chapter1 final

Gravity Anomalies, Flexure and

the Thermo-Mechanical Evolution

of the West Iberia Margin and its

Conjugate of Newfoundland

by

Tiago Cunha

Thesis submitted for the Degree of Doctor of Philosophy, University of Oxford

Wolfson College &

Department of Earth Sciences Trinity Term 2008

iGravity Anomalies, Flexure and the Thermo-Mechanical Evolution of

the West Iberia Margin and its Conjugate of Newfoundland

Tiago Cunha

Wolfson College & Department of Earth Sciences

Thesis submitted for the Degree of Doctor of Philosophy to the University of Oxford, Hilary

Term 2008

The West Iberia (WIM) and Newfoundland (NFM) continental margins formed over

a succession of rift events related to the opening of the North Atlantic between the Late

Triassic and the Early Cretaceous. They are characterized by a variable width Ocean-

Continent Transition (Zone) where exhumed and serpentinized mantle has been cored.

Results from 1-D well backstripping along the Portuguese shelf suggest 40-45% greater

extension factors () than previous estimates. In addition, the age and duration of both

the Late Triassic-earliest Jurassic and the Late Jurassic rifting episodes have been better

constrained. It appears, for example, that the Late Jurassic rift propagates northwards

along the margin, as inferred for the continental break-up during the Early Cretaceous.

Combined backstripping and gravity modelling techniques (POGM), together with

new bathymetry and sediment thickness grids, have been used to estimate the effective

elastic thickness, Te, of the lithosphere. Results along closely space profiles in the WIM

reveal that Te decreases from 15-40 km over unthinned Variscan basement to 15 km

over stretched continental and transitional crust. Along strike, a good correlation is found

between the modelled mechanical structure and the segmentation of the Variscan base-

ment onshore. Discrepancies between observed and calculated anomalies are discussed in

terms of other processes that affected the margin, such as serpentinization and inversion.

Thermo-mechanical rheological modelling techniques have also been tested. The com-

piled software inputs the s constrained from POGM to determine the thermal structure

of the margin, which, in turn, has been used to construct models of rheology and, hence,

Te. The models predict an increase in the Te of thinned continental lithosphere with

age since rifting, between < 20 and > 35 km, and suggest that processes such as ductile

shearing and mantle serpentinization might have permanently weakened the lithosphere.

Results from POGM along conjugate profiles off WI and NF reveal a greater long-

term strength of extended continental and transitional basement in the NFM (Tes of

10-40 km). This analysis further supports the hypothesis that the low Tes modelled in

the WIM might be related to its complex structural framework. In order to explain the

large asymmetries observed in the amount of basement subsidence (500-1300 m) between

the WI and NF margins a lithospheric-scale simple shear rift model is proposed for the

latest stages of rifting and continental break-up.

ii

Extended Abstract

The West Iberia (WIM) and Newfoundland (NFM) margins are arguably the best

studied pair of conjugate, amagmatic (or magma-poor) rifted continental margins in the

worlds oceans. During the last three decades they have been targeted by numerous gov-

ernment, academic and commercial groups. Both margins have been surveyed using high

quality seismic and sonar and have been subjected to both commercial and scientific

ocean drilling. The lack of significant sediment thicknesses beneath the continental slope

and rise provides the ideal setting to study the nature of the basement, where rotated

and half-graben rifted faulted blocks have been extensively imaged. In the WIM, sam-

ples of serpentinized upper mantle have been recovered at a number of basement highs,

confirming the existence of a region of transitional basement, commonly interpreted as

consisting of exhumed and serpentinized upper mantle, with a width that can vary from

50 km to 150 km.

The WI and NF conjugate margins formed during a number of rifting events related to

the opening of the North Atlantic between the Late Triassic and the Early Cretaceous.

The main episodes include: (1) the Late Triassic-earliest Jurassic rift, which resulted

in the establishment of a large evaporitic province, extending over most of the Central

and North Atlantic; (2) the Late Jurassic extensional episode, contemporaneous with

significant plate reorganizations and the break-up between the African and North Amer-

ican plates; and (3) the latest Jurassic-Early Cretaceous rift, which culminated with the

break-up between Iberia and the Grand Banks of Canada.

Throughout its post-rift history, the plate tectonic setting of Iberia changed consider-

ably, acting either as an independent plate or as an accreted terrane of the African or the

Eurasian Plate. During the latest Cretaceous-Paleogene and Neogene periods, the mar-

gin underwent compressional deformation related to the Pyrenean and Betic orogenies,

respectively.

Results from 1-D well backstripping along the Portuguese continental shelf reveal im-

portant aspects of WIM-NFM rift. For example, they suggest a northwards propagation

of the Late Jurassic rifting event, similar to that inferred for the late stages of rifting

and continental break-up during the Early Cretaceous. Furthermore, they highlight the

iii

importance of both regional and local (e.g. strike-slip and halokinesis) tectonic processes

in the evolution of the main Meso-Cenozoic depositional systems along the WIM.

By comparing the subsidence curves obtained from well backstripping to theoretical

thermal model predictions, constraints have also been placed in the amount of thinning

() at different stages of the margin evolution. For the Late Triassic-earliest Jurassic

rifting event the calculated s vary from 1.16 to 1.32 and represent between 60 to 80% of

the total amount of extension. The total s modelled in this study suggest 40-45% more

extension than previous estimates. Uncertainties in the determination of the porosity-

depth relationships, intense halokinesis during extensional and compressional episodes,

and flexural compensation of the sedimentary loads can, however, affect these estimates.

Most studied wells show an important stratigraphic hiatus which extends from the

latest Jurassic to the late Early Cretaceous. In some wells of the Lusitanian Basin this

hiatus extends over the entire Late Jurassic and part of the Middle Jurassic. A simple

estimate, based on the deviations between the calculated and theoretically predicted

tectonic subsidence curves, suggests that up to 1500 m of sediments have been eroded

from the shelf during this period.

Recently compiled datasets of bathymetry and sediment thickness, together with

Process Oriented Gravity Modelling (POGM) techniques, have been used to estimate

the flexural rigidity, or equivalent effective elastic thickness, Te, of the lithosphere along

the WIM. The bathymetric compilation is based on ship-track data from the National

Geophysical Data Center (NGDC) and soundings from the Portuguese Hydrographic

Survey (IHP), and its accuracy was estimated to be better than 100 m from cross-over

error analysis. The sediment thickness dataset is based on previously existing structural

maps, recently acquired high quality seismic data, and offshore and onshore well data.

Three distinct structural maps have been compiled: (1) depth to basement; (2) top

Jurassic-mid Aptian unconformity; and (3) mid Eocene-Oligocene unconformity. The

conversion from two-way travel time to depth used a number of velocity-depth functions

through the sediment column. These were constrained from borehole data and from

wide-angle and multi-channel seismics, and show a considerable spatial variation along

and across the strike of the margin.

The results from POGM along closely spaced profiles (25-50 km) have been analysed

in the context of the structural framework and the first-order segmentation observed in

the bathymetry and free-air gravity anomaly (FAA) signature of the WIM. In most of the

studied transects, a decrease in the Te used to modelled the effects of the sediment loading,

Te(sed), is observed between relatively undeformed Variscan basement (i.e. 1.5;

Te(sed) of 10-40 km) and highly stretched continental crust (i.e. 2; Te(sed) < 15

km). Along strike, the largest variations in the modelled Te(sed) occur between the

iv

Central Iberian Terrain in the north, where Te(sed) varies between 20 and 40 km, and

the Ossa Morena and Southern Portuguese terrains, south of the Porto-Tomar Variscan

fault, where Te(sed) varies between 0 and 25 km.

It appears, however, that a model which assumes zero strength during rifting, Te(rift)

= 0 km, and a spatially varying Te(sed) cannot explain a number of prominent features of

the WIM free-air anomaly gravity field. These include, for example, the high amplitude

edge-effect anomalies along the central and southern WIM, between the continental

shelf and rise, the broad negative anomalies over the Iberia and Tagus abyssal plains,

and the high amplitude positive anomaly (up to 80 mGal) along the lower continental

slope of the SW Portuguese Margin. Large isostatic anomalies (IA > 50 mGal) have

been computed associated to these features along a number of profiles.

A model which accounts for finite strength during rifting (i.e. Te(rift) > 0 km) and for

variations in the depth of stress maxima, or necking depth (Zneck), was therefore tested.

Along most of the profiles, such a model explains well the amplitude of the edge-effect

anomaly and improves significantly the fit between the modelled and the seismic Mohos.

Again, a good correlation was found between the modelled parameters and the zonation

of the Variscan orogen. North of the Porto-Tomar fault, Te(rift) varies between 5 and 10

km and the Zneck is shallower than that predicted by the Airy model of rifting (5.5-6.5

km). To the south, a Te(rift) of 10 km and a necking depth between 9 and 12 km was

consistently modelled.

The broad, high amplitude, negative FAAs over the Iberia and Tagus abyssal plains

cannot be explained from the contributions of rifting and sediment loading alone. For the

Southern Iberia Abyssal, where these anomalies are greater (up to -90 mGal), a model

which combines flexural deformation associated with the northwards thrusting of the

Estremadura Spur and the Madeira-Tore Seamount, and upper mantle serpentinization

and/or mantle depletion (5-10%) has been proposed. The model is consistent with the

available constraints on the structure and composition of the transitional basement and

suggests the existence of a 3-4 km layer of cooled basaltic melts overlying an 4 km

thick layer of < 25% serpentinized upper mantle.

