa is fa-

vored. These ndings further call into question the viability of nonplume models for

the formation of magmatic rifted margins, particularly models invoking edge ow in

the convecting mantle around cratonic keels.

15

Nyblade, A.A., 2002, Crust and upper mantle structure in East Africa: Implications for the origin of Cenozoic rifting and volcanism and the formation of magmaticrifted margins, in Menzies, M.A., Klemperer, S.L., Ebinger, C.J., and Baker, J., eds., Volcanic Rifted Margins: Boulder, Colorado, Geological Society of AmericaSpecial Paper 362, p. 1526.

INTRODUCTION

The origin of magmatic rifted continental margins is highlycontended. Mantle plume models have long been advocated(e.g., White and McKenzie, 1989; Richards et al., 1989; Hill etal., 1992), but nonplume models invoking passive rifting abovewarmer than normal regions of the sublithospheric mantle havealso been argued for vigorously (e.g., Anderson, 1994; King andAnderson, 1995, 1998).

Evaluating models for the formation of magmatic riftedmargins is difcult because there are few, if any, magmaticrifted margins developing today where crust and upper mantlestructure can be investigated in an active tectonic setting. Ceno-zoic rifting and volcanism in East Africa represent the earlieststage in the development of a rifted continental margin, andEast Africa is one of the few places where geodynamicprocesses that could lead to the development of a magmaticrifted margin may be active. Therefore, in this chapter crust andupper mantle structure in East Africa, together with the tectonichistory of the region, is used to evaluate geodynamic processescommonly associated with the formation of magmatic riftedmargins.

Specically, two questions concerning the formation ofmagmatic rifted margins are addressed. (1) Is ow in the con-vecting mantle around the edges of Archean cratons a viablemechanism for generating magmatic rifted margins? Manymagmatic rifted margins are similar to the East African rift sys-tem in that they developed adjacent to an Archean craton (e.g.,DeccanChagosLacadive Ridge, North Atlantic volcanicprovince, Karoo; Cofn and Eldholm, 1992). Even if the EastAfrican rift system never develops into a passive continentalmargin, its special tectonic setting nonetheless affords an excel-lent opportunity to evaluate models for magmatic rifted marginsinvoking convective ow in the mantle around the edges of cra-tons. (2) Can plume models adequately explain crust and uppermantle structures found beneath a possibly embryonic mag-matic rift? Plume models for the formation of magmatic riftedmargins are favored by many. Several different plume modelshave been proposed for East Africa, and by evaluating thesemodels it might be possible to gain further insights into the na-ture of mantle plumes and their role in the formation of mag-matic rifted margins.

CRUST AND UPPER MANTLE STRUCTURE

IN EAST AFRICA

Geology

The Precambrian basement of East Africa consists of theArchean Tanzania Craton, which is in the center of the region,and a number of Early to Late Proterozoic mobile belts sur-rounding it (Fig. 1). The Cenozoic rift valleys have developedalmost exclusively within the mobile belts, largely skirting thecratonic nucleus.

The Tanzania Craton consists mainly of granites, gneisses,and amphibolites; some greenstone belts are in the region northof 4.5S. The youngest dates for the craton (ca. 2500 Ma) comefrom granitic rocks near its eastern margin (Cahen et al., 1984).To the east of the Tanzania Craton is the Mozambique Belt,which has mainly north- to south-striking structures formed bymultiple collisional events dated between 1200 Ma and 450 Ma(Cahen et al., 1984; Shackleton, 1986; Key et al., 1989). Nd andSr model ages and Pb isotopes indicate that much of the Mozam-bique Belt crust initially formed in the Late Archean and EarlyProterozoic (Moller et al., 1998). The Tanzania Craton is bor-dered to the southeast and southwest by the Early ProterozoicUsagaran and Ubendian Belts, respectively (Lenoir et al., 1994;Theunissen et al., 1996). The northern part of the Ubendian Beltis truncated west of the Tanzania Craton by the northeast-trend-ing Late Proterozoic Kibaran Belt (Cahen et al., 1984). To thenorth of the craton is the Late Proterozoic Ruwenzori Belt (Ca-hen et al., 1984).

