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Geosphere

doi: 10.1130/GES00805.1 2012;8;1606-1631Geosphere

Ana Krueger, Mike Murphy, Ed Gilbert and Kevin Burke Barreirinhas Basin, BrazilDeposition and deformation in the deepwater sediment of the offshore

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1606 For permission to copy, contact [email protected]© 2012 Geological Society of America

Deposition and deformation in the deepwater sediment of the offshore Barreirinhas Basin, Brazil

Ana Krueger1, Mike Murphy1, Ed Gilbert2, and Kevin Burke1

1Department of Earth and Atmospheric Sciences, University of Houston, Houston, Texas 77204-5007, USA2Consulting Geologist, 22111 Joshua Kendall Lane, Katy, Texas 77449, USA

Geosphere; December 2012; v. 8; no. 6; p. 1606–1631; doi:10.1130/GES00805.1; 18 fi gures; 1 table.Received 25 March 2012 ♦ Revision received 21 August 2012 ♦ Accepted 5 September 2012 ♦ Published online 16 November 2012

ABSTRACT

The Barreirinhas Basin is an ideal loca-tion to study shale-dominated gravity-driven thrusting systems because of the limited areal extent of the deformed areas compared to other areas in the world. Regional seismic refl ection profi les across the Barreirinhas Basin on the Brazilian Equatorial margin show two major deepwater fold and thrust belts linked landward to extensional fault systems. Thrust faults are interpreted to be products of shortening caused by gravity-driven extension on the continental margin that involve rocks of both the shelf and the slope. Results show two main deformation events during the Cretaceous (between ca. 89 and 84 Ma) and several episodes dur-ing the Later Cenozoic (ca. 55–0 Ma). All events were characterized by displacement along a detachment fault linking a landward system of normal faults to a basinward sys-tem of folds and thrust faults. The Creta-ceous deformation involved a <1.5-km-thick sequence deformed in a 30-km-wide set of listric normal faults (an extensional domain) on the outer continental shelf and upper slope. Those normal listric faults merge sea-ward into a bed-parallel detachment surface forming a 30-km-wide translational domain and a 30-km-wide zone of imbricate thrust faults (a compressional domain) at the toe of the slope. The Cenozoic structural sys-tem involves a thick (over 4 km) sedimen-tary sequence of Turonian to Miocene age, which cross-cuts the preexisting Cretaceous deformed sequence. The Cretaceous and Cenozoic deformational events formed two discrete bowl-shaped fault systems that are linked at different stratigraphic levels. Plots of displacement versus time show normal- and thrust-fault movements during the same time intervals, confi rming the link between

extension on the continental shelf and com-pression on the slope. Deformation increased dramatically during the past 10 m.y., with movement on all earlier and some newly formed faults. The increased deformation coincided with tectonic paleogeographic and topographic changes in northern South America during the Late Miocene that led to an increase in sediment supply to the Bar-reirinhas Basin.

INTRODUCTION

Gravity gliding covers a wide range of tempo-ral and spatial scales, from small (up to several square kilometers in area) surfi cial geologically instantaneous slumps or slides, to long-lived giant submarine fold-thrust belts that can cover thousands of square kilometers (Morley et al., 2011). Because deformation rates can control the generation of mass fl ows (i.e., turbidites) and displace large masses of water, understand-ing the mechanisms and kinematics of deforma-tion in these systems can be critical to several fi elds, notably reservoir presence for hydrocar-bon exploration and local sequence stratigraphy (Gilbert et al., 2011).

Small-scale underwater slides are commonly triggered by catastrophic events such as storms, earthquakes, and high rainfall. They contrast with areally extensive gravity gliding in deep-water fold-thrust belts on passive margins that result from long-term sediment loading (Morley et al., 2011). The latter give rise to large-scale compressional provinces or deepwater fold-thrust belts at the toe of the continental slope, linked to shelf extensional provinces by a broad, relatively undeformed zone of lateral transport.

Large-scale gravity gliding systems devel-oped in sediment piles, such as the Niger Delta, can be described in terms of the critical taper angle thrust model (Davis et al., 1983; Dahlen, 1984), where the combination of a thick sedi-

ment pile, steep basinward surface slope, and a gentle landward basal slope commonly gener-ates shear stress to promote basinward sliding up the basal slope. Loss of shear stress at the toe of the slope generates compression and results in a deepwater fold-thrust belt at the toe of the slope. Such detachments usually occur in evap-orites or thick shales (Rowan et al., 2004), and shale detachments are normally associated with documented fl uid overpressure that also reduces the frictional resistance at the base of the sedi-ment pile (Mourgues et al., 2009). Examples include the Amazon fan (Cobbold et al., 2004), Niger delta (Corredor et al., 2005; Billotti and Shaw, 2005; Cobbold et al., 2009), and the Mexican ridges (Weimer and Buffl er, 1992). In order to evaluate the mechanisms that drive large-scale gravity glide systems, it is critical to know the geometry and sequence of faulting and their relationship to sedimentation.

In this paper, we present seismic interpretation across a poorly documented region of the Bra-zilian equatorial margin in order to: (1) assess the geometry and timing of faulting; (2) docu-ment the structural relationship between thrust faults at the toe of the slope and normal faults at the shelf edge; and (3) evaluate their relation-ship to the sedimentation history.

REGIONAL TECTONICS AND BASIN STRATIGRAPHY

The Barreirinhas Basin is one of a set of basins on the Equatorial Brazilian margin, including the Amazon Cone (the deepwater part of the delta), and lies west of the terminus of the Romanche fracture zone (Fig. 1). The Barreir-inhas Basin is separated on its southeast margin from the Ceará Basin by the Tutóia High, but the basin’s boundary to the north with the Pará-Maranhão Basin is poorly defi ned. Here we treat the Barreirinhas and Pará-Maranhão basins as a single basin.



Deposition and deformation in the deepwater Barreirinhas Basin

Geosphere, December 2012 1607

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The Brazilian Equatorial margin transitions from a transform continental margin in the Piauí-Ceará basins to an oblique-rifted margin in the Pará-Maranhão basins (Fig. 2). There is a zone of thinned continental transitional crust as narrow as 12 km underlying the continental slope, and a steep seafl oor slope where sediment prograde abruptly into deepwater. This forms an instable margin ideal for generating downslope mass transport by both sedimentary and struc-tural mechanisms.

The initial Aptian rifting (ca. 125 Ma) phase along the Equatorial margin had a strong dex-tral shear component that led to the creation of small pull-apart basins fi lled with thick, conti-nental sedimentary sequences (Trosdtorf et al., 2007) (Fig. 3). By Late-Albian time (102 Ma), Brazil had broken free from West Africa, end-ing dextral shear of continental crust along the margin (Antobreh et al., 2009). Oceanic waters invaded the basin from north to south during the Late Aptian (ca. 112 Ma); a lagoonal anoxic sequence, the Codó Formation (Fig. 3), overlain by the Albian (112–100 Ma) marine Canárias and Cajú Groups (Fig. 3) (Trosdtorf et al., 2007).