In the Southwest Portuguese Margin, the short wavelength, positive-negative IA cou-

ple observed over the continental slope and rise, respectively, has been modelled as caused

by intra-crustal thrust loading and associated flexural deformation. This interpretation is

consistent with the location of a recently interpreted major thrust front along the base of

the continental slope, here designated as the Tagus Abyssal Plain Thrust (TAP-thrust).

Tests using POGM show that the Tes associated to the emplacement of these loads,

Te(thrust) 35, are much greater that those inferred at the time of rifting, Te(rift)

10 km. Since most extensional deformation took place during the latest Jurassic-Early

vCretaceous, and the thrust loads were not emplaced until the mid Miocene to Recent

(i.e. during the Betic Orogeny), then these results suggest that the strength of thinned,

extended continental crust increases with age since rifting.

When combined, the seismic and gravity data in the SW Portuguese Margin suggest

that the mid Miocene to Recent compressional deformation is being accommodated over

an arcuate 100 km wide and 150 km long region, between the Gorringe Bank and the

Setubal Canyon. The recently mapped structural features might also have important

implications for the seismicity of the region. This study proposes a rupture mechanism

to explain the 1755 Great Lisbon Earthquake and Tsunami, which involves failure along

the TAP-thrust, the Marques de Pombal Fault and probably also the Horseshoe Fault;

all of which appear to be linked by major NW-SE trending transfer Zones.

The s constrained from POGM have been used as input in a recently compiled

thermo-mechanical rheological modelling software, the Mrift code. The software com-

bines pre-existing routines to compute the thermal structure of rift basins with analytical

solutions that calculate the Te based on yield strength envelope (YSE) considerations;

i.e. which take into account the rheology of the crust and lithospheric mantle as well as

the bending stresses caused by the sediment loading.

By considering realistic multi-stage continental rift scenarios, this modelling approach

has been used to further investigate the subsidence/uplift history of the WIM and un-

derstand the mechanisms that control the spatial and temporal variation in the Tes

recovered from combined backstripping and gravity modelling. Moreover, by integrating

some a priori information on the margin paleowater depths, the model also allows for the

sediment loading history to be determined. From the sediment loads, and the associated

flexural deformation, a theoretical basin stratigraphic profile is generated, which can then

be compared to the observed margin stratigraphy.

The Tes computed analytically from the thermal structure of the margin are con-

sistent with those estimated from POGM over relatively unstretched continental crust.

For the Galicia Margin, for example, the modelling results suggest a reduced radiogenic

heat-production within the crust, in good agreement with the existing heat-flow data

along the shelf. Over stretched continental crust, the calculated Tes increase from 10-20

km during rifting to > 35 km at 100 m.y. after rifting has ceased, as predicted from

POGM techniques. Higher overall Tes over stretched continental and transitional base-

ment suggest that processes such as widespread normal faulting, shearing along low-angle

detachments and upper mantle serpentinization might affect the long-term strength of

the lithosphere in these regions.

By combining low flexural strengths during rifting with shallow to intermediate pale-

owater depths between the Late Triassic and earliest Jurassic (up to 400 m) and between

vi

the Middle and Late Jurassic (up to 800 m), a reasonable good fit is obtained between the

modelled and the observed stratigraphy along the margin. From the Early Cretaceous

onwards, the predicted paleowater depths show an abrupt deepening of the WIM between

the Early Cretaceous and the early Late Cretaceous, followed by a period of more gentle

margin subsidence between the early Late Cretaceous and the Oligocene. The predicted

paleowater depths are in good agreement with paleoenvironmental data recovered from

ODP and DSDP drilling. Between the Oligocene and the present it appears that the

subsidence kept pace with the sedimentation along most of the margins deep offshore (

20 m/m.y.).

Results from POGM along conjugate profiles off West Iberia and Newfoundland show

some discrepancies in the way Te varies both along and across the margins strike. Over

highly extended continental and transitional basement, for example, the Tes in the New-

foundland Margin are comparatively high, varying between 10 and 40 km. These differ-

ences appear to reflect, to a certain extent, the more complex structural framework of the

WIM, marked by widespread normal faulting and conspicuous low-angle detachments.

Important asymmetries are also observed in the geometry of the Moho and the base-

ment total tectonic subsidence (TTS) between the WI and NF margins. The greater

basement depths in the WIM (between 500-1300 m deeper, approximately) extend from

thinned continental to unambiguous oceanic crust, across different widths of transitional

type basement, and cannot be explained by lateral variations in the composition of the

crust and/or upper lithospheric mantle. Furthermore, no consistent relationship has

been found between the margins state of flexure and the asymmetries observed in the

structural framework and the TTS across a number of conjugate transects.

Recently acquired wide-angle and refraction seismic data along the WI and NF mar-

gins show similar amounts of crustal thinning between the continental shelf and rise.

Crustal stretching factors recovered from both the gravity modelling and the seismic

data have been used to compute the margins thermal structure and subsidence history.

The results from the thermal models suggest that, along the NFM the mantle lithosphere

underwent significantly less stretching than that observed in the crust. In order to explain

the observations, a lithospheric-scale simple shear model, characterized by a deeping

to the east major shear zone, is proposed for the late stages of rifting and continental

break-up between the West Iberia and Newfoundland conjugate margins.

vii

Declaration

I hereby declare that this thesis, submitted in fulfilment of the requirements for the

degree of Doctorate of Philosophy, represents my own work. This work has not been

previously submitted to this, or any other institution, for any degree, diploma or other

qualification.

viii

To my mother and father, Sao and Joao

ix

Acknowledgements

In first place I would like to thank my supervisors, Tony Watts and Luis Menezes

Pinheiro, for their supervision, advice and guidance. In particular I would like to thank

Tony for helping me around a few difficult corners this project went through, and thor-

ough reviews of most of this thesis chapters, and to Luis for his great support over the

past 8 years we have worked together, continuous discussion on issues related to the West

Iberia Margin and help accessing different type of datasets, and thorough reviews of some

of this thesis chapters.

Financially, this work was supported by a PRAXIS scolarship from the Fundacao

para a Ciencia e Tecnologia, Portugal (Ref: SFRH/BD/5207/2001) and by an advance

studies scolarship from Fundacao Gulbenkian, Portugal (Num. 72515). Wolfson College

also provided financial support for some conference expenses.

I would like to thank TGS-Nopec for accessing their 2-D multi-channel seismic sur-

vey offshore Western Portugal, and in particular to Reidun Myklebust for the seismic

processing and final images of several sections in the SW Portuguese Margin. I would

also like to thank Dra. Teresinha Abecassis and Carlos Moita, from the Nucleo para

a Prospecao e Exploracao de Petroleos em Portugal, for establishing all the necessary

contacts with the TGS-Nopec Group and providing the permises for the interpretation

of the seismic datasets.

I also wish to thank some individuals who helped me at different stages of this project.

Thanks to Marta Perez-Gussinye for her friendly support, many fruitfull discussions in

a broad range of themes related to this thesis, and carefull reading of drafts of several

chapters. Thanks to Tiago Alves for providing me with good quality sediment thickness

data and clarifying many aspects of the Lusitanian Basin stratigraphy and evolution.

Thanks to Luis Matias for his continued help, great ideas and enthusiasm, and reading

of drafts of several chapters. To J.C. I am gratefull for his great skills in computer

languages and availability to solve any queries, at any time. I would also like to thank

Pedro Terrinha for his help in the interpretation of the TGS-Nopec seismic data.

The sediment thickness dataset compiled in this study used some previously existing

compilations and I would like to thank the authors who kindly provided me with the

xdata, namely: S. Dean and T. Minshull, T. Alves, E. Rasmussen and O. Vejbaek, C.

Moita and J. Carvalho.

Images within this thesis have almost entirely been created using the GMT software

of P. Wessel and W. Smith (http://www.soest.hawaii.edu/gmt).

I would like to thank my fellow members of the Marine Group over the years, Paul

Wyer, Mohammed Ali, Marta Perez-Gussinye, David Close, John Hillier, Tom Jordan,

Natalie Lane, Matt Rodger, and Oigvind Engen, for sharing their ideas on a broad range

of geology and geophysical issues, proof-reading of some early drafts of several chapters,

and many relaxing RA lunches.

I am especially gratefull to Natalie Lane who, during this last year, read thoroughly

and helped me improving the english of at least half of the thesis chapters.

I would also like to thank the staff of the Department of Earth Sciences for helping

at every point possible during this project. In particular I would like to thank Jenny

Colls, Yasuko Nakajima and Emma Brown, for not letting me loosing any extra time with

library, financial and administrative issues, and to Steve Usher and May Chung for their

continued technical support in keeping the computers up and running. Steve, thanks

for prospecting in the meanders of my laptop memory after I had accidently erased the

entire home directory. For a couple of hours my thesis was lost, as far as I could imagine.

On a personal note, I would like to thank my friends at Oxford, Bristol and Nothing-

ham for making my time here enjoyable, and in particular to the people of the Oxford

University Handball Club (OUHC) for the excellent friendship environment and some

travelling around the U.K., which eventually ended up somewhere in Finland.