Extensive late Paleozoic to Mesozoic rifts can be foundthroughout parts of East Africa and are generally oriented north-east-southwest or northwest-southeast. The present outlines ofthese rifts, now reduced from their original sizes by uplift anderosion, exceed the size of the Cenozoic rifts (Fig. 2). Twoepisodes of rifting are well documented (Kreuser, 1995). Karoorifting began during the Early Permian with the initial separa-tion of Madagascar from Africa and terminated in the EarlyJurassic, leaving a rift system that extended from the coastal re-gion of Kenya southward through Tanzania, Mozambique,Malawi, Zambia, and Zimbabwe. Shortly afterward, a newphase of rifting commenced in the Early Jurassic along the en-tire coast of East Africa and terminated at 165 Ma with the nalseparation of Madagascar from the African continent.

There is little volcanism associated with the Paleozoic andMesozoic rifts in East Africa (and also Madagascar). The onlyknown Paleozoic or Mesozoic volcanic units are some Late Cre-taceous sills in the Anza graben (Bosworth and Morley, 1994) anda few Early Jurassic sills in the Luwegu trough (Fig. 2) (Hankel,1987). Locally, the sills in the Luwegu trough attain thicknessesof ~120 m (Spence, 1957) and are similar in age and compositionto the extensive Karoo ood basalts of southern Africa (Kreuser,1995; Hankel, 1987). In addition to these sills, a number of vol-canic plugs and dikes, mostly carbonatitic, have been found insome of the Karoo basins, but they are predominantly early Ter-tiary in age (Hankel, 1987; Bosworth and Morley, 1994).

The Cenozoic rift system in East Africa consists of twobranches, the Western Rift and the Eastern Rift (Fig. 1). Exten-sion within the Eastern Rift in Kenya has led to the formation ofa narrow (5080 km wide) rift graben commonly referred to asthe Kenya or Gregory rift. At the terminus of the Eastern Rift innortheastern Tanzania, the graben structures found in Kenyagive way to a much wider zone (~300 km) of block faulting(Dawson, 1992; Ebinger et al., 1997; Foster et al., 1997). TheWestern Rift is composed of numerous en echelon fault-bounded basins (Ebinger, 1989). Many of these rift basins to the

16 A.A. Nyblade

south and southwest of the Tanzania Craton developed within oradjacent to Karoo rifts, and in some cases Karoo-aged faultsmay have been reactivated.

The timing of Cenozoic volcanism in East Africa is fairlywell dened. George et al. (1998) provided a summary of mag-matism in the Eastern Rift showing a clear southward progres-sion in onset age. Magmatism commenced at 4045 Ma insouthern Ethiopia, 3035 Ma in northern Kenya, 15 Ma in cen-tral Kenya, and 8 Ma in northern Tanzania. In contrast to theEastern Rift, there are only a few volcanic centers in the West-ern Rift (Fig. 1). Magmatism began ca. 12 Ma in the northerneruptive centers around Lake Kivu and at 8 Ma in the RungweProvince in southern Tanzania (Ebinger et al., 1989; Pasteels etal., 1989; Kampunzu et al., 1998). Rift faulting in the Eastern

and Western Rifts probably commenced about the same time asthe magmatism (Ebinger, 1989).

Crustal structure

Until recently, investigations of crustal structure in EastAfrica focused primarily on the Eastern and Western Rifts. Earlystudies used seismic refraction data and observations from tele-seismic and regional earthquakes to examine crustal structure(Bonjer et al., 1970; Grifths et al., 1971; Long et al., 1972;Mueller and Bonjer, 1973; Bram and Schmeling, 1975; Noletand Mueller, 1982; Hebert and Langston, 1985), yielding esti-mates of Moho depths of 4048 km beneath unrifted crust, andof 2032 km under the rift valleys. Work on crustal structure in

Crust and upper mantle structure in East Africa 17

Figure 1. Schematic map of EastAfrica showing political boundaries,large rift lakes, Precambrian terranes,Cenozoic rift system, and Cenozoicvolcanic provinces.

and around the Kenya Rift was undertaken by the Kenya Rift In-ternational Seismic Project (KRISP) using seismic refractionproles (Prodehl et al., 1994; Fuchs et al., 1997, and referencestherein). The KRISP group found that along the axis of the riftMoho depth shallows from 35 km in southern Kenya to ~20 kmbeneath northern Kenya. Away from the rift, crustal thicknessesof 3440 km beneath the Tanzania Craton and 3542 km be-neath the Mozambique Belt were obtained.