An oceanic connection between the waters of the Central Atlantic and South Atlantic was established during Cenomanian/Turonian time (ca. 100–90 Ma) (Antobreh et al., 2009). Fol-lowing establishment of a passive margin, local sedimentation and gravity tectonism was strongly but indirectly infl uenced by Andean tectonism. Overlying the Cajú Group is the Turonian (ca. 90 Ma) through Oligocene (ca. 22 Ma) Humberto de Campos Group, with deposi-tion of the time equivalent Areinhas Formation on the continent, the Ilha de Santana Forma-tion on the shelf, and the Travossas Formation in deepwater (Trosdtorf et al., 2007) (Fig. 3). Gravity tectonics have deformed the Travossas Formation and overlying units.

Sedimentation increases during the Miocene have been recognized north of the study area. Figueiredo et al. (2009) compared biostrati-graphic data with isotopic data to establish prov-enances and times of erosion and redeposition of sediment on the Amazon Fan and constructed paleogeographic maps for the Miocene (Fig. 4). Figueiredo et al.’s (2009) paleogeographic maps show a change of drainage direction linking the Western Amazonia wetlands to the Amazon Fan at ca. 6.8 Ma (Fig. 4).

DATA AND METHODOLOGY

Ten regional stratigraphic horizons, roughly corresponding to the main sequences described by Trosdtorf et al. (2007) (Fig. 3), were identifi ed in wells on the shelf and one well in deepwater, and tied to distinctive seismic refl ections (Fig. 2).



Krueger et al.

1608 Geosphere, December 2012

The data set consisted of a 4×8 km spaced grid of two-dimensional (2D) seismic lines cover-ing the whole deformational system. Interpreted horizons are (1) top of the sedimentary base-ment, (2) Albian/ Cenomanian (102–98 Ma), (3) Turonian (93.5–89.3 Ma), (4) Santonian (85.8–83.5 Ma), (5) Middle Campanian uncon-formity (82–78 Ma), (6) Maastrichtian uncon-formity (70.6–65.5 Ma), (7) Eocene (Lute-tian/Bartonian) unconformity (42–37.2 Ma), (8) Oligocene (Chattian) unconformity (23.03–28.4 Ma), (9) Miocene (Tortonian/Serravalain) unconformity (13.65–7.246 Ma), and (10) water-bottom.

One regional cross section 130 km long was depth converted using constant average interval

velocities for each of the map horizons. This cross section was restored using Lithotect (Geo-Logic Systems, Halliburton), assuming plane-strain deformation.

SEISMIC INTERPRETATION OF THE BARREIRINHAS BASIN

Top of Basement

The interpreted top basement surface is a combination of the top of the prerift mega sequence and crystalline basement on con-tinental crust described by Trosdtorf et al. (2007), and the top of the oceanic crust as defi ned seismically (Figs. 5–8). The basement

map (Fig. 9) shows that the transition from con-tinental crust to oceanic crust corresponds to the change from basement depths of <3000 ms on the shelf to depths of >6000 ms on the toe of the slope across a transitional zone 10–20 km wide (Fig. 5). The interpretation of the nature of the crust is based on changes in the free-air gravity anomalies (Fig. 1) and base-ment depths (Fig. 9). Rift-related basement faults formed horsts and grabens on what is now the continental shelf (Fig. 5). Rift faults cut through the top of the basement surface (red surface) and mostly terminate upward below the early drift sequence (yellow). The rift sequence can be approximated as the inter-val between the basement and the top of the

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basins

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10 m bathymetriccontour interval on shelf

250 m bathymetric contour interval on slope

wells

N

Figure 2. Location map showing onshore surface geology, offshore bathymetry, main physiographic features, and offshore basin out-lines. Saint Paul Fracture Zone (SPFZ) and Romanche Fracture Zone (RFZ) are represented on the map. Inset shows the study area with available wells, and seismic lines. Surface geology is based on the Geologic map of South America (Schobbenhaus and Bellizia, 2001) (www.cprm.gov.br). Bathymetric data is derived from ETOPO 1 grid (Amante and Eakins, 2009). Well locations and offshore basin limits are from Agência Nacional de Petróleo (www.anp.gov.br).





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Krueger et al.


Albian and is relatively thin in the study area, generally <1000 ms (Fig. 5). As previously noted, the abrupt transition zone between oce-anic and continental crust is appropriate to a transform to oblique margin (Fig. 1).

Albian (ca. 112–100 Ma)

The mapped Albian top (ca. 100 Ma; deeper yellow on the seismic lines in Figures 5–8) corre-sponds approximately to the top of the early drift sequence and is the fi rst truly time- correlative sequence in both the shelf and deepwater.

Cenomanian/Turonian (ca. 100–89 Ma)

The top of Turonian (ca. 89 Ma; purple on the seismic lines in Figures 5–8) is a continu-ous and high-amplitude surface that has a char-acteristic seismic character that can be easily correlated throughout the Equatorial margin. It corresponds to a maximum fl ooding surface at the top of the Cenomanian/Turonian condensed sequence, composed mainly of mudstones and marls (Trosdtorf et al., 2007) (Fig. 3).

The detachment surface beneath the Creta-ceous deformed rocks lies within the Turonian. Because of the steep shelf margin, there is a large difference in total depth between the con-tinental shelf and the toe of the slope within <20 km (Fig. 5).

Base of Coniacian to Top Santonian (ca. 89–83 Ma)

Two seismic units are included within this age span, the top of the Cretaceous deformed sequence horizon (cyan on the seismic sec-tions) and the top of the Santonian sequence (ca. 84–83 Ma; pink on the seismic sec-tions, Figs. 5–8). The age of the deformed Cretaceous section (cyan) is uncertain, but it lies below the top of Santonian and above the base of the Cenomanian/Turonian (Figs. 5–8) implying an age younger than the basal Turonian (ca. 89 Ma) and older than the basal Campanian (ca. 83 Ma).