Thanks also to all my friends and family back in Portugal, who have been supportive

and patient during this long thesis experience. Its great to be back

Lastly, but by no means not least, I want to thank my beautiful small family, Mar-

ianne, Clara and Edgar. We all had to go through a lot of extenuating changes and

difficult periods of sepparation... but gravity kept us together. My dear Marianne, I

do appologise and infinitly apprecite your patience during these last years of my Ph D.,

especially when trying to combine being a student, a husband, a father and an handball

player and coach at the same time. I recognise it was all a bit too much, and I would not

have done it without your courage and strength. This thesis is also dedicated to you.

Contents

Abstract i

Extended Abstract ii

Declaration vii

Dedicate viii

Acknowledgements ix

Contents xi

List of Figures xxi

List of Tables xxi

1 Introduction 1

1.1 Passive Continental Margins . . . . . . . . . . . . . . . . . . . . . . . . . . 1

1.1.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1

1.1.2 Kinematic Models of Rift Type Basins . . . . . . . . . . . . . . . . 5

1.1.3 Mechanics of Rifting and Modes of Lithosheric Extension . . . . 7

1.1.4 The No Magma Paradox . . . . . . . . . . . . . . . . . . . . . . . . 11

1.2 Isostasy, Flexure and Gravity Modelling . . . . . . . . . . . . . . . . . . . 11

1.2.1 Introduction to Isostasy . . . . . . . . . . . . . . . . . . . . . . . . 11

1.2.2 Flexure and Te . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 13

1.2.3 The Use of Gravity in the Study of Passive Continental Margins . 17

1.3 Thesis Outline . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 21

2 Methods 23

2.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23

2.2 Backstripping, Flexure, Subsidence and Gravity Modelling . . . . . . . . . 24

2.2.1 1-D Well Backstripping . . . . . . . . . . . . . . . . . . . . . . . . 24

2.2.2 Flexural Methods . . . . . . . . . . . . . . . . . . . . . . . . . . . . 27

xi

Contents xii

2.2.3 Tectonic Subsidence and Stretching Factor . . . . . . . . . . . . . 30

2.2.4 Calculation of Gravity Anomalies . . . . . . . . . . . . . . . . . . . 37

2.2.5 Process Oriented Gravity Modelling . . . . . . . . . . . . . . . . . 38

2.3 Thermo-Mechanical Rheological Models: Te Structure and Stratigraphy

of Rift Basins . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 41

2.3.1 Refinements to the McKenzie Rift Model . . . . . . . . . . . . . . 41

2.3.2 Analytical Calculation of Te . . . . . . . . . . . . . . . . . . . . . . 46

2.3.3 Model Setup and Parameterisation . . . . . . . . . . . . . . . . . . 51

2.3.4 Subsidence, Flexure and Stratigraphic Modelling: The Steers

Head Basin . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 53

2.3.5 Subsidence, Paleowater Depths and Stratigraphy at Passive Con-

tinental Margins . . . . . . . . . . . . . . . . . . . . . . . . . . . . 57

2.3.6 Multi-Stage Continental Rifting . . . . . . . . . . . . . . . . . . . . 62

2.3.7 Modelling Strategy and Aims . . . . . . . . . . . . . . . . . . . . . 64

3 West Iberia Margin and Data Compilations 66

3.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 66

3.2 Geological and Geophysical Setting of the WIM . . . . . . . . . . . . . . . 67

3.2.1 Plate Tectonic Setting and Seismicity . . . . . . . . . . . . . . . . 67

3.2.2 The Variscan Orogen . . . . . . . . . . . . . . . . . . . . . . . . . . 69

3.2.3 Margin Physiography, Segmentation and Structural Framework . . 71

3.2.4 Stratigraphy and Rifting History of the Margin . . . . . . . . . . . 76

3.2.5 Crustal Structure and Style of Rifting . . . . . . . . . . . . . . . . 79

3.2.6 The Ocean-Continent Transition (Zone), Magnetic Anomalies and

the Onset of Seafloor Spreading . . . . . . . . . . . . . . . . . . . . 86

3.2.7 Free-air and Isostatic Gravity Anomalies at the WIM . . . . . . . 91

3.3 Bathymetric Compilation . . . . . . . . . . . . . . . . . . . . . . . . . . . 94

3.3.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 94

3.3.2 Data Sources . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 96

3.3.3 Cross-Over Error Analysis of Ship-track Data . . . . . . . . . . . . 98

3.3.4 Data Gridding, Statistical Distribution and Spectral Analysis . . . 102

3.4 Sediment Thickness Compilation . . . . . . . . . . . . . . . . . . . . . . . 103

3.4.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 103

3.4.2 Data Sources and Stratigraphic Correlations . . . . . . . . . . . . 104

3.4.3 Sediment Velocity Structure in the West Iberia Margin . . . . . . . 107

3.4.4 Compiled Structural Maps . . . . . . . . . . . . . . . . . . . . . . . 112

Contents xiii

4 Well Backstripping and Subsidence Analysis 117

4.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 117

4.2 Structural Setting and Stratigraphy of the Shelf and Onshore Rift Basins 118

4.2.1 Structural Setting . . . . . . . . . . . . . . . . . . . . . . . . . . . 118

4.2.2 Stratigraphy and Basin Evolution . . . . . . . . . . . . . . . . . . 119

4.3 Backstripping . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 127

4.3.1 Horizon Tops, Ages and Paleowater Depths . . . . . . . . . . . . . 128

4.3.2 Porosity-Vs-Depth and Compaction/Decompaction . . . . . . . . . 130

4.3.3 Sea-level . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 135

4.4 Model Parameterisation and Sensitivity Analysis . . . . . . . . . . . . . . 138

4.5 Subsidence Analysis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 143

4.6 Multi-Phased Continental Rifting and Erosion . . . . . . . . . . . . . . . . 153

4.7 Summary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 161

5 Gravity Anomalies and Flexure in the West Iberia Margin 163

5.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 163

5.2 Data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 164

5.3 Process-Oriented Gravity Modelling, POGM . . . . . . . . . . . . . . . . 167

5.3.1 Methodology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 167

5.3.2 Results Along Selected Profiles . . . . . . . . . . . . . . . . . . . . 175

5.3.3 Summary of Results and Analysis of Residuals . . . . . . . . . . . 184

5.4 Depth of Necking and Strength During Rifting . . . . . . . . . . . . . . . 191

5.5 Underplating . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 200

5.6 Thrust Loading in the Southwest Portuguese Margin . . . . . . . . . . . . 203

5.7 Contributions from Thrust Loading and Upper Mantle Serpentinization

and Depletion in the Southern Iberia Abyssal Plain . . . . . . . . . . . . . 211

5.8 Summary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 221

6 Thermo-Mechanical Modelling and Predicted Stratigraphy 222

6.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 222

6.2 Margin Segmentation and Stretching Factors . . . . . . . . . . . . . . . . 223

6.3 Modelling . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 226

6.3.1 Modelling Strategy and Parameterisation . . . . . . . . . . . . . . 227

6.3.2 Margin Stratigraphy, Stretching Factors and Subsidence . . . . . . 230

6.3.3 Thermal Structure, YSEs and Strength of the Lithosphere . . . . 234

6.3.4 Backstripping and Paleowater depths . . . . . . . . . . . . . . . . . 237

6.3.5 Predicted Stratigraphy and Mechanical Structure Along Profile 7 . 241

Contents xiv

6.4 Predicted Stratigraphy and Mechanical Structure Along Profiles 2 and 8 . 250

6.4.1 Central Portuguese Margin: Results along Profile 8 . . . . . . . . . 251

6.4.2 Galicia Margin: Results Along Profile 2 . . . . . . . . . . . . . . . 254

6.5 Discussion of Results . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 258

6.5.1 Margin Subsidence/Uplift History . . . . . . . . . . . . . . . . . . 258

6.5.2 Paleowater Depths and Margin Evolution . . . . . . . . . . . . . . 261

6.5.3 Thermo-mechanical Structure, YSEs and Depths of Necking . . . 261

6.6 Summary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 265

7 The West Iberia-Newfoundland Conjugate Margins 267

7.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 267

7.2 Central North Atlantic Margins and Plate Tectonic Setting . . . . . . . . 268

7.2.1 Present Setting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 268

7.2.2 Rotation Poles, Coordinate Systems and Plate Reconstructions . . 270

7.2.3 West Iberia-Newfoundland Reconstructions at M0 . . . . . . . . . 272

7.3 Overview of the West Iberia-Newfoundland Rift . . . . . . . . . . . . . . . 276

7.3.1 Regional Setting, Structural and Stratigraphic Framework, and