A more recent investigation of crustal structure in EastAfrica by Last et al. (1997) using teleseismic receiver functionsand Rayleigh wave dispersion measurements has provided ad-ditional information about crustal structure in the mobile beltsand craton. Earthquakes used in this study were recorded by theTanzania Broadband Seismic Experiment (Nyblade et al.,1996). For the Tanzania Craton, Last et al. (1997) obtainedMoho depths of 3742 km, a mean crustal shear velocity (Vs) of3.79 km/s, and Poissons ratios of 0.240.26. For the Mozam-bique Belt, Last et al. (1997) obtained a Vs of 3.74 km/s, Mohodepths between 36 and 39 km, and Poissons ratios between 0.24and 0.27. Results obtained by Last et al. (1997) from the Uben-dian Belt indicate a Vs of ~3.74 km/s and Moho depths between

40 and 45 km. Based on a comparison of these results and theKRISP results to global averages for Precambrian crust, it canbe concluded that crustal structure away from the narrow rift val-leys has not been modied to any signicant extent by the Ceno-zoic tectonism in East Africa.

Upper mantle structure

Uppermost mantle structure beneath the Kenya Rift was in-vestigated by the KRISP group using seismic refraction proles(Prodehl et al., 1994; Fuchs et al., 1997, and references therein).The pattern of Pn velocities beneath and adjacent to the KenyaRift is fairly simple (Fig. 3). Low Pn velocities of 7.57.8 km/sare found under the axis of the rift, and Pn velocities of 8.18.3km/s are found under the unrifted areas of the Mozambique Beltand Tanzania Craton. The transition from low to high Pn veloc-ities is abrupt and coincides with the main rift border faults, in-dicating that thermal modication of uppermost mantle struc-ture is conned in Kenya to the rift proper. Models derived frominverting teleseismic travel time residuals, however, suggest thatdeeper in the upper mantle under Kenya the zone of modied

18 A.A. Nyblade

Figure 2. Schematic map of East Africashowing locations of Permian-Creta-ceous rift basins, main Cenozoic riftfaults, and Archean Tanzania Craton(Permian-Cretaceous rifts taken fromKreuser, 1995; Bosworth and Morley,1994).

lithosphere may extend over a broader region (Achauer et al.,1994; Slack et al., 1994; Green et al., 1991). Nolet and Mueller(1982) examined mantle structure beneath the Western Rift bysimultaneously inverting teleseismic body wave traveltimes andsurface wave phase and group velocities. They found a thin (~20km), high-velocity lid beneath the Western Rift underlain by alow-velocity channel.

Upper mantle structure beneath the craton and rifted mobilebelts in Tanzania has been investigated in a number of studiesusing data from the Tanzania Broadband Seismic Experiment(Nyblade et al., 1996). The P wave traveltimes from regionalearthquakes were inverted for long wavelength (>100 km) Pnvelocity variations beneath Tanzania by Brazier et al. (2000) us-ing a generalized inverse algorithm. Brazier et al. found Pn ve-locities of 8.408.45 km/s beneath the center of the TanzaniaCraton, 8.308.35 km/s beneath the Mozambique Belt where theEastern Rift terminates, and 8.358.40 km/s beneath the West-ern Rift (Fig. 3). These velocities are high for Precambrian litho-sphere and thus indicate that there are no broad (>100 km wide)

thermal anomalies in the uppermost mantle beneath the cratonor rifted mobile belts in Tanzania.

Structure deeper in the upper mantle beneath the Mozam-bique Belt and craton in Tanzania was imaged by inverting rel-ative traveltimes from teleseismic P and S waves for upper man-tle seismic velocity variations (Ritsema et al., 1998). Thepatterns of P- and S-wave velocity variation obtained are simi-lar, and so only the S-wave velocity model is discussed here(Fig. 4A). The model shows higher than average velocities be-neath the Tanzania Craton and predominantly lower than aver-age velocities beneath the rifted mobile belts surrounding thecraton. The low-velocity region under the Eastern Rift extendsvertically to depths greater than 400 km and laterally over a re-gion ~300 km wide. The lithospheric keel beneath the craton, asdened by the relatively fast velocities, extends to a depth of~200 km (the continuation of fast velocities to 300400 kmdepth in Fig. 4A is due to limited vertical resolution). Betweendepths of 200 and 300 km the low-velocity structure associatedwith the rifts begins to extend westward under the fast structure

Crust and upper mantle structure in East Africa 19

Figure 3. Pn velocities across East Africa from Brazier et al. (2000) and KRISP group. Geological features as inFigure 1. Numbers next to KRISP refraction lines give Pn velocities (in km/s).

Figure 4. A: Vertical slice through S wavevelocity model for East Africa from Rit-sema et al. (1998). Geographic location ofvertical slice is shown in B. Uncertaintiesin horizontal and vertical dimensions ofvelocity structures are ~50 and 100 km, re-spectively. Velocity structure above 100km and below 500 km is

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