Base Campanian to Top Maastrichtian (83.5–65.5 Ma)

The Campanian sequence onlaps the top of the deformed Santonian sequence (Fig. 10). At the end of the Campanian (ca. 83 to ca. 75 Ma), sea level dropped causing a change from mainly transgressive to mainly regressive sequences (Trosdtorf et al., 2007, after Haq et al., 1987). The base of the regressive sequence is the top Campanian horizon (ca. 83 Ma; yellow on the seismic sections, Figs. 5–8) and the top of the

S

B

F

AS

Middle Miocene~ 14 Ma

Late Miocene ~ 6 Ma

A

B

GEOCHRONOLOGICAL PROVINCES

Sunsas mobile belt1.25 - 1.0 Ga

Maroni-Itacaiunas mobile belt 2.2 - 1.95 Ga

Central Amazonian province > 2.3 Ga

Ventuari-Tápajos Domain 1.95 - 1.8 Ga

Rio Negro-Juruena magmatic arc1.8 - 1.55 Ga

Rondonian-Santo Ignáciomobile belt 1.5 - 1.3 Ga

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Andes

And

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Brazilian shield

Guiana shield

Brazilian shield

Figure 4. Paleogeography for northern South America (Figueiredo et al., 2009) and Ama-zonian geochronological provinces (Almeida et al., 2007). (A) During the Middle Miocene (ca. 14 Ma) the Purus arch formed a continental divide separating east from west Amazonia. Rivers west of the Purus arch formerly fl owed to the Pebas wetlands and rivers east of the Purus arch formerly fl owed to the Atlantic Ocean (Figueiredo et al., 2009). (B) Starting in Late Miocene time (ca. 6 Ma) a drainage rearrangement caused all rivers as far west as the Andes to fl ow east rendering the Pebas wetlands dry and leading to the birth of the Amazon river in its modern shape (Figueiredo et al., 2009).This drainage rearrangement is also the probable cause for the formation of the São Luís delta and the increased sedimentation on the Barreirinhas Basin. F—Amazon delta in the Foz do Amazonas Basin. B—São Luís delta in the Barrerinhas Basin. S—Solimões Basin; A—Amazonas Basin.





Fig

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—A

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Krueger et al.


Fig

ure

6. R

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ack

tim

e m

igra

ted

seis

mic

line

B–B

′. L

ocat

ion

of t

he li

ne is

sho

wn

on F

igur

e 2.

Ver

tica

l sca

le is

in s

econ

ds a

nd h

oriz

onta

l sca

le is

in k

ilom

eter

s,

aver

age

vert

ical

exa

gger

atio

n is

~1:

5. L

ine

is c

ourt

esy

of W

este

rnG

eco.

(A) S

eism

ic li

ne w

itho

ut in

terp

reta

tion

. (B

) Int

erpr

eted

sei

smic

line

sho

ws

a sy

stem

of s

mal

ler

nor-

mal

fau

lts

on t

he s

helf

def

orm

ing

a th

in C

reta

ceou

s se

quen

ce f

rom

Tur

onia

n (p

urpl

e ho

rizo

n) t

o Sa

nton

ian

age

(pin

k ho

rizo

n) c

onne

cted

to

a se

ries

of

thru

st im

bric

ates

lo

cate

d w

ithi

n th

e sa

me

litho

uni

t on

the

toe

of th

e sl

ope.

The

det

achm

ent s

urfa

ce fo

r th

e no

rmal

and

thru

st fa

ults

is a

bed

-par

alle

l det

achm

ent c

lose

to th

e pu

rple

hor

izon

of

Tur

onia

n ag

e. T

his

Cre

tace

ous

syst

em is

cro

sscu

t by

a yo

unge

r C

enoz

oic

faul

t sys

tem

. Int

erpr

eted

hor

izon

s ar

e la

bele

d on

the

line:

MIO

—M

ioce

ne, O

LIG

—O

ligoc

ene,

E

OC

—E

ocen

e, M

AA

—M

aast

rich

tian

, CA

MP

—C

ampa

nian

, SA

N—

Sant

onia

n, C

EN

/TU

R—

Cen

oman

ian/

Turo

nian

, AL

B—

Alb

ian.





Fig

ure

7. R

egio

nal p

rest

ack

tim

e m

igra

ted

seis

mic

line

C–C

′. L

ocat

ion

of th

e lin

e is

sho

wn

on F

igur

e 2.

Ver

tica

l sca

le is

in s

econ

ds a

nd h

oriz

onta

l sca

le is

in k

ilom

eter

s, a

ver-

age

vert

ical

exa

gger

atio

n is

~1:

5. L

ine

is c

ourt

esy

of W

este

rnG

eco.

(A

) Se

ism

ic li

ne w

itho

ut in

terp

reta

tion

. (B

) In

terp

rete

d lin

e sh

ows

larg

e no

rmal

fau

lts

on t

he s

helf

and

co

ntin

enta

l slo

pe li

nked

at d

epth

to a

maj

or fo

ld a

t the

bas

e of

the

cont

inen

tal s

lope

. FZ

—lo

cati

on o

f an

ocea

nic

frac

ture

zon

e. I

nter

pret

ed h

oriz

ons

are

labe

led

on th

e lin

e:

MIO

—M

ioce

ne, O

LIG

—O

ligoc

ene,

EO

C—

Eoc

ene,

MA

A—

Maa

stri

chti

an, C

AM

P—

Cam

pani

an, S

AN

—Sa

nton

ian,

CE

N/T

UR

—C

enom

ania

n/Tu

roni

an, A

LB

—A

lbia

n.



Krueger et al.


Fig

ure

8.

Reg

iona

l pr

esta

ck

tim

e m

igra

ted

seis

mic

li

ne

D–D

′. L

ocat

ion

of t

he l

ine

is

repr

esen

ted

on F

igur

e 2.

Ver

-ti

cal

scal

e is

in

seco

nds

and

hori

zont

al s

cale

is

in k

ilom

e-te

rs,

aver

age

vert

ical

exa

gger

-at

ion

is ~

1:5.

Lin

e is

cou

rtes

y of

Wes

tern

Gec

o. (

A)

Seis

mic

li

ne

wit

hout

in

terp

reta

tion

. (B

) In

terp

rete

d li

ne.

Our

in

terp

reta

tion

sh

ows

link

age

betw

een

the

norm

al a

nd t

hrus

t fa

ults

. T

he l

ine

repr

esen

ts t

he

spat

ial

exte

nt

of

defo

rmed

C

enoz

oic

rock

s th

at o

verp

rint

s th

e pr

eexi

stin

g C

reta

ceou

s de

form

atio

n.

Inte

rpre

ted

hori

zons

ar

e la

bele

d on

th

e lin

e:

MIO

— M

ioce

ne,

OL

IG—

O

ligoc

ene,

E

OC

—E

ocen

e,

MA

A—

Maa

stri

chtia

n, C

AM

P—

C

ampa

nian

, SA

N—

Sant

onia

n,

CE

N/T

UR

— C

eno

ma

nia

n/

Turo

nian

, A

LB

—A

lbia

n.





fi rst regressive package is the top of Maastrich-tian (ca. 65 Ma; green on the seismic sections, Figs. 5–8). The overall regression is character-ized by progradation of the shelf break from its position at the top of the Campanian to its more seaward position at the top of Maastrichtian (Figs. 5–7).