Margins Segmentation . . . . . . . . . . . . . . . . . . . . . . . . . 276

7.3.2 Crustal Structure, Velocity Models and Constraints on the Style

and Geometry of Rifting . . . . . . . . . . . . . . . . . . . . . . . . 281

7.3.3 Sediments, Backstripping and Subsidence Analysis . . . . . . . . . 289

7.4 Results from POGM on Conjugate Profiles . . . . . . . . . . . . . . . . . 292

7.4.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 292

7.4.2 The Galicia and Flemish Cap margins . . . . . . . . . . . . . . . . 294

7.4.3 The Intermediate Segment . . . . . . . . . . . . . . . . . . . . . . . 296

7.4.4 The North Newfoundland Basin and Central Portuguese Margin . 298

7.4.5 The South Newfoundland Basin and Southwest Portuguese Margin 300

7.4.6 Summary and Analysis of Results . . . . . . . . . . . . . . . . . . 300

7.5 Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 302

7.5.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 302

7.5.2 Dynamic topography . . . . . . . . . . . . . . . . . . . . . . . . . . 304

7.5.3 Differential Crustal Stretching and Lower Crustal Flow . . . . . . 305

7.5.4 Flexural Stresses and Isostatic Rebound . . . . . . . . . . . . . . . 307

7.5.5 Magmatism and Mantle Serpentinization . . . . . . . . . . . . . . 308

7.5.6 Lithospheric-Scale Simple Shear . . . . . . . . . . . . . . . . . . . 309

7.5.7 Rift Width, Symmetry and Modes of Lithospheric Extension . . 318

7.6 Summary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 320

Contents xv

8 Conclusions and Future Work 322

8.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 322

8.2 Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 323

8.3 Future Work . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 327

A Thermal Structure of Extended Continental Lithosphere 329

A.1 1-D Heat Flow . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 329

A.2 2-D Heat Flow . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 331

B User Guide to the Mrift code 333

C Final Gravity and Crustal Models 336

Bibliography 344

List of Figures

1.1 Worldwide distribution of volcanic and magma-poor passive continental

margins . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2

1.2 Typical structure of volcanic and magma-poor rift margins . . . . . . . . 4

1.3 McKenzie [1978] pure shear model of lithospheric extension . . . . . . . . 6

1.4 Pure shear and simple shear models of lithospheric extension . . . . . . . 7

1.5 Numerical model predictions on the modes of lithospheric extension . . . 9

1.6 Comparison of rift geometries obtained from analogue sandbox modelling 10

1.7 The No Magma Paradox in passive continental margins . . . . . . . . . . 12

1.8 Airy and Pratt models of isostasy . . . . . . . . . . . . . . . . . . . . . . . 13

1.9 Local versus regional models of isostasy . . . . . . . . . . . . . . . . . . . 14

1.10 Te versus age of the oceanic lithosphere at the time loading . . . . . . . . 15

1.11 Predicted and observed stratigraphy at rift basins: elastic versus viscoelas-

tic plate models . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16

1.12 Sediments and gravity anomaly data at passive continental margins . . . . 19

1.13 Te estimates at passive continental margins . . . . . . . . . . . . . . . . . 20

2.1 Rational of the 1-D backstripping technique . . . . . . . . . . . . . . . . . 26

2.2 Multilayer 1-D backstripping . . . . . . . . . . . . . . . . . . . . . . . . . 27

2.3 Flexural response function, (k) . . . . . . . . . . . . . . . . . . . . . . . 29

2.4 Subsidence and thermal structure of the lithosphere: the McKenzie [1978]

two-stage rift model . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 32

2.5 Introduction to the concept of depth of necking (Zneck) . . . . . . . . . . 35

2.6 2-D gravity calculations . . . . . . . . . . . . . . . . . . . . . . . . . . . . 38

2.7 3-D gravity calculations . . . . . . . . . . . . . . . . . . . . . . . . . . . . 39

2.8 Process Oriented Gravity Modelling . . . . . . . . . . . . . . . . . . . . . 40

2.9 Predicted subsidence in rift type basins: effects of lateral heat flow and

finite rifting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 43

2.10 Predicted subsidence for single and multi-stage continental rift scenarios . 45

2.11 The yield strength envelope; YSE . . . . . . . . . . . . . . . . . . . . . . . 48

2.12 Fiber stresses in a thin bending plate and the YSE failure criterion . . . . 49

xvi

Contents xvii

2.13 Sensitivity of the calculated Tes to variations in the crust rheology, strain

rates and radiogenic heat production . . . . . . . . . . . . . . . . . . . . . 50

2.14 Temperature structure, subsidence, Te and flexure at rift basins . . . . . . 55

2.15 The steers head basin: Tes, strain rates and predicted stratigraphy . . . 56

2.16 Predicted stratigraphy and calculated Tes for different thermal and rheo-

logical modelling settings . . . . . . . . . . . . . . . . . . . . . . . . . . . 58

2.17 Subsidence and paleowater depths in passive continental margins . . . . . 59

2.18 Predicted stratigraphy and Tes in passive continental margins . . . . . . 61

2.19 Finite rifting and multi-stage continental rifting passive margin scenarios:

Predicted stratigraphy and Tes . . . . . . . . . . . . . . . . . . . . . . . . 63

3.1 Plate tectonic setting of the Iberian Plate since continental break-up . . . 68

3.2 Physiography of the West Iberia Margin (WIM), sedimentary basins and

structural elements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 71

3.3 Schematic stratigraphic sections along the WIM . . . . . . . . . . . . . . . 74

3.4 Stratigraphy of the main sedimentary basins in the WIM . . . . . . . . . 78

3.5 Low-angle detachments and geometry of the West Iberia-Newfoundland rift 81

3.6 Velocity models from three wide-angle/refraction seismic sections . . . . . 85

3.7 Magnetic anomalies off West Iberia . . . . . . . . . . . . . . . . . . . . . . 89

3.8 Free-air gravity and isostatic anomalies at the West Iberia Margin . . . . 92

3.9 Statistical distribution of bathymetric values . . . . . . . . . . . . . . . . 95

3.10 Shipboard bathymetric data . . . . . . . . . . . . . . . . . . . . . . . . . . 97

3.11 Analysis of shipboard bathymetric data . . . . . . . . . . . . . . . . . . . 100

3.12 Statistics of cross-over errors . . . . . . . . . . . . . . . . . . . . . . . . . 101

3.13 The new WIM bathymetry: Statistical distribution and spectral analysis . 103

3.14 Sediment thickness data in the WIM . . . . . . . . . . . . . . . . . . . . . 106

3.15 Sediment velocities in the WIM and TWTT-velocity functions . . . . . . 109

3.16 Sediment velocity structure in the WIM . . . . . . . . . . . . . . . . . . . 111

3.17 Compiled structural maps . . . . . . . . . . . . . . . . . . . . . . . . . . . 113

3.18 Compiled sediment thickness grid for the WIM and comparisons with ex-

isting data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 115

4.1 Sedimentary basins, main structural elements and commercial wells along

the Portuguese Margin . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 119

4.2 Stratigraphy of the onshore and continental shelf basins . . . . . . . . . . 122

4.3 Synthetic stratigraphic profile along the Portuguese continental shelf . . . 125

4.4 Depositional environments in the Lusitanian and Porto basins . . . . . . . 129

Contents xviii

4.5 Porosity data and lithologies from commercial wells . . . . . . . . . . . . . 133

4.6 Porosity-depth decompaction curves . . . . . . . . . . . . . . . . . . . . . 135

4.7 Relative sea-level changes in the geological past . . . . . . . . . . . . . . . 137

4.8 Well backstripping sensitivity analysis . . . . . . . . . . . . . . . . . . . . 141

4.9 Results from well backstripping in the Porto Basin . . . . . . . . . . . . . 145

4.10 Results from well backstripping in the Northern Lusitanian Basin . . . . . 147

4.11 Results from well backstripping in the Southern Lusitanian Basin . . . . . 150

4.12 Summary of results from well backstripping . . . . . . . . . . . . . . . . . 152

4.13 Multi-stage continental rifting in wells Lu-1 and Pe-1 . . . . . . . . . . . . 154

4.14 Extension factors and duration of the Late Triassic-Early Jurassic rifting

event from RMS analysis in 7 boreholes . . . . . . . . . . . . . . . . . . . 157

4.15 Erosion estimates in the Northern Lusitanian Basin . . . . . . . . . . . . . 159

5.1 Bathymetry of the WIM, seismic and well data, and location of the profiles

used for gravity modelling . . . . . . . . . . . . . . . . . . . . . . . . . . . 165

5.2 Long-wavelength gravity field . . . . . . . . . . . . . . . . . . . . . . . . . 167

5.3 Process-Oriented Gravity Modelling, POGM, flow chart . . . . . . . . . . 168

5.4 Gravity anomalies, flexure and geometry of the rift margin . . . . . . . . . 170

5.5 Comparison between observed and modelled gravity for various Te(sed)

scenarios . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 172

5.6 Calculated gravity anomalies, residuals and crustal model for a laterally

varying Te(sed) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 173

5.7 Sensitivity of POGM to variations in Te(sed), s, m and tc . . . . . . . . 174

5.8 Results from POGM in the Galicia Margin, Profiles 1-4 . . . . . . . . . . 177

5.9 Results from POGM in the Intermediate Segment, Profiles 4-7 . . . . . . 179

5.10 Results from POGM in the Central Portuguese Margin, Profiles 8-10 . . . 181

5.11 Results from POGM in the Southwest Portuguese Margin, Profiles 11-14 . 183

5.12 Airy and flexural isostatic anomalies along Profiles 1 to 7 . . . . . . . . . 186

5.13 Airy and flexural isostatic anomalies along Profiles 8 to 14 . . . . . . . . . 188

5.14 Gravity anomalies and crustal structure along three wide-angle/refraction

profiles . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 190

5.15 Effects of varying Zneck, Te(rift) and Te(sed) on the calculated gravity

anomalies, isostatic residuals and rift geometry for a typical WIM transect 193

5.16 Sensitivity of the isostatic anomaly to Te(rift) and Zneck . . . . . . . . . . 194

5.17 Finite strength during rifting; synthesis of results along the WIM . . . . . 196

5.18 Underplating in the Central Portuguese Margin; Line IAM-9 . . . . . . . 202

5.19 Structural elements in the SW Portuguese Margin . . . . . . . . . . . . . 204

Contents xix

5.20 Gravity anomalies from thrust loading and associated flexure . . . . . . . 207

5.21 Isostatic anomalies power spectra in the SW Portuguese Margin . . . . . 210

5.22 FAA and tectonic setting of the Southern Iberia Abyssal Plain . . . . . . 212

5.23 Thrust loading, flexure and gravity anomalies in the Southern Iberia Abyssal

Plain . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 214

5.24 Relationships between partial melting, mantle density and thickness of

depleted upper mantle . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 217

5.25 Gravity signature from upper mantle serpentinization and depletion . . . 219

6.1 Stretching factors along the West Iberia Margin . . . . . . . . . . . . . . . 224

6.2 Flow chart of the multi-rift thermo-mechanical modelling procedure . . . 227

6.3 Margin stratigraphy, stretching factors and predicted subsidence along

Profile 7 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 232

6.4 Predicted thermal structure, YSE, Te and strain rates . . . . . . . . . . . 234

6.5 Cumulative backstrip, total tectonic subsidence and paleowater depths . . 239

6.6 Paleowater depths along Profile 7 and associated uncertainties . . . . . . . 241

6.7 Paleowater depths, predicted stratigraphy and mechanical structure along

Profile 7 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 243

6.8 Effects of varying the paleowater depths on the margin stratigraphy . . . 245

6.9 Tests on the computed margin stratigraphy . . . . . . . . . . . . . . . . . 246

6.10 Paleowater depths, thermo-mechanical evolution and preferred stratigraphic

model along Profile 7 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 248

6.11 Paleowater depths, thermo-mechanical evolution and predicted stratigra-

phy along Profile 8 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 252

6.12 Paleowater depths, thermo-mechanical evolution and predicted stratigra-

phy along Profile 8 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 256

6.13 Comparison between 1-D well backstripping and 2-D flexural models pre-

dicted subsidence . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 259

6.14 Stretching factors, thermal structure and YSEs . . . . . . . . . . . . . . . 263

7.1 North Atlantic bathymetry and the conjugate margins of West Iberia and

Newfoundland . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 269

7.2 Geocentric versus geographic coordinates in pole rotations . . . . . . . . . 271

7.3 Central North Atlantic plate reconstructions at the time of M0 . . . . . . 274

7.4 Bathymetry, tectonic subsidence, free-air and isostatic anomalies at the

West Iberia-Newfoundland conjugate margins . . . . . . . . . . . . . . . . 278

7.5 Conjugate velocity models, free-air and isostatic anomalies . . . . . . . . . 284

Contents xx

7.6 Basin profiles and total tectonic subsidence . . . . . . . . . . . . . . . . . 290

7.7 Results from POGM on conjugate profiles: Lines SCREECH-1 and ISE-1 295

7.8 Results from POGM on conjugate profiles: Line SCREECH-2 and Profile 7 297

7.9 Results from POGM on conjugate profiles: Lines Lithoprobe 85-4, SCREECH-

3 and IAM-9 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 299

7.10 Results from POGM on conjugate profiles: Lines Lithoprobe 85-2 and

IAM-5 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 301

7.11 Rifting geometry, crustal extension and load anomalies along seismic pro-

files SCREECH-3 and IAM-9 . . . . . . . . . . . . . . . . . . . . . . . . . 306

7.12 Subsidence and thermal models along the SCREECH-3 and IAM-9 profiles 310

7.13 Simple shear model proposed to explain the formation of the West Iberia

and Newfoundland conjugate margins . . . . . . . . . . . . . . . . . . . . 314

A.1 Model proposed by McKenzie [1978] for the formation and evolution of

rift type basins . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 330

C.1 Process-Oriented Gravity Modelling results along Profile 1 . . . . . . . . . 337









C.10 Process-Oriented Gravity Modelling results along Profile 10 . . . . . . . . 341





List of Tables

2.1 Reference parameters for the thermo-mechanical models . . . . . . . . . . 52

2.2 Rheological parameterisation of the crust and lithospheric mantle . . . . . 53

3.1 Seismic data used to constrain three velocity models along the WIM . . . 83

3.2 Data from previous works used in the new sediment thickness compilation 105

4.1 Parameterisation of the thermal model . . . . . . . . . . . . . . . . . . . . 139

5.1 Segments defined along the WIM for the analysis of the gravity modelling

results . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 175

6.1 Rheological parameterisation of the crust and lithospheric mantle . . . . . 229

7.1 Reconstruction poles for the African, Iberian and Eurasian plates relative

to North America . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 273

7.2 Crustal sections constrained from seismic data and gravity modelling along

the West Iberia and Newfoundland margins . . . . . . . . . . . . . . . . . 282

xxi

Chapter 1

Introduction

1.1 Passive Continental Margins

1.1.1 Introduction

Passive continental margins are global scale features that mark the transition between

oceans and continents without an intervening plate boundary. They are the most widesp-

read manifestation of Mesozoic and Cenozoic extensional deformation in continental litho-

sphere and, on the present day Earth, they bound nearly all continental masses which

border the Atlantic, Indian, Southern and Arctic oceans (Figure 1.1). Passive margins,

also known as divergent, rift or Atlantic type margins, develop by thinning of the con-

tinental lithosphere, followed by continental break-up and the formation of new oceanic

basins, under an overall tensional stress regime that is generally associated with the de-

velopment of normal faulting, magma generation and emplacement, vertical movements

of uplift and subsidence, erosion and sediment deposition (Sleep [1971]; Walcott [1972];

Watts & Ryan [1976]; Montadert et al. [1977]; de Charpal et al. [1978]; McKenzie [1978];

White et al. [1987]; Uchupi & Emery [1991]; Eldholm et al. [1995]).

Extreme lithospheric thinning and continental break-up can lead to the formation

of a wide range of passive margin settings. A common distinction is based upon the

amount of magmatism during rifting and the early stages of seafloor spreading. The two

end members of this classification are the volcanic and non-volcanic (or magma-poor)

passive continental margins. Although some authors consider the possibility of a wide

spectrum of magmatic activity during rifting, with the existence of intermediate volcanic

passive margins (e.g. Symonds et al. [2000]) and transitional passive margin segments

(Dehler et al. [2003]), others highlight the fact that they represent distinct processes of

continental break-up (e.g. Geoffrey [2005]).

Volcanic margins are commonly associated with anomalously hot mantle conditions

(100-250C; Skogseid [2001]); i.e. where heat has been advected at the base of the

lithosphere (White et al. [1987]; White & McKenzie [1989]; Hill [1994]). They are clearly

dominant, for example, in the eastern North Atlantic and South Atlantic Oceans, and

along the western Indian Coast (Figure 1.1), where their formation has been associated

1

Chapter 1. Introduction 2

Figure 1.1: Worldwide distribution of volcanic (pink) and magma-poor (blue) passive

continental margins. Modified from Menzies et al. [2002]. Volcanic passive margins

where sequences of intrabasement seaward dipping reflectors (SDRs) have been identified

are marked by orange dots (according to Coffin & Eldholm [1994]). The dashed grey

rectangles highlight the margins studied in this thesis, namely, the magma-poor passive

margins of West Iberia and Newfoundland, in the eastern and western central North

Atlantic, respectively.

with the activity of mantle plumes (Hill [1994]; Courtillot et al. [1999]; Skogseid [2001]

and references therein). Globally, volcanic margins may represent up to 90% of the oceans

rifted margins (pink lines in Figure 1.1; Menzies et al. [2002]).

One of the most distinctive features of these margins is the existence of a thick (>

3 km) zone of extrusive basaltic complexes, characterized by a smooth opaque surface

and internal wedges of seaward dipping reflectors (e.g. Eldholm et al. [1995]; the SDRs

in Figure 1.2a). These complexes resulted from the emplacement of multiple sub-aerial

basalt flows interbedded with thin sediment layers, that were subsequently rotated and

tilted as the margin subsided (Skogseid & Eldholm [1987]). This type of sequence was

first identified by Hinz [1981]), and later confirmed by drilling in a number of North

Atlantic margins (Coffin & Eldholm [1994]; Eldholm et al. [1995]; Menzies et al. [2002]

and references therein). Other geological features commonly observed at rifted volcanic

margins include (e.g. White et al. [1987]; Coffin & Eldholm [1994]; Geoffrey [2005]):

(1) a thick, wedge shape, underplated lower crustal body (up to 15 km thick) of high

P-wave velocities (> 7.3 km s1; Figure 1.2a); (2) sills and low-angle dikes within pre-

rift sediments; (3) thicker than normal oceanic crust adjacent to the ocean-continent

boundary (OCB); (4) and, in some cases, onshore continental flood basalts.