The Maastrichtian was a time of tectonic quiescence in the basin and the Maastrichtian sequence buries the Campanian ponded mini-basins. The top of the Maastrichtian is the youngest map horizon deposited prior to the

onset of Cenozoic folding and faulting, and is the oldest sequence without growth structures (Fig. 5).

Paleocene/Eocene (ca. 66–37 Ma)

No Paleocene horizon was mapped in this study because of limitations of the seismic resolution, therefore it is unclear if deforma-tion actually commenced during the Paleocene. The top Eocene horizon (ca. 42 Ma; light green horizon on the seismic sections, Figs. 5–8) cor-

responds to the base of the fi rst sequence that can be identifi ed as a structural growth sequence for Cenozoic fault movement.

Oligocene (ca. 34–23 Ma)

During the Oligocene sea-level rise resulted in a change from regressive in the Rupelian (ca. 34–28 Ma) to transgressive in the Chattian (ca. 28–23 Ma) forming the Upper Oligocene uncon-formity (Fig. 3) (Trosdtorf et al., 2007, after Haq et al., 1987) (orange horizon; Figs. 5–8).

3000

35002500

7000

6000

5500

4500

6500

5000

7500

2000

8000

1500

1000

1000

7000

1500

3000

2000

2500

7000

2000

2500

8000

43°0'0"W

43°0'0"W

43°15'0"W

43°15'0"W

43°30'0"W

43°30'0"W

43°45'0"W

43°45'0"W

44°0'0"W

44°0'0"W

0°15'0"S 0°15'0"S

0°30'0"S 0°30'0"S

0°45'0"S 0°45'0"S

1°0'0"S 1°0'0"S

1°15'0"S 1°15'0"S

0 10 20 30 405 km

Legend

Contour Interval: 500ms

Time (ms)

840 - 1,500

1,500 - 2,000

2,000 - 2,500

2,500 - 3,000

3,000 - 3,500

3,500 - 4,000

4,000 - 4,500

4,500 - 5,000

5,000 - 5,500

5,500 - 6,000

6,000 - 6,500

6,500 - 7,000

7,000 - 7,500

7,500 - 8,000

normal faults

N

Figure 9. Basement map showing a narrow continental slope with depths to basement changing from 2500 ms (~3500 m) near the shelf break to 7000 ms (~7500 m) in the abyssal plains over a distance of <20 km.



Krueger et al.


Miocene (ca. 23–11 Ma)

The “top Miocene” horizon (ca. 11 Ma; cyan horizon on the seismic sections, Figs. 5–8) cor-responds to a rapid transgressive event on the whole Brazilian Equatorial margin that starved the region of clastic input and gave rise to a large carbonate ramp (Fig. 3) (Trosdtorf et al., 2007).

Upper-Miocene to Present (ca. 11–0 Ma)

In the latter Miocene the region experi-enced the onset of another major clastic infl ux. The sequence above the late Miocene horizon

(ca. 11 Ma) is a prograding sequence of deep-water unconsolidated ponded basin mud and clay sequence that fi ll accommodation spaces created by the folded structures (Figs. 5–7). The present-day seafl oor is a folded, faulted, and eroded surface cut by numerous Pleistocene seafl oor canyons (Figs. 5–7).

SEDIMENTATION RATES

Sedimentation rates are a critical factor in gravity-driven fold-thrust belts because high sedimentation rates lead to steepening of the continental slope, and downslope movement

3

4

5

6

7seco

nds

(tw

t)

10 km0 km

CC C′C C′

SW NE

C C′C C′

SW NE3

4

5

6

7seco

nds

(tw

t)

Santonian mini-basins

Campanian mini-basins

Santonian mini-basins

Campanian mini-basins

Figure 10. Detail of prestack time migrated dip seismic line C–C′. Vertical scale is in sec-onds and horizontal scale is in kilometers, average vertical exaggeration is ~1:5, red arrows show lapout locations. We interpret that accommodation space was being created on the paleo seafl oor, and infi lled by the pink sequence during the Santonian. Cyan corresponds to the Coniacian/Santonian deformed sequence. Sequence in pink was probably deposited during a deformational hiatus and onlaps the cyan sequence, forming a series of ponded mini-basins fi lling the accommodation space created by faulting and folding. Sequence in yellow corresponds to lower Campanian sediment and infi lls remaining paleo water-bottom topography at Campanian time. Campanian sequence was deposited postdeformation of Coniacian/Santonian sequences in cyan and pink. Notice that sediment landward of the last fold at the bottom of the sequence represented in yellow onlap the pink sequence, but not the basinward sequence, suggesting movement on the folds continued longer landward.

serves to reduce the oversteepened surface gradient. Sedimentation rates were calculated for the single deepwater well location, high-lighted in red (zoom inset, Fig. 2), using unit thicknesses of faunally defi ned age horizons that correspond to the refl ectors in Figure 5. Depth-converted seismic stratigraphic horizons were used to estimate sedimentation rates shal-lower and deeper than the faunally defi ned hori-zons; decompaction was not taken into account. Sedimentation rates were also calculated for a point on the shelf using seismic horizons only. The results for both locations are probably per-turbed by slumping and erosion, but are the only data available. The calculated sedimenta-tion rates are summarized in Table 1 and plotted on Figure 11.

Sedimentation rates as high as 200 m/m.y. in deepwater and 90 m/m.y. on the shelf prevailed during the Coniacian (ca. 89–83 Ma; Table 1 and Fig. 11) interval. The Coniacian-Santonian is a mud-dominated interval with a sedimenta-tion rate similar to that of the modern Amazon Fan (Nancy Engelhardt-Moore, 2009, personal commun.). Fossil recovery from well cuttings was poor, consistent with a rapid deposition mass wasting environment. These high rates of sedimentation coincided with the Cretaceous rapid deformation of the Travossas Formation in deepwater (ca. 89–83 Ma).

Maastrichtian (ca. 70–67 Ma) sedimentation rates were also high, 86 m/m.y. in deepwater, and 66.7 m/m.y. on the shelf (Table 1 and Fig. 11), likely associated with an episode of tec-tonic uplift that affected the whole Brazilian Equatorial margin, as described by Zalan (2004) extending as far south as the Camamu Basin (Cobbold et al., 2010). During Late Cretaceous to Paleocene time two erosional events are observed in the deepwater but not on the shelf, and are represented on Figure 11 as extremely low sedimentation rates. We estimate that the events occurred between 83.5 and 70 Ma (7.5 m/m.y.) and 67–55 Ma (15 m/m.y.), and interpret them as local erosion on the top of large fold structures (Fig. 7).

From the end of the Paleocene to Mid Mio-cene time (55–10 Ma), rates of sedimentation were very low on the shelf (4–7.1 m/m.y.), but higher in deepwater (22.2–40 m/m.y.) areas as sediment on the shelf was eroded or bypassed, and deposited in the deep basin.