Magma-poor passive continental margins, on the other hand, form away from the

influence of hotspot activity. Due to the lack of significant magmatism, either extruded


at the surface or underplated under the crust, these margins are ideal settings to study

the mechanisms of rifting and continental break-up, particularly when the basement is

not obscured by a thick sediment coverage.

In the late 70s, Montadert et al. [1977] and de Charpal et al. [1978] observed a

series of tilted and rotated blocks, bounded by listric normal faults, in the Bay of Biscay

and Rockall Plateau margins. These authors postulated that lithospheric stretching was

accommodated by brittle failure within the upper crust and ductile flow in the underlying

lower crust and mantle. In the Bay of Biscay Margin, de Charpal et al. [1978] identified a

strong, approximately horizontal seismic reflector lying at the base of the normal faults,

the S reflector, and interpreted it as marking the brittle-ductile transition.

Since then, the deep offshore of a number of non-volcanic passive continental margins

has been investigated by several academic and government-led groups. These include

(see Figure 1.1): the margins of West Iberia (Boillot et al. [1988b]; Whitmarsh & Swayer

[1996]), Newfoundland (Reid [1994]; Hopper et al. [2006]), most of the Labrador Sea

(Chian & Louden [1994]; Chian et al. [1995]), the northern Red Sea, part of the southern

and NW Australia (Fynlayson et al. [1999]; Driscoll & Karner [1998]), the Gulf of Aden

(dAcremont et al. [2005]), and the NW Moroccan margin (Contrucci et al. [2004]).

Although sequences of rotated faulted blocks have been clearly imaged in some of

these settings, such as in the NW Portugal (Figure 1.2b) and Galicia margins (e.g. Sibuet

[1992]; Krawczyk et al. [1996]), others show no significant upper crustal faulting between

relatively undeformed basement, in the shelf and onshore, and highly stretched continen-

tal crust, in the continental slope and rise (e.g. Reid [1994]; Chian et al. [1995]; Dean

et al. [2000]; Shillington et al. [2006]). Moreover, it appears that in most passive con-

tinental margins the amount of stretching measured from faulting in the upper crust is

much lower than that obtained for the whole crust and lithosphere from subsidence and

gravity modelling (e.g. Driscoll & Karner [1998]; Davis & Kusznir [2004]).

Another relevant observation from non-volcanic passive continental margins is the

existence of a region of transition between typical continental and unambiguous oceanic

type crust, the Ocean-Continent Transition (Zone), OCTZ. In some places, along the

West Iberia and Newfoundland margins, an OCTZ of over 100 km width has been sug-

gested (e.g. Pickup et al. [1996]; Dean et al. [2000]; Sibuet et al. [2007]). Based on the

reflectivity of the basement, the seismic velocities inferred from wide-angle and refraction

data, the magnetic signature, and on samples recovered from a number of basement highs

during 3 West Iberia and 1 Newfoundland Ocean Drilling Program (ODP) legs, this re-

gion is commonly interpreted as consisting of exhumed and serpentinized sub-continental

lithospheric mantle (Boillot et al. [1988a]; Whitmarsh et al. [1990]; Pinheiro et al. [1992];

Reid [1994]; Pickup et al. [1996]; Dean et al. [2000]; Whitmarsh & Wallace [2001]; Lau


Figure 1.2: Typical structure of volcanic (a) and magma-poor (b) rift margins. a) The

Vring volcanic margin, off Norway (modified from Skogseid & Eldholm [1995]), shows

the typical SDR sequences and the High-Velocity lower crustal Body (HVB). b) The

non-volcanic margin example is from NW Portugal (modified from Alves et al. [2006]).

It shows the typical rotated faulted blocks sequence (Montadert et al. [1977]) and the

serpentinization of the lithosphere mantle peridotites under highly thinned continental

crust, which will eventually become exhumed as extension proceeds (e.g. Boillot et al.

[1988a]; Sibuet [1992]). Also indicated are the distinct morphological sectors which char-

acterize passive continental margins in general (i.e. the continental shelf, slope, rise and

abyssal plain; see text).

et al. [2006a]; Sibuet et al. [2007]). Under the currently accepted rift models, however,

the existence of such a wide region of exhumed mantle without significant magmatism is

paradoxal (see discussion in Section 1.1.4). Buck [2004], for example, referred to this as

the No Magma paradox.

Despite the large differences in the amount of magmatism and in the structure of

the break-up zone, most passive continental margins can be described in terms of similar

conventional morphological features (Figure 1.2). These include (after Heezen [1974]):

- The continental shelf, characterized by a gently dipping slope between the coast

line and the shelf break. The depth of the shelf break in passive continental margins

varies in function of the latitude. In the West Iberia Margin, for example, the shelf break

varies narrowly between 160 and 200 m (Vanney & Mougenot [1981]). In some passive


continental margins, a wide coastal plain (e.g. east coast U.S.A.; e.g. Grow & Sheridan

[1988]) is found landward of the coast line, with slopes similar to those found in the

continental shelf.

- The continental slope, defined between the shelf and the continental rise slope

breaks. Continental slope gradients in passive continental margins can reach up to 20%;

- The continental rise, a region gently dipping oceanwards (< 1%) that can extend for

more than 100 km, overlying extended continental, transitional and oceanic type crust;

- The abyssal plain, where the regional gradients are inverted, towards the conti-

nents, and typically do not exceed 0.1%. The abyssal plains cover most of the worlds

oceans, where seamounts, isolated or occurring in groups, locally disrupt the otherwise

monotonous relief.

1.1.2 Kinematic Models of Rift Type Basins

The existence of large thicknesses of shallow water sediments at passive continental mar-

gins, and rift basins in general, was only recently understood. Sleep [1971], for example,

proposed that the tectonic subsidence of continental margins was due to crustal thinning,

after thermal expansion, uplift and erosion, followed by thermal contraction of the litho-

sphere. The thermal contraction model of Sleep [1971] is consistent with the observations

at a number of boreholes along the East Coast, U.S.A., where the inferred tectonic sub-

sidence decreases exponentially with time following rifting (Sleep [1971]; Sleep & Snell

[1976]; Watts & Ryan [1976]), in a manner similar to that of the oceanic basement away

from the mid-ocean ridges (Parsons & Sclater [1977]). It presented, however, two major

difficulties. First, it did not explain the initial heating of the continental lithosphere.

Second, the subsidence of the margin beyond the initial, pre-deformation reference level

assumed large amounts of sub-aerial basement erosion, which was not supported by the

data.

McKenzie [1978] proposed an alternative, two-stage rift model to explain the devel-

opment and evolution of sedimentary basins (Figure 1.3). According to this model, the

lithosphere is in a first stage instantaneously stretched, producing crustal and lithospheric

mantle thinning and passive upwelling of hot asthenosphere (Figure 1.3b). This stage

is associated with block faulting within the crust and rapid subsidence due to isostatic

re-equilibration of the thinned crust. In a second stage, as heat is lost through vertical

conduction to the surface, and the lithosphere regains its original, steady state thickness,

further subsidence occurs (Figure 1.3c).

The McKenzie pure shear stretching model, and subsequent modifications which

accounted for the effects of lateral heat flow and finite rifting (e.g. Watts et al. [1982];


Figure 1.3: Illustration of the

McKenzie [1978] pure shear

model of lithospheric exten-

sion and thermal evolution

for the formation of rift type

basins. is the stretching fac-

tor, by which the lithospheric

mantle (Lm) and crust (C) are

initially stretched. A is the as-

thenosphere, w is the initial

width of the basin, L is the

initial thickness of the litho-

sphere and t is the time since

rifting.

Cochran [1983]) can explain, in broad terms, the subsidence/uplift history and heat-flow

record, as well as the geometry and thickness of the sediment strata in many rift type

basins, both in passive continental margins and intracontinental rift settings.

One limitation of the McKenzie type pure shear kinematic models is that they predict

highly symmetric rift basins (Figure 1.4a), whilst rifts are often asymmetric (e.g. Wer-

nicke [1981]; Ebinger et al. [1991]; Lister et al. [1991]). Wernicke [1981], and later Wer-

nicke & Burchfiel [1982] and Wernicke [1985], proposed a different conceptual model of

rifting, where large amounts of lithospheric extension are accommodated along low-angle

normal faults (or detachment faults) which cut through the crust and mantle lithosphere

(Figure 1.4b). They invoked the model to explain the formation of metamorphic core

complexes, where low and high grade metamorphic rocks are juxtaposed near the sur-

face, as observed in the Basin and Range Province, Western U.S.A. (e.g. Wernicke [1981];

Armstrong [1982]). This model, commonly known as the simple shear, Wernicke or

the Lithospheric Wedge stretching model (e.g. Lister et al. [1991]), is associated with

lateral variations in the amount of crust and lithospheric stretching and the formation

of an upper plate margin, in the hangingwall of the detachment surface, and a lower

plate margin in the footwall. The subsidence and uplift history of the upper and lower


plate margins predicted by such models will be fundamentally different, with more initial

(i.e. isostatic) subsidence in the lower plate and greater thermal subsidence in the upper

plate margin (Wernicke [1985]).

Figure 1.4: Different models

of lithospheric extension pro-

posed to explain the forma-

tion of rift basins and passive

continental margins. a) Pure

shear. b) Simple shear. c)

Combination of lithospheric-

scale pure shear and simple

shear stretching within the

crust. Modified from Keen

et al. [1989].