From Late Miocene to the Present (10–0 Ma; Table 1), sedimentation rates increased both on the shelf (40 m/m.y.) and in deepwater (37.5 m/m.y.). A Late Miocene, post–10 Ma, high sedimentation pulse coincides with re arrangement of the drainage system east of the Andes, and the birth of the modern Amazon drainage (Figueiredo et al., 2009), indicating that





Late Miocene drainage rearrangement affected areas farther south than previously recognized.

KINEMATIC ANALYSIS OF THE BARREIRINHAS BASIN

The postrift structural evolution of this part of the Barreirinhas Basin is dominated by a series of collapse systems involving detachment surfaces within shale units. Other fold struc-tures comparable to those analyzed have been mapped in the Barreirinhas and Pará-Maranhão basins (Zalan, 2011) and in the Foz do Amazo-nas (Cobbold et al., 2004; Araujo et al., 2009; Perovano et al., 2009), and each is a highly complex three-dimensional system. Variations in the exact timing of fault movements within the basin seem likely. Our work focuses on one representative set of structures, and does not necessarily depict timing or structural details in other parts of the basin.

Structural Palinspastic Restorations

To better understand deformation rates on the faults, fault propagation, and linkage between faults, a present-day deformed sec-tion was restored to four earlier confi gura-tions (Fig. 12). The restorations in each case assume a continuous and planar sea-bottom slope from undeformed shelf sediment to undeformed basin-fl oor sediment, following a seismic time horizon correlated to biostrati-graphic data and the seismic-stratigraphic model. The method does not account for pos-sible minor variations in sea-bottom topogra-phy, but provides an adequate basis for struc-tural analysis.

Restorations were constructed preserving bed-lengths and assuming fl exural slip/fl ow kinematics. The sections are subperpendicular

TABLE 1. SEDIMENTATION RATE OF BARREIRINHAS BASIN

Time intervalsAge

(m.y.)

Duration of deposition

(m.y.)

Environment of depositionDeepwater Shelf

Rate of deposition (m/m.y.)Mid-Miocene to Present 10 to 0 10 37.5 40Mid-Oligocene to Mid-Miocene 28 to 10 13 22.2 5.5Mid-Eocene to Mid-Oligocene 42 to 28 11 35.7 7.1Top Paleocene to Mid-Eocene 55 to 42 21 40 4Maastrichtian to top Paleocene 67 to 55 12 7.5 13.6Base Campanian to Maastrichtian 70 to 67 3 86 66.7Top Santonian to top Campanian 83.5 to 70 6.5 15Base Santonian to top Santonian 84.6 to 83.5 1.1 57Base Coniacian to base Santonian 89 to 84.6 4.4 200Top of Turonian to top Campanian 89 to 70 90Top Albian to top Turonian 102 to 89 3 76.9 28.8

Note: Sedimentation rates are in meters/million years and are calculated for a point at the toe of the slope in the middle of the study area. Sedimentation rates for the Cretaceous stratigraphy were calculated using biostratigraphic data from a deepwater well in the area. Cenozoic sedimentation rates were calculated using the age and thickness of the seismically defi ned units. The result of the different resolutions of the methodology is a higher frequency curve for the Cretaceous and a lower frequency curve for the Cenozoic.

0

1000

2000

3000

4000

5000

60000 20 40 60 80 100 120

DE

PT

H (

m)

TIME (my)

Rate of Deposition (m/my)

37.5

22.2

35.740

7.5 86

15

200

76.9

40

5.5 7.14

13.666.7

90

28.8

Figure 11. Rate of deposition plot for two locations: (1) deepwater well location (gray curve) and (2) a location on the shelf (black curve). Points in black correspond to depth of measured horizons in meters versus their age in million years. Points in red correspond to rock samples’ depths on the well and interpreted ages based on the biostratigraphy. Points are plotted for their age in millions of years versus their current depth in meters. The shal-lowest point being the seafl oor (1279 m in deepwater and 50 m on the shelf) and the deeper point the Turonian (detachment surface) depth. Steep slopes correspond to high sedimen-tation rates and low angle slopes correspond to low sedimentation rates. The sedimentation rates calculated for each interval are plotted on the curves in meters/million years.



Krueger et al.


0 km

10 k

m

5 km

0 km

10 k

m

5 km

0 km

10 k

m

5 km

5 km

FN-3FN-3FN-2FN-2

FN-1

FN-1

FR-1

FR-1

FR-2

FR-2

FN-4FN-4FN-3FN-2

FN-1

FR-1

FR-2

0 km

10 k

m0

km75

km

FN-4

S=

~21

69 m

E=

~14

94 m

S=

~53

4 m

E=

~42

4 m

S=

~43

3 m

E=

~17

8 m

S=

~43

3 m

E=

~17

8 m

0 km

10 k

m

5 km

83.5

Ma

42 M

a

27 M

a

10 M

a

0 M

a

Fig

ure

12. P

alin

spas

tic

rest

orat

ions

. Dur

ing

the

Sant

onia

n (c

a. 8

4 M

a) a

pro

grad

ing

shel

f an

d hi

gh s

edim

enta

tion

rat

e in

dee

pwat

er (

200

m/m

.y.)

ca

used

slo

pe in

stab

ility

and

tri

gger

ed t

he f

orm

atio

n of

a s

erie

s of

list

ric

norm

al f

ault

s on

the

she

lf a

nd t

hrus

t fa

ults

at

the

toe

of t

he s

lope

. Dur

ing

the

Eoc

ene

(ca.

42

Ma)

a s

econ

d de

form

atio

n ev

ent

star

ted

to d

evel

op a

s th

e sh

elf

mar

gin

colla

psed

. Def

orm

atio

n co

ntin

ued

duri

ng t

he O

ligoc

ene

(at

ca. 2

7 M

a) w

ith

mot

ion

on b

oth

norm

al a

nd t

hrus

t fa

ults

, but

def

orm

atio

n ra

tes

wer

e sl

ow. D

urin

g th

e M

ioce

ne, a

fter

ca.

10

Ma,

def

orm

atio

n ra

tes

incr

ease

d si

gnifi

cant

ly, w

ith

maj

or n

orm

al g

row

th f

ault

ing

on t

he s

helf

mar

gin,

and

fol

ding

at

the

toe

of t

he s

lope

. The

re a

re a

ctiv

e an

ticl

ines

on

the

sea

fl oor

. You

nger

nor

mal

fau

lts

deve

lope

d in

the

foo

twal

ls o

f pr

eexi

stin

g no

rmal

fau

lts.

Maj

or c

anyo

n sy

stem

s w

ere

inci

sed

into

the

she

lf

mar

gin,

cut

ting

bot

h no

rmal

fau

lts

and

grow

th f

olds

pro

babl

y du

ring

the

last

10

Ma.





to the trends of folds, thrusts, and normal faults. Note that there are cross-structures that represent shortening along the b-kinematic axis of thrust sheets, but along-strike strain is generally <1%, and therefore negligible for our purposes.

Present (0 Ma)At present most of the deformation in the

study area is distributed among four normal faults and two large thrust faults. The normal faults (FN-1, FN-2, FN-3, and FN-4) and thrust

faults (FR-1 and FR-2) are labeled in Figure 12. There is also a smaller back-thrust fault associated with a fold collapse feature but that is a minor structure (Fig. 12). The present-day shortening measured on our representative cross section is ~2200 m and the present extension is ~1500 m (Fig. 12).

Miocene (10 Ma), Oligocene (27 Ma), and Eocene (42 Ma)

Displacement on the four normal faults and the two thrust faults was restored to the

paleogeometries of 10, 27, and 42 m.y. ago. Restorations of the normal faults resulted in extensions of 420 m (10 Ma) and 180 m (27 and 42 Ma). Restoration of the thrust faults resulted in shortening of 530 m (10 Ma) and 430 m (27 and 42 Ma). Most of the deforma-tion (75% of the shortening and 72% of the extension) could be restored at 10 Ma (Fig. 12). After the second restoration time-step (27 Ma) 80% of the shortening and 90% of the extension is resolved, no change resulted from 27 to 42 million years (Fig. 12).

Figure 13. Detail of the prestack time migrated dip seismic line B. Vertical scale is in seconds and horizontal scale is in kilometers, average vertical exaggeration is ~1:5. Cenozoic deformation is highlighted, the pregrowth sequences are in green and the growth sequences are in yellow and blue.



Krueger et al.


Santonian (83.5 Ma)During the Coniacian-Santonian (Fig. 12) a

prograding shelf and a high sedimentation rate in deepwater (226 m/m.y.) caused slope instabil-ity and triggered the formation of a set of normal listric faults on the shelf and thrust faults at the toe of the slope. This wide linked extensional-compressional system developed very rapidly, all the deformation and the infi lling of the deformed seafl oor took place during the Santo-nian (ca. 86–84 Ma) within the sequences repre-sented in blue and pink in Figure12. Because of limited stratigraphic resolution within this inter-val, and poor seismic imaging due to subsequent deformation by a younger fault system (Fig. 13), detailed interpretation of individual faults is not possible. Depiction is schematic for this event (Fig. 12).

Cenozoic Fault Analysis

Integration of detailed structural analysis with a well-defi ned stratigraphic age model provides the opportunity to determine fi nite (i.e., long-interval) fault motion rates over the Cenozoic history of the basin (Fig. 13). For each fault the fault-parallel displacement was measured for a representative pregrowth section, i.e., the top of Cretaceous (Fig. 14).

The identifi cation of the timing of fault motion also allows us to infer linkages of vari-ous faults in the systems; shelfal normal faults moving simultaneously with basinal reverse faults can reasonably be assumed to link across the intervening translational zone.

The displacement versus time plot (Fig. 15) demonstrates a variable deformation rate

4000

3600

FR

-2F

R-2

FR-1

FR-1

FR

-2

4000

3600

3200

2800

4800

5400

800

43° 45′ W

1° 1

5′ S

1° 0

′ S

1° 1

5′ S

1° 0

′ S

43° 30′ W

43° 45′ W 43° 30′ W

0 10 km

B A

C′

B′

A′D

D′

FR-1

FN-1

FN-1

FN

-2F

N-2

FN-3FN-3

FN-4

FN-4

FN-1

FN

-2

FN-3

FN-4

1800

1600

1600

thrust faults normal faults

C

1200

48004400

5200

Figure 14. Structural map on the top of the Cretaceous section (top of pregrowth section) showing the location of seismic lines used in this study. White dashed line represents updip and downdip limits of the earlier Turonian-Santonian deformation system. The thrust faults (FR-1 and FR-2) and normal faults (FN-1, FN-2, FN-3, FN-4) used in our structural analysis are labeled on the map.

through time. Deformation that began in the Eocene (ca. 42 Ma) continued during the Oli-gocene with motion on both normal and thrust faults, but deformation rates were always slow (Fig.15). Deformation rates increased sig-nifi cantly in the Miocene as indicated by the expanded Miocene section on the downthrown side of shelf-margin normal faults (Fig. 3). In post-Miocene time, deformation rates continued to increase (Fig.15), with major normal growth faulting forming synclines on the shelf margin, and fold crests rising toward the sea surface at the toe of the slope (Fig.13). Additional normal faults (FN-3 and FN-4) developed in the foot-wall of preexisting normal faults (FN-1 and FN-2) (Fig. 13).

DISCUSSION

Coniacian/Santonian Bed-Parallel Gravity Gliding

The very brief time interval of ca. 89–84 Ma (duration ca. 5 Ma) corresponds with a period of eustatic sea-level fall (Trosdtorf et al., 2007, after Haq et al., 1987; Fig. 3) and increased tectonism in the Andes (Zalan, 1998), which resulted in a very high sediment infl ux of 226 m/m.y. to the basin. The combination of a steep basement slope (10°–15°) above the narrow continental to oceanic crustal transi-tion zone with this period of high sedimenta-tion rate and consequent oversteepening of the surface slope led to instability of the slope. This Santonian slope instability generated a set of linked listric normal faults on the shelf and thrust faults at the toe of the slope. On the shelf a thin deformed sequence is character-ized by a 30-km-wide zone of listric normal faults detached within the underlying sequence of Cenomanian to Turonian marls and shales (Fig. 10). The extensional domain is linked by a 30-km-wide translational domain without visible internal deformation to a compressional domain. On the toe of the slope the sequence is deformed by a set of landward dipping thrust faults forming a belt of imbricate thrust sheets with 30 km in the dip direction (Fig. 5) and 30 km in the strike direction (Fig. 8).

Zalan (2011) concluded that Cretaceous gliding and thrusting marked the onset of deformation that was semicontinuous through-out the Cenozoic. However, consistent age of sediment infi ll of local basins created by nor-mal faults on the shelf and by thrusts in the basin (Fig. 16), and cross-cutting relationships in which Cenozoic faults sole at a deeper strati-graphic level and rotate the Cretaceous thrusts (Fig. 13), indicate two discrete tectonic events separated in time.





Syndeformational Santonian age sediment was deposited on the top of both the rotated blocks of the shelf and the folds at the toe of the slope. This onlapping sequence is highlighted in pink in Figure 10. The isochron map for the inter-val between the top of the Turonian and the top of the Santonian (Fig. 16) includes the thickness of the Cretaceous syndeformational sequence and part of the postdeformational sequence; most of the sediment, up to 1000 ms, accumu-lated on the slope in pocket mini-basins similar to those described by Hooper et al. (2002) and Corredor et al. (2005) in the Niger Delta.