A number of rift models which combine the effects of large displacements along litho-

spheric and/or crustal scale penetrative major shear zones, symmetrical extension on the

ductile lower crust and mantle lithosphere, and the isostatic response of the lithosphere

to the asymmetric redistribution of loads during stretching, have been since proposed

(e.g. Coward [1986]; Lister et al. [1986]; Weissel & Karner [1989]; Lister et al. [1991]).

The importance of these models is that they can predict not only the asymmetry of the

rift basins, but also a great variety of features observed in Atlantic type margins and in-

tracontinental rifts, such as: (1) permanently uplifted rift flanks (e.g. Weissel & Karner

[1989]; Ebinger et al. [1991]); (2) large post-rift regional subsidence, in relation to that

predicted from the amount of stretching measured from faulting in the upper-middle

crust (Boillot et al. [1988a]; Driscoll & Karner [1998]; Davis & Kusznir [2004]); and (3)

marginal plateaus (Etheridge et al. [1989]; Lister et al. [1991]).

1.1.3 Mechanics of Rifting and Modes of Lithosheric Extension

The models discussed above provide an essential kinematic framework that explains the

formation and, to a certain extent, the diversity and geometry of rift type basins, based

on simple principles of isostasy and heat conduction. They are, however, limited in the

sense that they do not take into consideration the rheological structure and the strain

distribution within the lithosphere during stretching.

A number of studies into the dynamics of rifting show that different styles, or modes,


of continental extension can be produced by varying the initial thermal and mechanical

structure of the lithosphere (e.g. crustal thickness, steady state geotherm and rheologies

of the crust and mantle) and the rate of lithospheric extension (e.g. England [1983];

Kusznir & Park [1987]; Sonder & England [1989]; Buck [1991]; Bassi et al. [1993]). Buck

[1991], for example, used a thin-sheet model approximation (i.e. where a vertically aver-

aged rheology is determined at every point of the model) to calculate the balance between

the effects of stretching in the intrinsic strength of the lithosphere, and the gravitational

stresses generated by lateral variations in the thickness of the crust during rifting. Ac-

cording to this author, the style and width of the rift basin is mainly determined by the

steady state geotherm and the crustal thickness (Figure 1.5a). Narrow rifts will form

if the lithosphere is initially cold, whereas a wide or core complex (i.e. very wide) rift

mode is favoured for a thick crust and high initial Moho temperatures. As shown in

Figure 1.5a, Bucks model predicts little changes in the style of rifting associated with

variations in the velocity of the rifting process (i.e. the strain rate).

Thin-sheet approximation models, such as those developed by Buck [1991] and others

before him (e.g. England [1983]; Kusznir & Park [1987]), are restricted in that they

do not account for the way strain varies across the lithosphere. Moreover, the analysis

of Buck [1991] only considered a small amount of extension (< 25%), whereas seismic

sections across many passive continental margins show > 250% of crustal stretching over

widths of less than 100 km (e.g. Dean et al. [2000]; Hopper et al. [2006]; Lau et al.

[2006a]).

Bassi et al. [1993], for example, used finite element modelling techniques to investi-

gated the evolution of passive margin settings and the conditions for continental break-up.

In particular, these authors have shown that for realistic rifting velocities and durations,

similar to those inferred at a number of North Atlantic margins, the style of rifting will

depend not only on the initial conditions but also on the strain rates. As depicted in

Figure 1.5b, lower geothermal gradients and harder rheologies (i.e. higher plasticity),

combined with higher rifting velocities, will determine a rapid localization of the strain

near the rift axis, favouring a narrow style of rifting. On the other hand, creep dominant

rheologies and slower rift velocities are associated with a gradual necking of the litho-

sphere and wider rift settings. For a warmer and thinner lithosphere, the models of Bassi

et al. [1993] predict a runaway thinning stretching mode. This mode is essentially

characterized by a lateral migration of the locus of extension caused by the cooling and

associated hardening of the thinned lithosphere (e.g. England [1983]; Kusznir & Park

[1987]; Sonder & England [1989]).

It is important to note, however, that according to the modelling results of Bassi

et al. [1993], and to previously publish dynamic models of rifting (e.g. Bassi [1991];


Figure 1.5: Numerical modelling predictions on the style, or mode, of continental rifting.

a) Results from numerical models which use a thin sheet approximation to determine the

evolution of the lithospheric intrinsic strength and gravitational buoyancy forces during

rifting (after Buck [1991]). b) Finite element models developed to investigate the style

of rifting at passive continental margin settings (after Bassi et al. [1993]).

Lin & Parmentier [1990]), the lithosphere becomes mechanically unstable for stretching

factors () > 3-4 (i.e. 150-200% of extension). This leads to rapid necking and continental

rupture, thus preventing the formation of a very wide region of highly stretched (i.e.

> 3) continental and/or transitional basement.

Over the past decade, both numerical (e.g. Hopper & Buck [1996]; ter Voorde et al.

[1998]; Govers & Wortel [1999]; Burov & Poliakov [2001]; Behn et al. [2002]; Davis &

Kusznir [2002]) and analogue (Brun & Beslier [1996]; Michon & Merle [1998]) models

of rift basin evolution have become increasingly more complex, integrating multi-layer

crustal and mantle rheologies. Of particular interest for this thesis are those studies

which focus on the mechanical effects of coupling-decoupling between crust and mantle

competent layers (ter Voorde et al. [1998]; Burov & Poliakov [2001]), and on the devel-

opment of necking instabilities, which control both the subsidence/uplift history and the

overall style of lithospheric thinning of rift basins (Braun & Beaumont [1989]; Brun &

Beslier [1996]; Govers & Wortel [1999]).


Figure 1.6: Comparison of rift geometries obtained from analogue sandbox modelling.

The experiments predict the formation of two main shear zones, in the ductile portions of

the crust and mantle (dark grey areas). In experiment (a) the extension rates are small,

and the model predicts a narrow, essentially asymmetric rift, with the development of

a lithospheric-scale detachment structure; Experiment (b) uses higher extension rates,

which results in a wider rift basin and the development of crustal-scale detachments

faults, superimposed on a lithospheric-scale symmetric rift.

Brun & Beslier [1996], for example, using a sandbox analogue modelling apparatus,

showed that by applying a constant rate of deformation at the edges of the model,

a combination of pure (regional necking) and simple shear (grabens and half-grabens)

related structures would be generated within the same rift system. Brun & Beslier [1996],

and later Michon & Merle [1998], further suggest that both the geometry and width of

the rift vary as a function of the extension rate, with higher extension rates favouring the

development of wider and more symmetric rifts (Figure 1.6). These results are, therefore,

at odds with the predictions from the numerical modelling experiments of Bassi et al.

[1993] discussed above.

Govers & Wortel [1999], on the other hand, using finite element modelling techniques,

argued that the depth at which lithospheric necking occurs (Zneck), which corresponds

to levels of maximum strength in the crust and mantle (e.g. Braun & Beaumont [1989]),

is not fixed, as commonly assumed in kinematical models of rift basin formation (e.g.

Weissel & Karner [1989]; Watts & Stewart [1998]), but evolves both in time and space.

However, according to their modelling predictions Zneck becomes stationary after rela-

tively small amounts of extension, < 1.5. In most passive continental margins, this

amount of stretching is typical for the upper continental slope, and reached within the

early stages of rifting. Therefore, the Znecks inferred from the kinematic models, based

on the stratigraphy, subsidence and gravity anomaly data, probably reflect an average

mechanical structure of the lithosphere during the rifting process, as suggested by Keen

& Dehler [1997].


1.1.4 The No Magma Paradox

One observation which is difficult to explain in light of models which assume passive

rifting and necking (e.g. McKenzie type rift models), is the existence of wide regions

of highly extended continental and/or transitional basement where, in places, the upper

lithospheric mantle has been exhumed at the surface, without significant amounts of

melting (e.g. Reid [1994]; Whitmarsh & Wallace [2001]; Desmurs et al. [2004]). According

to predictions from simple thermal models, more than 4 km of melt would be expected

at the Ocean-Continent Transition (Zone), OCTZ, from adiabatic upwelling of normal

1300 C asthenosphere and a > 10 (Bown & White [1995]; Minshull et al. [2001]).

Minshull et al. [2001], and more recently Reston & Morgan [2005], based on numerical

models of rifting which assume depth uniform stretching (i.e. pure shear), have shown

that in order to produce less than 2 km melt, as inferred for the WIM (Whitmarsh

et al. [2001b]; Minshull et al. [2001]), the mantle potential temperature needed to be

lowered by 50-100C (Figure 1.7a). Perez-Gussinye et al. [2006], using improved mod-

elling techniques, suggested that very slow rifting velocities (< 6 mm yr1) and/or highly

depleted sub-continental lithospheric mantle compositions (> 10% depletion) could also

explain the lack of melt at the West Iberia Margin wide OCTZ (Figure 1.7b). There

is, however, no strong evidence that such conditions prevailed during the late stages of

rifting in most of the studied magma-poor passive continental margins.