The Cretaceous allochthon is ~1 km thick (Fig. 10) and ~70 km in downdip extent (Fig. 14). Gliding was facilitated by a combination of a low basal slope and high pore pressures in the Turonian shales, but even with an extremely effi cient detachment the area-to-thickness ratio

of the allochthon is unusual. Stratigraphic resolution is limited within this thin interval, and seismic imaging is poor due to subsequent deformation by the Cenozoic fault system (Fig. 13). Detailed interpretation of individual faults is not possible, and depiction is schematic for this event (Fig. 12).

Maastrichtian through mid-Eocene Structural Quiescence

During this time interval, the shelf margin continued to prograde into deepwater, but was structurally stable with the exception of reac-tivation of one basement-involved fault in the narrow zone of extended continental crust (Fig. 6), an observation also made by Zalan (2011).

The isochron map between the Maastrichtian and the Eocene horizons (Fig. 13) shows thick

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and thin axes of deposition of the Paleocene/Eocene sediment (ca. 65–37 Ma). On the shelf break, sediment eroded from the footwall and redeposited on the hanging wall of the normal faults, forming elongated depocenters parallel to the shelf break (Fig. 17). On the continen-tal slope, sediment was deposited while thrust faults moved, resulting in the establishment of a growth sequence. Sediment deposited on the hanging wall of the thrust faults and formed ponded mini-basins. Thrust fault movement formed folds on their hanging walls that created anticlines on the Paleocene/Eocene seafl oor. These anticlines were soon partly eroded and sediment redeposited on the thrust fault foot-walls (Fig. 17).

Mid-Eocene Onset of Nonbed Parallel Gravity Sliding

During mid-Eocene the shelf margin col-lapsed, developing two normal faults (FN-1 and FN-2) and a thrust fault (FR-2 in Fig. 13). The two normal faults linked to thrust faults FR-1 and FR-2, forming concave detachment faults (Fig. 13).

Deformation in the system persisted for at least another 40 m.y., until the Present. The Ceno-zoic deformed area is less extensive (~30 km) both downdip and along strike (Fig. 14) than the Cretaceous deformed area (Fig. 18).

The Eocene fault systems cross-cut bedding and form two bowl-shaped fault systems at two different depths (Fig. 13). The bowl geometry is different from the classic shale detached deep-water fold and transform belt observed in Nige-ria (Corredor et al., 2005; Cobbold et al., 2009) and described here for the Cretaceous section (Fig. 18).

The isochron map (Fig. 17) of the thin mid-Eocene to Oligocene interval suggests that dur-ing the Oligocene transgression, most of the sediment was trapped on the continental shelf and slope. On the continental slope, sediment was trapped in ponded mini-basins created by normal and thrust faults. The largest fold was partly breached by erosion, allowing sediment to bypass to the abyssal plain (Fig. 17).

The Oligocene to mid-Miocene unconformity isochron map (Fig. 17) indicates a time of ero-sion on the shelf and deposition in deepwater. The isochron map shows erosional channeling of the continental shelf and slope, and breach-ing of the major anticline continued to allow bypassing of sediment onto the abyssal plain.

Mid-Miocene Acceleration of Deformation

Deformation rates increased dramatically in mid-Miocene time, as indicated by the expanded pre–10 Ma Miocene section on the



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Figure 16 (Continued on following pages). Cretaceous isochrons representing fault growth in the drift units and related sedi-mentary evolution. Three Cretaceous isochrons are represented in the fi gure: Turonian (89 Ma) to Santonian (84 Ma), Santonian (84 Ma) to Campanian (78 Ma), and Campanian (78 Ma) to Maastrich-tian (65.5 Ma) litho units. Thin isochrons areas are associated with canyons, gullies, and other erosional features and thick iso-chrons areas are associated with deposi-tional fairways, fans, ponded mini-basins, fault growth, and prograding wedges. The interpreted features are shown below each isochron map.





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downthrown side of the shelf-margin normal faults (Fig. 14). This accelerated fault motion corresponds in time to increased sedimentation rate (Fig. 11), and a transition from a carbon-ate ramp to prograding siliciclastics along the entire Equatorial margin.

The Late-Miocene to Present rapid deposi-tion, highlighted in blue in Figure 13, generated a thick, progradational sequence depicted in the Miocene to Present isochron map (Fig. 17). The accumulation is the thickest sequence deposited in the ponded mini-basins, and caused mini-basins to coalesce. Localized depocenters are thickest near the Miocene shelf break, above the hanging walls of the normal faults, and in basins that formed behind the major anticline on the thrust faults at the toe of the slope (Fig. 17). The fold was locally breached, allowing sediment to be deposited onto the abyssal plain.

It is likely that the high sedimentation and deformation rates observed in the Barreirinhas and other basins during the Late Miocene are an indirect consequence of the major drainage reor-ganization in northern South America that began ca. 11 Ma (Altamira-Areyan, 2009) and diverted the drainage west of the Purus arch from the Caribbean and into the present-day Lower Ama-zon Basin (Fig. 4) (Figueiredo et al., 2009). That may also be the case for other large deepwater

fold and thrust structures developed on the Bra-zilian Equatorial margin from the Amazon Cone (Araujo et al., 2009; Perovano et al., 2009) to the Barreirinhas Basin (Zalan, 1998, 2004, 2005, 2011; Gilbert, 2006, 2011; Krueger and Gilbert, 2009; Krueger et al., 2011).

Deformation rates continued to increase, with major normal growth faulting on the shelf mar-gin, and uplift of folds at the toe of the slope (Fig. 13). Approximately 80% of the net strain in the area took place within the last 10 m.y. (Fig. 12).

Two new normal faults (FN-3 and FN-4) developed in the footwall of preexisting nor-mal fault FN-2, and linked to the existing FN-2/FR-1 system (Fig. 13). Fault motion rates increased signifi cantly (Fig. 15) as indi-cated by the expanded Miocene section on the downthrown side of the shelf-margin normal faults (Fig. 13).

Major Pleistocene to Holocene canyon sys-tems were subsequently incised into the shelf margin, cutting both normal faults and growth folds (Fig. 17). The three-dimensional effects of the deformation can be seen in the Miocene (10 Ma) to Present isochron map (Fig. 17) that shows development of a thick ponded mini-basin bounded by normal faults landward and thrust faults basinward.

EXTENSION SHORTENINGTRANSLATION

EXTENSION SHORTENINGTRANSLATION

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Figure 18. Geometry of the two fault linked systems that form the Barreirinhas fold and thrust belt in our work area. (A) Cenozoic age fault system. Normal faults on the continental shelf link at depth with thrust faults on the continental slope together forming a concave detachment system. The main detachment fault cross-cuts stratigraphy and has normal dis-placement updip and reverse displacement downdip. (B) Cretaceous fault system. Normal and thrust faults detach at a bed parallel surface.

Driving MechanismsShale-detached deepwater fold systems,

similar to those of the Barreirinhas Basin, have been described at several continental margins, particularly the Niger Delta (Damuth, 1994; Hooper et al., 2002; Rowan et al., 2004; Krueger and Gilbert, 2006, 2009; Sultan et al., 2007), the Pará-Maranhão and Barreirinhas basins (Zalan, 2005, 2011; Gilbert, 2006), and offshore Namibia (Butler and Paton, 2010). Two factors must be considered in the kinematics of these thrust belts, the generally accepted mechanism that allows thrust motion, and the “triggering event” that initiates deformation.

The critical taper wedge model of Davis et al. (1983) and Dahlen (1984) is often cited as the causative mechanism for thrusting, but it must be borne in mind that the ultimate cause of deformation is the condition or conditions that created the key tapered-wedge requirements: a basal detachment that slopes toward the hinter-land, and a surface slope toward the foreland. In deepwater fold belts, it is generally accepted that these conditions are largely driven by rapid sediment progradation linked to a sea-level drop or tectonic events. In addition, deepwater fold and thrust belts are invariably linked to gravita-tional collapse and updip extension on the shelf, and therefore are not strict analogs to the fore-land fold and thrust belts and subduction zones from which the model was developed.

In areas of very thick sediment accumula-tion such as the Amazon Cone (Araujo et al., 2009; Perovano et al., 2009) and Niger Delta (Corredor et al., 2005; Billotti and Shaw, 2005; Cobbold et al., 2009), the sediment weight depress the lithosphere (Morley et al., 2011) causing the slope of the basal contact of the sedimentary pile to dip landward, which is also observed to a lesser extent in the Barreirinhas Basin (Zalan, 2011).

In the Barreirinhas Basin, a steep continental slope combined with rapid sediment prograda-tion generated an unstable surface slope. The evolution of the sedimentary wedge top through time was measured on the seismic lines and on the structural restorations: ~4° during the Santonian deformation, 3.3° in mid-Miocene time, 3.6° in the mid-Oligocene, 3.8° in mid- Miocene, and steepening abruptly to 5° in the mid-Miocene to Present, the time of major deformation (Fig. 12).

The compression on the toe of the slope is caused by friction at the detachment level and cohesion of the sliding rocks. Zalan (1998) associated compression at the toe of the slope with a slowdown in the gravity-driven move-ment of the sediment due to either (1) a change in the gradient of the detachment layer or (2) the buttressing effect of a more rigid body, such as



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an igneous intrusion, an ancient volcano, or a protruding rift-phase domino-type fault block. Where observed, buttressing effects are rela-tively isolated, and have the localized effect of forming imbricate fans, and locally defl ect the trend of the fold belt (Fig. 8). In the absence of any of the other features, we interpret the change in gradient as the causative factor for the compression at the toe of the slope.

At the toe of the slope, compressional defor-mation by thrusting and folding occurs when the taper conditions are no longer suffi cient to sup-port basal shear on the sediment pile (Davis and Kusznir, 2004; Dahlen, 1984). The change in basal gradient typically occurs at the top of the oceanic crust. The loss of water-bottom gradient can be the result of accumulation of sediment at the toe of the slope via thrusting and folding or redeposition of eroded sediment from the shelf (Fig. 6). In the Barreirinhas Basin the basal slope is close to zero, therefore sediment surface is the key element to allow further thrusting.

The ultimate origin of the deformation is more problematic. Zalan (2011) speculated that earthquakes associated with reactivation of nearby oceanic fracture zones could have been a triggering mechanism for deformation. Undoubtedly such events could trigger discrete movement events, but given the long (40 m.y.) duration of the Cenozoic faulting, and the direct temporal correspondence of high deformation rates with high sedimentation rates, sediment loading is favored here as the controlling fac-tor. Distal tectonic events in the Andes may have infl uenced Equatorial margin gravity deforma-tion indirectly by impacting sediment input.

SUMMARY AND CONCLUSIONS

The structural architecture of the Barreirinhas Basin is dominated by two major deepwater fold and thrust belts linked landward to extensional fault systems, but the Cretaceous and Cenozoic faults and folds have markedly different charac-teristics. The short-lived Cretaceous system is <1.5 km thick, and involved listric normal faults and small stacked imbricate thrusts linked by a bed-parallel décollement. The Cenozoic fault system cuts through 4 km or more of Cretaceous and Cenozoic sediment, and cross-cut the preex-isting Cretaceous deformed sequence. Normal faults connect to the thrust faults at depth, and cut across bedding, forming two bowl-shaped “mega-slumps.”

Integration with an age model developed from sequence stratigraphy and paleontological data allows the determination of fi nite fault motion rates for the Cenozoic system. Movement on both normal and thrust faults began at the same time, and fault motion rates for both shortening

and extension varied simultaneously, suggest-ing linkage between normal and thrust faults in deformation rate and net strain. This supports the idea that extensional deformation on the shelf is being accommodated by shortening on the toe of the slope. In addition, the results dem-onstrate a temporal relationship between sedi-mentation rate and fault motion rates, verifying that gravitational loading by rapid sedimenta-tion is likely the driving mechanism.

Relative fault timing indicates that the exten-sional province propagated landward, and that thrust faults raised the elevation of the lower slope, lowering the slope gradient. As more faults were introduced into the system the loca-tion of the shelf break moved landward, main-taining the slope gradient at or below 5°.

Establishing deformational mechanisms and strain rates helps to fi ll the gap between our knowledge of small-scale, geologically instan-taneous gravity-driven submarine slumps or slides (a sedimentological phenomenon) and the larger-scale, slower gravity-driven thrusts (a structural phenomenon). Closing this gap may aid in understanding the generation of mass-fl ow deposits at steep margins, with implica-tions for sedimentology, basin analysis, and hydrocarbon exploration.

ACKNOWLEDGMENTS

WesternGeco kindly provided permission to pub-lish the four seismic lines depicted here. Much of the mapping and structural restoration was done through the auspices of Devon Energy, and we especially acknowledge permission to use the interpreted seismic grids that are the basis of the isochron maps. Structural restorations were performed using LithoTect software. Special thanks to Michael Hankins (HRT America, formerly of Devon); to Pedro Zalan, Ivo Trosdtorf, and Jorge Picanzo Figueiredo (Petrobras); and to Dale Bird (Bird Geophysical) for discussions of the local and regional geology. Nancy Engelhardt-Moore of Devon Energy provided key information on the impli-cations of the paleontological data to both age-dating and sedimentation rates. Also special thanks to Laura Unverzagt (University of Houston) for help with Arc-GIS, and lastly, the editors and reviewers, especially Peter Cobbold, who provided much-needed sugges-tions for revisions.

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