Alternatively, Latin & White [1990] have shown that if a simple shear model of

rifting is assumed, then the volume of melt generated from adiabatic upwelling of the

asthenosphere during extension is significantly lower, even for very large extension factors

( > 10; Figure 1.7c).

1.2 Isostasy, Flexure and Gravity Modelling

1.2.1 Introduction to Isostasy

In the rifting model put forward by McKenzie [1978], crustal thinning and passive up-

welling of higher density asthenosphere is locally compensated by the subsidence of the

rift basin, which is infilled by lower density air, water and/or sediments. During post-

rift thermal relaxation, the lithosphere thickens, and further subsidence is produced (see

Figure 1.3). The model assumes that isostatic compensation is preserved throughout and

accounts for the variations in the density of the rocks associated with the temperature

effects, i.e. thermal isostasy.

The term compensation, in relation to the way masses are distributed near the Earths

surface was first utilized by the astronomer and mathematician R. G. Boscovich, in the


Figure 1.7: The No Magma Paradox in passive continental margins: Predictions from

thermal numerical models. a) Predicted melt thicknesses assuming depth uniform stretch-

ing. Results are shown in function of mantle potential temperature and rift duration

(indicated in the curves). For a rift duration of 5-15 m.y., mantle potential tempera-

tures < 1250C would be needed to justify the estimated melt thicknesses (< 2 km) in

the Iberia Abyssal Plain (IAP). Modified from Minshull et al. [2001]. b) Results from

numerical models which also take into account the heat losses and gains associated with

melting and serpentinization, respectively (Perez-Gussinye et al. [2006]). According to

these authors, a slow rifting velocity (red and green lines) and/or high mantle depletions

(> 10%; dotted line) could explain the low melt productions at the OCTZ. c) Predictions

from simple thermal models showing the effectiveness of a simple shear model of rifting

(solid black lines; labels give the amount of stretching, ) in reducing the amount of melt

predicted from a pure shear model (dashed grey lines). Mantle potential temperature is

1280C. The blue lines between the liquidus and solidus show the melt fraction by weight

(McKenzie & Bickle [1988]). After Latin & White [1990].

mid eighteenth century. His remarks concerned the observations of Pierre Bouguer a few

years earlier, who suggested that the gravitational attraction of the Andes was much

smaller than expected from the mass of matter represented in the mountains. According

to Boscovich, the mountains were supported by the ... thermal expansion of the material

in depth ... .

The first models of isostasy, however, were only formally conceptualised approxi-

mately one century later (Pratt [1855]; Airy [1855]), to explain a difference of more than

5 between the positions determined using geodetic and astronomical methods in north-

ern India (i.e. just south of the Himalayas mountain range). According to the model

of Pratt, the differences in height of the mountains were compensated by lateral varia-

tions in the average density of the rocks, in a way that the lithostatic pressure would

be maintained at a certain reference level (i.e. the compensation depth; Dc in Figure


2.67

r3.3

2 km 4 km 6 km

h

3 km -5 km

30 km

25 km

Dc

2.67g/cm3

2.62 2.57 2.52 2.59 2.67 2.81

2 km 4 km 6 km

h

3 km -5 km

100 km

Dc

b)a)

Figure 1.8: Pratt (a) and Airy (b) models of isostasy. In Pratts model, topography

(h) is supported by density contrasts above a compensation depth (Dc). Airys model of

isostasy involves a lower density root (r) underneath an excess of mass due to topography.

1.8a). Pratts calculations, however, predicted discrepancies between the geodetic and

astronomical methods approximately three times greater than observed.

Airy, on the other hand, believed that the outermost rigid layer of the Earth overlies

a fluid of greater density, and that the excess in mass caused by the mountains was

compensated in depth by an equivalent mass deficiency, or a root, as represented in

Figure 1.8b. A concise historical record on the developments of the concept and models

of isostasy is given in Watts [2001b].

1.2.2 Flexure and Te

The Pratt and Airy models of isostasy, although based on contrasting assumptions, are

similar in that they assume that an excess in mass due to topography is locally com-

pensated. This view was later contested in a series of papers by J. Barrell in 1914 (see

Watts [2001b] for summary) who, based on relief and gravity considerations, argued that

the Earths outermost solid crust, or the lithosphere as he referred to it, was capable of

rigidly supporting loads associated with prominent geological features, such as mountain

ranges and river deltas. Barrell also predicted the existence of an underlying week, vis-

cous layer, which accommodates the bending of the lithosphere by lateral flow. He called

this layer the asthenosphere.

Although the concept of regional, rather then local compensation, had already been

suggested in previous works (e.g. Gilbert [1889]), it was Vening Meinesz [1931] who first

proposed a model to quantitatively account for the effects of regional isostasy, where the

radius of compensation, R, is proportional to the strength of the crust (Figure 1.9a).

Throughout the late 30s and the 40s, R. Gunn published a series of papers where he


Figure 1.9: Local versus regional (i.e. flexural) models of isostasy. a) Schematic diagram

showing a comparison between the local and regional models of isostasy. R represents

the radius of regional compensation according to Vening Meinesz [1931] (modified from

Watts [2001b]). b) Calculated versus observed flexural deformation under the Island

of Boa Vista (Cape Verde) for different values of the plates flexural rigidity, Te. The

observed flexural deformation corresponds to the moat infill. From Ali et al. [2003].

studied the role of lithospheric strength in explaining the departures from local models of

isostasy (see Watts [2001b] for summary). Importantly, he showed that the major grav-

ity anomalies associated with different types of geological loads (e.g. mountain ranges,

seamounts, passive margins and island arcs-trenches systems) could be explained assum-

ing that the lithosphere behaves as a thin elastic plate overlying a weak, inviscid fluid.

According to Gunn [1943],

Dd4w

dx4+ (m infill)gw = 0 (1.1)

where D is the plates flexural rigidity, m is the density of the underlying fluid, infill

is the density of the material infilling the deflection (w), and g is the acceleration due to

gravity.

The flexural rigidity of the plate can then be expressed in terms of an equivalent

elastic thickness, Te,

D =ET 3e

12(1 2)(1.2)

where E and are the Youngs modulus and the Poissons ratio, respectively, which

reflect intrinsic mechanical properties of the materials.

As shown in Figure 1.9b, by varying the Te associated with the emplacement of a load,

both the amplitude and wavelength of the flexural deformation will change considerably.

The resulting gravity anomaly (i.e. load+flexure) will thus be a consequence of Te.


Figure 1.10: Te versus age

of the oceanic lithosphere at

the time loading. Although

Te estimates from distinct tec-

tonic environments are com-

pared, the values appear to

cluster between the depths of

the 300 and 600C isotherms,

as predicted by Watts [1978].

N is the number of plotted es-

timates. After Watts [2001b].

A great number of studies have since been published with estimates of Te in dis-

tinct tectonic environments, such as seamounts (Watts [1978]; Calmant [1987]), foreland

basins (Karner & Watts [1983]), passive continental margins and intracontinental rift

basins (Stewart [1998] and references therein), trenches (McNutt [1984]) and cratonic

terranes Forsyth [1985]. These estimates are based either on purely elastic models (i.e.

Equation 1.1), or on more or less complex viscoelastic plate models, which account for

the rheological stratification of the lithosphere (e.g. McNutt [1984]; McNutt et al. [1988];

Watts & Zhong [2000]). The general equation for the flexure of a viscoelastic plate

overlying an inviscid substrate is given by (e.g. Nadai [1963]),

Dd4w

dx4+ (m infill)g(w + w) = 0 (1.3)

where w is the first derivative of w with respect to time and is the Maxwell relaxation

time; i.e. when the accumulated elastic strain equals the viscous strain.

In the oceans, a strong correlation has been found between the age of the lithosphere

at the time of loading and the Tes calculated using a simple elastic plate model. This was

first noticed by Watts [1978], using cross-spectral modelling techniques to analyse the

relationship between gravity and bathymetry along the Hawaiian-Emperor Seamount

Chain. According to Watts [1978], the Te of the Pacific Plate follows a predictable

pattern, and corresponds to the depth of the 450C 150C isotherm. As shown in

Figure 1.10, the predictions of Watts [1978] have been confirmed over both recent and

old oceanic lithosphere, and in distinct tectonic settings, throughout the worlds ocean

basins (see Watts [2001b] for references).


Watts et al. [1982] compared the stratigraphy and gravity anomalies observed at a

number of intracratonic rift basins to the predictions from both elastic and viscoelastic

plate models. The modelling results suggest that, overall, an elastic plate model where

the Te increases with age since rifting explains better the observations. As depicted in

Figure 1.11a, this model is characterized by a widening of the basin depocentre with time

and the onlap of the younger sediment units onto the basin flanks (the steers head basin

model; Watts et al. [1982]). In contrast, the viscoelastic plate model predicts that the

younger sediments are confined to the basin centre, due to the increasing stress relaxation

with the age of the load, and show more subsidence than older units (Figure 1.11c).

Figure 1.11: Predicted

and observed stratigra-

phy at intracontinental

rift settings: elastic

versus viscoelastic plate

models (from Watts

et al. [1982]). a) Elastic

plate model. b) Cross

section from the North

Sea Central Graben. c)

Viscoelastic plate model.

The example shown in Figure 1.11b is a cross section from the North Sea Basin (see

Watts et al. [1982] for references), where the maximum subsidence

chapter1 final

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